Salty Matters

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

John Warren - Sunday, June 16, 2019

Introduction

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).

    References

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

    John Warren - Saturday, May 18, 2019

    Introduction

    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."

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    Evaporite interactions with magma Part 3 of 3: On-site evaporite and major extinction events?

    John Warren - Saturday, April 13, 2019

     

    Introduction

    The previous two articles in this series dealt with heating evaporites, volatiles expelled into the atmosphere, and major biotal extinction events. I argued that short-term heating of a megaevaporite mass during emplacement of a Large Igneous Province (LIP) or heating of evaporities at the site of a large bolide impact, will move vast volumes of sulphurous and halocarbon volatiles, as well as solids, CO2 and CH4 into the earth's upper atmosphere (Figure 1a). The resulting catastrophic climatic effects link in time and probable causes to earth-scale major extinction horizons. (Figure 1b). In this article shall examine how three of the five major Phanerozoic extinction events have an evaporite association, starting with the most intense extinction event of the Phanerozoic; the end-Permian and its link to LIP emplacement into two separate sequences of massive bedded evaporite (Cambrian or Devonian mega-salts) in the Tunguska Basin, Siberia.


    End-Permian - Saline interactions during emplacement of Siberian Traps

    The Siberian Traps LIP is of significant size (~7 × 106 km2) and total volume (~4 × 106 km3) (Ivanov et al., 2013 and references therein). It is, however, smaller than the Late Cretaceous Deccan Traps and has a volume that is about a half of the Late Triassic Central Atlantic Magmatic Province (CAMP). All three of these continental LIPs are dwarfed by the Early Cretaceous marine Ontong-Java LIP (≈20 × 106 km3). So, it seems that the volume of igneous material in a LIP does not directly relate to the intensity of the extinction event (Figure 1b).

    The Siberian Traps include ultramafic alkaline, mafic and felsic rocks that erupted in different proportions within a vast region extending over several thousands of square kilometres across Western and Eastern Siberia (Figure 2a). The Siberian Traps are considered have been emplaced atop a hotspot in a relatively short time frame (≈1 million years), when a large volume of deep mantle-derived igneous material was intruded and erupted at the Permo-Triassic boundary (Burgess et al., 2017).


    Trap geology

    Near Noril'sk, lava outflows reach thicknesses of over 3 km, while further to the northeast in the Maymecha-Kotuy region, half of the total lava pile is composed of ultramafic rocks including magnesian rich meimechites (Figure 2a). The very high MgO contents (8-40 wt %) of the meimechites in such low-degree melts indicates that the site of initial melting was very deep, as much as 200 km, and either in the lowermost continental lithosphere or in the underlying asthenosphere (Arndt et al., 1995). Melting probably was linked with the arrival of a mantle plume that was in its turn the source of the Siberian basaltic flood volcanism.

    Thickness of volcaniclastic material in the Siberian Traps ranges from intercalated layers less than a meter thick on the Putorana Plateau to hundreds of meters near the base of the volcanic sections in the Angara and the Maymecha-Kotuy areas (Figure 2a). The total volume of mafic volcaniclastic material has been estimated at >200,000 km3 or >5% of the total volume of the Siberian Traps (Black et al., 2015). Volcanic rocks of this age are also present in drillcore in the West Siberian Basin (Ivanov et al., 2013).

    Magma-sediment and magma-water interactions active during emplacement of the Siberian Traps in the upper lithosphere encompass a variety of heated evaporite interactions: batholith metal-evaporite interactions, lava-water interactions and intense phreatomagmatic explosions via vents and breccia pipes that formed saline-igneous volatile fountains reaching the upper atmosphere. The positions of these fountains are perhaps indicated by vent-related iron-rich diatremes (Figure 2a; Svensen et al., 2009). All these interactions are critical inputs to the End-Permian extinction event that links vast volumes of altered evaporites with the heating mechanisms inherent to Siberian Trap geology.

     

    Evaporite basins (Devonian and Cambrian)

    The Siberian Traps region is not only significant because of its vast extent and its deep nickel-prone mantle source, but also in that the immense volumes of igneous rocks that making up the traps were emplaced into two chemically prone saline giants with differing dominant mineralogies and ages; 1) Cambrian mega-halite sediments in the south, with interlayers of hydrated potash salts (mostly carnallitite) and 2) Devonian megasulphates in the north, containing two 50-100m beds of anhydrite (Figures 2b, 5). The interactions with the two types of salt basins, one halite-dominant, the other anhydrite-dominant, gives rise to two distinct meta-evaporite indicator associations. In the North, the interaction of picritic magmas with bedded thick anhydrites formed the supergiant Noril'sk nickel deposit, while in the south the LIP emplacement formed numerous magnetite-rich explosive breccia pipes, sourced at the stratigraphic level of the Cambrian salts (Figure 2b).


    Norils'k region & Devonian evaporites

    In the northern part of the Tunguska Basin the evaporite sediments hosting the intrusives of the Siberian Traps are a combination of Devonian anhydrites and carbonates, with overlying Carboniferous coals. Trap basalts, now cover this sedimentary sequence (Figure 4a), while sill-like tholeiitic intrusions, varying in composition from subalkaline dolerite to gabbro-dolerite are emplaced in the sediment pile and were part of the feeder system to the flood basalts (Figures 4b, 5, 6).


    The region of Devonian evaporites contains the Noril'sk-Talnakh ore deposit, the largest Phanerozoic nickel deposit in the world (Figures 3, 4; Naldrett 2004). In the mine area, ore-bearing gabbroic-dolerites are differentiated, whereby picrite and picritic dolerite are overlain by more felsic differentiates. The Cu-Ni-platinoid mineralisation at Noril'sk forms relatively persistent stratabound horizons of massive sulphides in the lower portions of the three mineralised intrusions (Noril'sk, Talnakh, Kharaelakh), which are made up of segregations and accumulations of pyrrhotite, pentlandite and chalcopyrite (Figures 5, 6).


    At the world-scale, the supergiant Permian Noril'sk-Talnakh deposit is an unusual Cu-Ni deposit. It did not form in the Precambrian, and so is unlike almost all the world's other supergiant magmatic nickel-sulphide deposits (Figure 3). It formed at the end of the Palaeozoic and straddles the Permo-Triassic boundary (Black et al., 2014a). Magmatic nickel ores at Noril'sk crystallised outside the influence of the reducing planetary atmosphere that typifies Archaean Ni flood basalt deposits and is not tied to greenstone terranes and the athenospheric transition to more sialic plate-scale conditions. (Figure 3). The high temperatures and near complete assimilation of Devonian sulphate evaporite blocks within the Noril’sk magma mean that this is one of the more enigmatic (“salt is elsewhere”) styles of evaporite-related high-temperature ore deposits (Warren, 2016, Chapter 16). Notions of evaporite assimilation for ore deposits tied to igneous-evaporite interactions are usually only one of multiple possible explanations of a magmatic ore but, in my opinion, for Noril’sk this is the most likely scenario. So, I emphasise the evaporite connection for the Noril’sk-Talnakh deposit in this article. Alternate non-evaporitic orthomagmatic explanations can be found in papers such as Wooden et al., (1992); Lightfoot et al. (1997), and Krivolutskaya (2016). Independent of the mode of nickel-ore fixation, most authors working in the Tunguska Basin agree that the emplacement of the trap intrusives drove the escape of a huge pulse of sediment-derived volatiles into the Earth's atmosphere.


    Regional structure of the Noril’sk district is dominated by NNE-NE Permo-Triassic block faulting, which was coeval with magmatic activity. Individual faults may be over 500 km in length with throws of up to a kilometre (Figure 4b; Naldrett, 1997). Mineralised intrusions radiate outward and upward from intrusive centres and penetrate all levels of the overlying sedimentary sequence. Most intrusive centres are associated with prominent block faulting and fault intersections. The main Noril’sk-Kharaelakh fault occurs within the Siberian Platform, but is parallel to the main fault system that defines the boundary between the platform and the nearby Yenisei Trough. The Kharaelakh-Noril’sk fault guided the main upwelling magma body (Figures 4b, 6). Individual sills splay off this fault control and are interlayered with sulphate evaporite beds to can attain lateral lengths of 12 km, widths of 2 km and thicknesses of 30 to 350 m.

    Mineralogical compositions of the Devonian sediments interlayered with these sills are of great importance in understanding the geological responses to heating by intrusive igneous sills in the Noril'sk-Talnakh area (Figures 5, 6). Based on their lithological features and paleontological character, the intruded Devonian succession is subdivided into the Yampakhtinsky, Khrebtovsky, Zubovsky, Kureysky, and Razvedochninsky Formations (Lower Devonian), the Manturovsky and Yuktinsky Formations (Middle Devonian), and the Nakohozsky, Kalargonsky and Fokinsky Formations (Upper Devonian) (Figure 5; Krivolutskaya, 2016; Naldrett, 2004). The two main evaporite levels are the Middle Devonian and Lower Devonian anhydrite-dominant successions, both deposited in a subsealevel transitioning rift (Figure 5, 6; Naldrett, 2005; Warren, 2016).

    The Yampaktinsky and Khrebtovsky Formations consist of Lower Devonian carbonates interbedded with abundant gypsum (in outcrop) and anhydrite (subsurface), along with some of the oldest lenses of celestite in the area (Figure 5). The total thicknesses of these two CaSO4 units are around 100 and 80 m, respectively. The Lower Devonian Zubovsky Formation is composed of grey-colored dolomitic marls interbedded with argillaceous dolomites, mudstones, and anhydrite with a total thickness of 110–140 m. The Zubovsky Formation unconformably overlies the Lower Devonian Khrebtovsky Formation in the Noril’sk region. The Lower Devonian Kureysky Formation consists of mottled dolomite and calcareous mudstones and marls with rare siltstone and limestone. The thicknesses of all units in the outcrop section remain stratiform and vary within 50–60 m. The contacts with the overlying and underlying formations are conformable.

    The Lower Devonian Razvedochninsky Formation is dominated by siltstones, sandstones, and conglomeratic sandstones with a thickness that regionally does not exceed 110–150 m, but reaches 150–235 m in troughs, and decreases sharply to the south until fully wedging out.

    The Middle Devonian Manturovsky Formation overlies the eroded Razvedochninsky Formation and consists of a terrigenous-carbonate section with abundant salt-bearing strata, most of which consist of rock salt or brecciated equivalents. This formation’s thickness is 100-210 m but ranges up to 500 m (Figure 6). The Middle Devonian Yuktinsky section is dominated by clastic–carbonate sediments ranging from 12 to 40 m thick, while in the troughs the thickness of interlayered sulphate rocks reaches 55 m. The contacts with the underlying and overlying Middle Devonian Manturovsky deposits are considered comformable. The Upper Devonian Nakokhozsky Formation consists of folded calcium-sulphate-rich variegated shale–carbonate rocks with a thickness of 2–60 m that increases in the troughs to 80–130 m (Figure 5). The Upper Devonian Kalargonsky Formation is characterised by a grey-colored terrigenous-carbonate section that includes dolomites, dolomitic marl, dolomite–limestone, and anhydrite dominate in the basins. This formation’s thickness is 170–270 m. The Kalargonsky Formation unconformably overlies the Middle Devonian Nakokhozsky sediments and the contact is typically a breccia (Figure 5).

    The Middle Devonian Fokinsky Formation (as distinct from the mineralised Fokinsky intrusions) consists of evaporite sulphate-rich clastic–carbonate sequences, primarily within the troughs, and anhydrite, dolomitic marls interbedded with limestone lenses of rock salt, and clay–carbonate breccias (Krivolutskaya, 2016). The thickness of this formation is 220–420 m (approximately 500 m in the western part of the Vologochansky Trough).

    The Fokinsky Formation is not recognised by all authors working in the region. This disparity in stratigraphic recognition across the region underlines a problem inherent in the litho-stratigraphic descriptions of many bedded evaporite regions worldwide, where it is assumed that a layer-cake stratigraphy/correlation is present pre- and post-intrusion. Thereby the effects of evaporite collapse dissolution, bed wedge-out and possible salt flow are not quantified. In my opinion, sedimentary breccias in such regions are more likely to be diagenetic and laterally discontinuous (see Warren, 2016; Chapter 7).

    In summary, the Devonian stratigraphy in the vicinity of the Noril'sk Mine retains significant thicknesses (50-100m) with variations centred on transitions in and out of bedded anhydrite. There is a strong likelihood that the current outcrop geology interpretations under-illustrate former thicknesses of bedded evaporites during to ongoing dissolution, collapse and possible flowage.

    The anomalous Phanerozoic age of the Noril’sk-Talnakh ore deposits, compared with the Precambrian ages of other magmatic Ni-Cu deposits, and its relative enrichment in Ni, Cu, Pt and Pd compared with Sudbury and Jinchuan (Figure 3), is thought to reflect the anomalously high volumes of sulphur in the parent magma. Additional sulphur entered the evolving magma chamber via intrusion and assimilation of CaSO4 blocks and associated hydrothermal solutions altering and dissolving adjacent thick-bedded anhydrite successions (Figure 7; Naldrett 1981, 1993, 1997; Pang et al., 2013). Noril'sk-Talnakh's rich sulphur supply contrasts with that of the komatiitic Archaean Cu-Ni deposits, where the sedimentary sulphur supply came from more ubiquitous, less-focused sulphur sources sometimes entrained in widespread sedimentary pyrite (Figure 3). Such pyrite characterises a significant portion of fine-grained sediments accumulated under an anoxic reducing Archaean to Palaeoproterozoic atmosphere.


    Abundant crystals of magmatic anhydrite today typify the olivine-bearing (picritic) gabbros in the Kharaelakh intrusion, which is located in the basin stratigraphy at the level of the Devonian anhydrites (Figure 6; Li et al., 2009 Spiridonov, 2010). Along with disseminated sulphides, the anhydrite crystals are characterised by planar boundaries with co-associated olivine and augite. Dihedral angles of ~120°, characteristic of simultaneous crystallisation, are common throughout the anhydrite-augite assemblages. Inclusions of anhydrite in augite and vice-versa are also typical.

    Rounded and subrounded sulphide inclusions composed of pyrrhotite, pentlandite, and chalcopyrite, that crystallised from immiscible sulphide liquid droplets in the magma, are commonplace within the magmatic anhydrite crystals and in the contact aureoles (Figure 7). Visual estimates by Li et al. (2009), based on five polished thin sections, indicate that the ratio of anhydrite to sulphide in mineralised samples varies from 0.05 to 0. The observation of abundant wollastonite in contact aureole rocks at this stratigraphic level suggests that reactions such as CaSO4 + SiO2 + H2O = CaSiO3 + H2S + 2O2 occurred, and that sulphate was likely reduced to sulphide before incorporation into the magma (Ripley et al., 2007).

    Picritic magmas in mantle plumes can have melt temperatures as high as 1600°C (Hezberg et al., 2007). Assimilation of anhydrite via partial melting of a cooler basaltic magma at shallower depths can be more difficult, owing to the high melting point of pure anhydrite (melt temperatures typically rang between 1360 and 1450°C, although this is significantly lowered in the presence of organics and water). Rather than only melting anhydrite enclosed by picritic magma, additional fluxing mechanisms likely move additional anhydrite-derived sulphur into the melt, either by hydrothermal leaching of sulphate followed by partial reduction, or via a process involving the dissolution of anhydrite during thermochemical sulphate reduction (TSR; Warren, 2016; Chapter 9). The latter process requires heat, anhydrite and organics (generally in the form of hydrocarbons or kerogen).

    Some authors use the euhedral outline of anhydrite in mineralised sills, as seen in Figure 7, to argue blocks anhydrite country rock was not assimilated. This is a specious argument as this type of anhydrite was precipitated during cooling of an already sulphur-saturated magma, the euhedral spary outline does not relate to the source of the sulphur, which is more clearly indicated by its sulphur-isotope signature (Figure 8a - also Warren, 2016; Chapter 8).


    Isotopic analysis of δ34S in the magmatic anhydrites and associated metal sulphides in the Kharaelakh intrusives require the assimilation of externally-derived high-δ34S sulphur from the adjacent country rock (Figure 8: Ripley et al., 2007, 2010). Where complete sulphate reduction occurred, the δ34S values require mixtures of some 60% anhydrite-derived evaporitic marine sulphur (δ34S values near 20‰), with 40% mantle-derived sulphide (δ34S of 0‰) to produce the required measured magmatic sulphide values ≈12‰ (Figures 8a, b).

    The sulphur isotope data and the nature of the sampled contact aureoles suggest intense intracontinental rifting in the Noril’sk region brought deeply-sourced mafic magmas into contact with supracrustal sulphur from evaporitic sulphates at the level of the Kharaelakh intrusion. Sulphur isotope data show the mineralised intervals at Noril’sk are anomalously heavy in δ34S (Figure 8a, b). These data are inconsistent with sulphur derived from mixing of the mantle magma sulphur (δ34Svalues near zero) with sulphur from an evaporitic sulphate source (Godlevski and Grinenko, 1963; Grinenko, 1985; Li et al., 2009; Pang et al., 2012; Black et al., 2014a).

    Sulphur isotope values from Paleozoic evaporites vary between +10 and +35‰ (Figure 8b; Claypool et al., 1980). Cambrian evaporites, including the major Irkutsk basin salts in Siberia, are the most 34S-enriched evaporites in the Phanerozoic, with mean δ34SVCDT = +30‰ (Claypool et al., 1980; Black et al., 2014a). Two-member mixing curves between meimechite and anhydrite sulphur (with δ34S = +20 to +35‰) convincingly reproduce the observed δ34S trends for the Noril'sk ores (Figure 8b; Black et al., 2014a).


    As the magma rose through the sedimentary cover, it penetrated and assimilated sulphur from extensive Devonian anhydrite layers (Figure 9). Sulphur in calcium sulphate was reduced to sulphide, CaO entered the magma, and iron from the magma reacted with reduced sulphur so that the end result was droplets of immiscible iron sulphide dispersed through the melt (Naldrett and Macdonald, 1980). These droplets acted as collectors for Ni, Cu and the platinum group elements, which are now so enriched in the Noril’sk ores.

    Naldrett (1991, 1997, 2005) concluded that prehnite + biotite + anhydrite + carbonate + zeolite + chlorite ± sulphide globules, which typify chromite agglomerations in the picrite of the Noril’sk intrusions, represent remnants of partially assimilated sulphate-rich country rock. Assimilation of anhydrite-rich rocks, coupled with the reduction of sulphate to sulphide, would have introduced considerable oxygen into the silicate melt, which then drove precipitation of chrome-spinel minerals (chromite - FeCr2O4; mangnesiochromite - MgCr2O4). Inclusions of anhydrite-rich material, floating in the magma, would have served as loci for chromite crystallisation, thus giving rise to the association between the agglomerations and the globules. Tarasov (1970) pointed out that Middle Carboniferous coal measures were also assimilated and may have supplied organics that assisted in the reduction of sulphur in the magmas (Figures 6, 7).

    Evidence of the assimilation of large volumes of anhydrite and coaly organics into the magma mass has implications beyond the formation of the Noril’sk-Talnakh ore deposits. Li et at. (2009) identified magmatic anhydrite-sulphide assemblages in a subvolcanic intrusion associated with the Siberian Traps. The δ34S values of anhydrite and coexisting sulphide crystals analysed by ion probing are 18‰–22‰ and 9‰–11‰, respectively, are much higher than the anhydrite-contaminated ore values shown in Figure 8). To obtain this level of fractionation means more than 50% of the total sulphur in the intrusion was derived from marine evaporites in the footwall strata. The contaminated magma was highly oxidised and able to dissolve up to one order of magnitude more sulphur than pure mantle-derived basaltic magma. Such sulphur-contaminated magma, when erupted, would have released vast volumes of SO2 into the atmosphere (Black et al., 2012, 2014b). That is, the eruption of the anhydrite-contaminated magma that is the Siberian Traps in the northern Tunguska Basin can help explain the intensity of the end-Permian extinction.

    In summary, such igneous - sulphate sediment interaction explains, at least in part: (1) the vast amount of sulphide melt in the Noril’sk-Talnakh ore field; (2) the heavy quasi-anhydrite isotopic composition of sulphur in sideronitic and massive nickel ores; (3) the reduced contents of noble metals in these ores (compared with the drop sulphides that occur toward base of the intrusions and have a likely mantle sulphide source); and (4) the high contents of radiogenic (crustal) osmium in sideronitic and massive ores (Spiridonov, 2010; Walker et al., 1994). In summary, the reserves of the world-class Ni-PGE deposit at Noril’sk-Talnakh, with its anomalous Phanerozoic age, likely reflect a fortuitous occurrence of thick Devoninn anhydrites (ultimate sulphide source) atop an active later set of deep mantle-tapping rift grabens that drove the LIP outlined by the Siberian Traps. Wherever these magmas vented into the Earth's atmosphere they carried significant volumes of sulphurous volatiles.


    Cambrian evaporites, potash & breccia pipes

    Salt deposits of late Vendian to Early Cambrian age in East Siberia cover an extensive area (ca. 2 million km2) located to the north-west of Lake Baikal with an extent showing it extends across much of the Permian Siberian Traps (Figure 2b). The thickness of this upper Vendian-Lower Cambrian evaporite succession is 2.0–2.5 km in the southern, western, and central parts of the basin, and 1.3–1.5 km in the NE part (Nepa-Vilyui). This saline giant (total volume of upper Vendian–Lower Cambrian evaporites is 785,000 km3; Zharkov, 1984) is characterised by the occurrence of fourteen regional marker carbonate units and 15 salt units (Figure 10; Zharkov, 1984, with references therein). Five major phases of salt deposition are distinguished, namely the late Vendian (Danilovo) and Early Cambrian (Usolye, Belsk, Angara, and Litvintsevo) salt basins (Figure 10a; Zharkov (1984), Kuznetsov and Suchy. (1992).

    Average thicknesses of the Cambrian evaporite deposits decreases with time (Figure 10b) as does the area (Figure 10a). The area of the oldest Cambrian basin, the Usolye salt basin is almost 2 million km2, and the average thickness of deposited salt around 200 m (Zharkov, 1984), while the area of the youngest, Litvintsevo salt basin is 0.5 million km2 and the average thickness of its evaporite bed (rock salt and anhydrite) is 50 m (Figure 10; Zharkov, 1984).

    Most of the petroleum reservoirs in the region are located in the Cambrian carbonates. The post-Cambrian stratigraphy contains major erosional breaks. As we saw in the Noril'sk discussion, Devonian evaporites are rare in the south but abundant in the north, whereas Ordovician rocks (limestones, marls) are locally abundant in the central parts of the basin. Cambrian salt deposition is interpreted as mostly taking place in a deeper water basin: Petrichenko (1988) concluded that at the termination of halite deposition the final brine depth was 50–260 m, and at the onset of potash deposition it was ≈10–50 m.


    Lower Cambrian Angara evaporites host the largest known bedded potash deposit in Russia, which is not yet produced (Figure 11; Garrett, 1995; Warren 2016, Chapter 11). Potash salts occur at the base of the Angara Formation in what is called the sixth halite series (Table 1). This intracratonic potash basin is one of the larger potash-entraining salt sumps in the world, it is several times larger than the Permian Upper Kama deposit and approaches the Prairie Evaporite in aerial extent, but not in lateral continuity, thickness or purity (Figure 11) due in large part to the effects of igneous disburbance.

    Plans were made in 1986 under the old Soviet regime to initiate a mining program in a section of this basin called the Nepskoye deposit but were never fully implemented, although some ore was extracted in the mid 1980s (Andreev et al., 1986). The proposed potash development region is located near the towns of Nepa and Ust-Kut (300 km apart) in Irkutsk State. Regionally, the dominant potash mineral is carnallite, but high-grade sylvinite is intersected at depths of 600-1,000 m in beds some 1.5-5 m thick over an area ≈ 1,000 km2 (Garrett, 1995). The lower Bur or K1 bedded potash horizon lies at a depth of 750 - 960 m and is 2-18 m thick (4-6 m in the central area (Figure 11a; Table 1). Two sylvinite zones in this horizon were mapped, with the central one being 16-26 km long and 6-8 km wide (Figure 11a). In the lower horizon (K1) the sylvinite was 1.5-3 m thick, and averaged 15-50% KCl, 0.05-0.5% MgCl2,with 0.5% insolubles. The overlying K2 potash zone (Tunguaka) also entrains several sylvinite beds and is some 679-880 m deep and 2.5-20 m thick. It has a 15-45% KCl content and comparatively low MgCl2 and insoluble contents. This zone represents the major potash reserves of the deposit. In the upper potash beds (K2) the sylvinite strata become more discontinuous, but some reasonably thick, high grade and extensive zones exist (Andreev et al. 1986). The sylvite ore sits in a more regional potash succession composed of a combination of carnallitite and sylvinite (Figure 11b). The broader Nepa potash region as generally mapped in Figure 2 has two interesting characteristics; 1) The igneous trap rocks as defined in the drill-controlled cross sections of Malykh and Geletii (1988) sit below the potash level (Figure 11b), 2) There is a paucity of magnetitic explosive breccia pipes in the Nepa potash region (Figure 2b).

    To the south and west, between Irkutsk and Taseyevo some 400 km to the west, other large potash occurrences have been reported in the same general but poorly delineated evaporite basins. For instance, in the Kanak-Taseyevo basin, potash beds (sylvite-carnallite containing 3-24% K2O) have been intersected at depths of 1,240-1,415 m (Garrett, 1995). Potash beds at these depths would require a solution mining methodology, but the at-surface climate would mean either cryogenic pan processing or evaporators, making recovery more difficult and expensive (Warren, 2016; Chapter 11).

    Basaltic breccia pipes, Tunguska Basin Siberia

    Basalt pipes form a rim to the main basalt body of the Siberian Traps and are genetically linked to trap emplacement (Figure 2b; Polozov et al., 2016). The pipes pierce through all sedimentary strata, even dolerite sills higher in the Permo-Carboniferous portion of the basin stratigraphy, and are considered to be a type of diatreme. Importantly, the basalt pipes with magnetite cores tend to occur across the southern Tunguska Basin, while unmineralised basalt pipes are more widespread (Figure 2b). Some of the basalt pipes bearing magnetite mineralisation are of commercial grade and are mined for their iron ore.

    Regionally, it is difficult to estimate the total number of pipes (both “barren” basalt and magnetite-enriched) because repeated glaciations have flattened relief, while thick taiga forest covers significant parts of Siberia. Thus, many pipes are hidden by swampy coniferous forests and so are difficult to map. However, conservative estimates based on prospecting surveys for iron mineralization in the southeern portion of Tunguska Basin, and geological mapping elsewhere, suggest there are more than three hundred magnetite-bearing basalt pipes. This includes 6 large (>100 Mt of iron ores), 14 medium (20–100 Mt) and 19 small <20 Mt) sized iron deposits. All other mineralised basalt pipes are currently of sub-economic grade or underexplored (Polozov et al., 2016).

    The magnetite deposits are consistently located in the Tunguska Basin region underlain by Cambrian evaporites and mainly defined by subvertical and cylindrical breccia bodies with magnesio-ferrite and magnetite as the primary ore minerals (Figure 2b). In many ways these deposits are similar to iron oxide, gold and copper (IOGC) deposits worldwide, but are classified in the Russian literature as Angara–Ilim type deposits, named after the two rivers where a large number of iron- mineralised basalt pipes crop out (Soloviev, 2010; Warren 2016, Chapter 16).

    Korshunovsky (Korshunovskoye) region, Siberia

    This region, in the Irkutsk district, is the eighth largest iron ore producer in Russia, with an annual output of 5 Mt of iron ore concentrate. Across the region, the pipes are sourced in the Cambrian evaporite part of the basin stratigraphy and pierce younger Paleozoic sediments composed of argillites, limestones, marls, siltstones, sandstones and clays of the late Cambrian Lena, Ust'kut, Mamyr and Ordovician Bratsk groups and overlying Early Carboniferous limestone.


    We shall focus one of the largest magnesite deposits in the region, the Korshunovshoe (Korshunovsky) magnetite breccia pipe, with an estimated reserve of 1.5 Gt of ore to a depth of 1700 m (Soloviev, 2010; Polozov et al., 2016). It is mined (open pit) and so the interior structures and relationships are well documented (Figure 12). The currently mined pipe is adjacent to another explosion pipe to the immediate south-east, with the mineralised breccias sourced mainly at the level of the Cambrian evaporites (halite, potash and anhydrite; Mazurov et al., 2007). At outcrop and in the pit little evidence, other than secondary textures (dissolution-collapse and brecciation), remains of the primary minerals of the mother saline layer, although remnant, recrystallised evaporite clasts (including halite and anhydrite) typify the mineralised breccia in the lower parts of the pipe (Mazurov et al., 2007, 2018). Textures at the evaporite level in the diatremes are not unlike those seen in regions of Eocene sill interaction with hydrated salts in the Zechstein potash mines of East Germany (Schofield et al., 2014; Warren 2016 and part 1 in this series of Salty Matters articles).

    The Korshunovshoe pipe is filled with tuff breccias and fragmentals composed of the surrounding saline country rocks which have undergone considerable metasomatic alteration. They incorporate fragments and larger blocks of sedimentary (60 to 80 vol.%; sandstones, siltstones, limestones, evaporite residues and argillites) and igneous (10 to 40 vol.%; gabbro-dolerites, dolerites and basalts) rocks, cemented by essentially chloritic material as well as by fine-grained carbonate (Figure 13). The central part of the magnetitic diatreme characterised by intense multiple brecciation, with rock fragments in the breccias represented mostly by variably-altered dolerites. They are cemented by a finely-dispersed matrix, entirely replaced by skarn, post-skarn alteration assemblages and iron oxides.


    Outside of this zone, intense fracturing has occurred, locally with brecciation in altered sedimentary rocks. The fractures are filled with magnetite, accompanied by chlorite and calcite. Finally, the outermost zone is characterised by weak, predominantly sub-horizontal fractures within sedimentary host rocks, locally replaced by skarns. Steeply-dipping dykes of gabbro-dolerite, dolerite, dolerite-porphyry, and basalt-porphyry are present, both within and outside the breccia pipes, while sub-horizontal dolerite sills occur at depth (Soloviev, 2010; Mazurov et al., 2007, 2018).

    Magnetite pipe orebodies at Korshunovshoe are texturally and mineralogically complex (Figures 12, 14) and are composed of: i) Banded masses of metasomatic magnetite that are within, and conformable to saline to calcareous members of the host sedimentary wall rocks (dominantly in dolomitic limestones, marls, calcareous argillites and sandstones with a calcareous or limy matrix, but only to a minor degree in sediments without a saline carbonate component) at a depth of some 700 to 1500 m from the surface; ii) Stock-like, lensoid, layered and columnar bodies of magnetite within the altered pyroclastics of the breccia pipe; and iii) Steeply dipping vein-like masses in zones of intense brecciation and replacement by skarns.

    Together these mineralisation styles form two large continuous bodies in the Korshunovshoe pipe (Figure 12). The main deposit has the form of a sub-vertical breccia pipe with plan dimensions of approximately 2400 x 700 m. Mineralisation has been traced by drilling to a depth of 1200 m, and by geophysical data to at least 3 km below the surface (Soloviev, 2010).

    The bulk of the ore is associated with brecciation and occurs within sediments, tuffs and igneous rocks and are demonstrably due to the partial replacement and alteration of the host. Massive and banded ores are less well developed. The mineralisation is mostly magnetite (≈82% of iron resources), with minor magno-magnetite, hematite and martite. The main orebody comprises vertically overlapping zones, with variable amounts of hematite and martite in the upper layers, calcite and magnetite in middle layers, and halite and magnetite in lower layers. The magnetite of the upper to middle zone is accompanied by pyroxene, chlorite and minor epidote with lesser amphibole, serpentine, calcite and garnet, and rare quartz, apatite and sphene and occurs as oolites, druses, masses and disseminations. Calcite increases downwards to 20 to 30%. In the lower part of the deposit, halite, amphibole and Mn-magnetite are more abundant. Pyrite, chalcopyrite and pyrrhotite are found throughout. Much of the magnetite is magno-magnetite which contains up to 6% MgO.

    Across the region of magnetite breccia pipes, ore is extracted from magmatic diatremes that completely penetrated the highly evaporitic lower Phanerozoic succession (Figure 14; Mazurov et al., 2007, Polozov et al., 2016). Early work on this intrusive magnetite style, which surrounds brecciated diatreme-like pipes, classified it as a skarn association, forming a halo around a set of explosive pipes that accompanied regional trap magmatism (Ivashchenko and Korabel’nikova, 1960).

    Characteristic spinel-forsterite magnesian skarns are confined to the overdome parts of large doleritic bodies and are the result of interactions of massive evaporitic and petroliferous dolomites with fluids released from liquid magma (Mazurov et al., 2007). Magnesian skarns of the postmagmatic stage are localised in the marginal parts and on the front (outwedged portions) of doleritic sills, apophyses, and the branches of intrusive bodies hosted at the level of the Cambrian carbonate-evaporite successions (Figure 14). The skarns penetrating the evaporite levels have a banded or layered structure and resemble gravel conglomerates, with carbonate cements. The round fragments (metasomatic pseudo-conglomerates) are composed of globules of disintegrated doleritic porphyrite, completely or partially substituted by zonal magnesian skarns. Their mesostasis is cryptocrystalline, and early phenocrysts of olivine, plagioclase, and pyroxene have undergone dispersion and substitution. Unaltered cores of the metasomatic ‘conglomerate’ are in contact with a fassaite zone, which passes outward into a spinel-fassaite zone and then into a forsterite-magnetite and calciphyre zone.


    The geometry of pipe emplacement is broken down into three related styles; i) Root zone, ii) Diatreme zone, iii) Crater zone (Figure 14). The upper crater zone is sometimes complicated by the presence of reworked crater-lacustrine deposits (Polozov et al., 2016). The root zone is typically brecciated with pseudoconglomerated and other saline volatilisation textures described in the previous paragraphs. The root zone can be traced out from the pipe stem as disturbed zones with considerable lateral extents at the level of the Cambrian evaporite beds. Subhorizontal brecciated dolerite “sills” of the Kapaevsk iron deposit were cemented with calcite, magnetite and halite in various ratios and traced down to deep levels close to the root zones in some basalt pipes In the Korshunovsk iron-ore deposit, such a brecciated body extends from the main diatreme pipe some 5 km to the west and 9 km to the south-west (Von der Flaass and Nikulin, 2000).

    Although not discussed in terms of a volatilisation mechanism in the published literature, I would argue that the lateral apophsyes are indicative of the former presence of hydrated salt layers, probably carnallitite beds showing similar responses to those seen in the potash mines of East Germany (Shofield et al., 2014; or part 1 in this current series of Salty Matters articles).

    The diatreme chimney atop the root zone indicates the rapid rise of a overpressured and upward flowing gas-charged rock mass. Basalt magma served as the ultimate source of iron for the magnetite in the breccia pipes. Extraction of iron from the melt and its transition and accumulation took place in the presence of chlorine-rich fluids, which were formed in the course of thermal decomposition of halite-hosted hydrated salt beds (carnallite). In the later stages of ore formation, some chlorine was fixed in scapolites, while sodium was fixed in albitites and scapolitites (dipyres). In the tuffs of a number of diatremes and paleovolcanoes of the Siberian Platform, native iron can form metal balls in association with moissanites and diamonds (Goryainov et al. 1976). The occurrence of such phases, as well as bitumen in calderas and carbonaceous matter in pisolite tuffs, points to the migration of hydrocarbon fluids through the volcano-tectonic structures (Ryabov et al., 2014).


    Hydrocarbons are abundant in the Cambrian and Ordovician sections of the Tunguska Basin, while coals are widespread in the Permo-Carboniferous Tunguska Series sediments (Figure 15). The juxtaposition of a vast volcanic province with its dykes, sills and diatremes interacting with extensive intracratonic saline Cambrian beds containing evaporites sealing substantial oil accumulations and interacting with coal-bearing deposits, likely produced massive quantities of halocarbons along with methane and CO2. Notably, contact metamorphism with hydrothermal systems rich in chlorine, created during pressure dissolution and dehydration of the surrounding evaporites, potentially synthesized large amounts of the organohalogens methyl chloride (CH3Cl) and methyl bromide (CH3Br) (Beerling et al., 2007; Visscher et al., 2004; Svensen et al., 2018).

    In terms of rapid transfer of volatiles to the atmosphere, the phreatomagmatic-sediment pipes (diatremes) generated tall, explosive volatile-rich eruption columns, which at times reached the stratosphere (Svensen et al., 2009). Such features simultaneously promote removal of highly soluble volcanic gases, such as HCl and SO2, and potentially deliver large volumes of sulphur, halocarbons water, methane and CO2 to the upper atmosphere (Black et al., 2015).

    Timing of trap emplacement

    Siberian Traps magmatic activity at the end-Permain is segmented into three distinct emplacement stages (Figure 16; Burgess et al., 2017). Stage 1, beginning just before 252.24±0.1 Ma, was characterised by initial pyroclastic eruptions followed by lava effusion. During this stage, an estimated two-thirds of the total volume of Siberian Traps lavas were emplaced (>1×106 km3). Stage 2 began at 251.907±0.067 Ma, and was characterised by cessation of extrusion and the onset of widespread sill-complex formation. These sills are exposed over a >1.5 × 106 km2 area and form arguably the most aerially extensive continental sill complex on Earth. Intrusive magmatism continued throughout stage 2 with no apparent hiatus. Stage 2 ended at 251.483±0.088 Ma, when extrusion of lavas resumed after an ~420 ka hiatus, marking the beginning of stage 3. Both extrusive and intrusive magmatism continued during stage 3, which lasted until at least 251.354 ± 0.088 Ma, an age defined by the youngest sill dated in the province. A maximum date for the end of stage 3 is estimated at 250.2 ± 0.3 Ma.


    Integration of LIP stages with the record of mass extinction and carbon cycle at the Permian-Triassic Global Stratotype Section and Point (GSSP) shows three important relationships (Burgess et al., 2017). (1) Extrusive eruption during stage 1 of Siberian LIP magmatism occurs over the ~300 kyr before the onset of mass extinction at 251.941 ± 0.037 Ma. During this interval, the biosphere and the carbon cycle show little evidence of instability. (2) The onset of stage 2, marked by the oldest Siberian Traps sill, and cessation of lava extrusion, coincides with the beginning of mass extinction and the abrupt (2–18 kyr) negative δ13CPDB excursion immediately preceding the extinction event (Figure 16a). The remainder of LIP stage 2, which is characterised by continued sill emplacement, coincides with broadly declining δ13CPDB values following the mass extinction. (3) Stage 3 in the LIP begins at the inflexion point in δ13CPDB composition, after which the carbon reservoir trends positive, toward pre-extinction values.

    Explosive volcanism in the Siberian Traps can be classified in three distinct groups: 1) deep-rooted sediment–magma interactions and pipe eruption where feeder sills are emplaced in evaporites (Cambrian and Devonian country rock), 2) shallower magma-water interactions in areas with abundant groundwater or hydrated salts, and 3) lava flows and lava fountaining during the main stage of effusive volcanism (Jerram et al., 2016a,b). Each stage has a differing set of expressions in terms of the interacting evaporites and the landscape expression of these interactions.

    Outcomes of the end-Permian igneous evaporite interplay

    A unusual aspect of the Siberian trap eruption compared to many but not all LIPs is the saline and kerogen-rich nature of regional geology in the Siberian platform that interacted with the LIP magmas. The main lithologies of the region are large volumes of Devonian anhydrites in the north, Cambrian halite and hydrated-potash salts in the south, hydrocarbon source rocks and evaporite-sealed hydrocarbons, and coals in the Permo-Carboniferous portions of the stratigraphy sitting directly below the basaltic otflows. Notably, contact metamorphism and the development of hydrothermal systems rich in chlorine (produced from the pressure dissolution and volatilisation of the surrounding evaporites, kerogens, coals and hydrocarbons with evaporite seals) potentially synthesized large amounts of the organohalogens methyl chloride (CH3Cl) and methyl bromide (CH3Br) along with vast volumes of sulphurous gases, CH4 and CO2 (Figure 17).


    End-Triassic extinction event - Saline interactions with CAMP magmas

    The Central Atlantic Magmatic Province (CAMP) was emplaced at the end of the Triassic (≈201 Ma) in a region created by the tectonic unzipping (rifting-breakup) of the Pangean supercontinent (Figure 18; Marzoli et al., 2018). CAMP extends across the former Pangaea from modern central Brazil northeastward some 5000 km across western Africa, Iberia, and northwestern France, and from Africa westward for 2500 km through eastern and southern North America and as far west as Texas and the Gulf of Mexico (Figure 18 - dashed red line). The Province is composed of basic igneous rocks emplaced in a combination of shallow intrusions and erupted large lava flow fields extending over a land surface area in excess of 10 million km2. During its emplacement, sill intrusions into evaporites are particularly widespread in the vast Amazonas and Solimões intracratonic basins (≈1 ×106km2), representing up to 70% of the total CAMP sill volume (Svensen et al., 2018).


    Sedimentary rocks intruded by sills in the Amazonas and Solimões basins include a lower (Ordovician–Mississippian) and upper (Pennsylvanian–Permian) Paleozoic series (Milani and Zalán, 1999). The lower Paleozoic series consists of sandstones and shales, some of which are particularly organic-rich (total organic content up to 8wt.%; Milani and Zalán, 1999; Gonzaga et al., 2000). The upper Paleozoic series is dominated by evaporite and carbonate deposits of varying abundances, interlayered with clastics. Sills are widespread within the upper Paleozoic evaporitic sequence, extending almost continuously from the western margin of the Solimões Basin to the eastern margin of the Amazonas Basin (Fig.19c). Sills within the lower Paleozoic unit are restricted to the eastern part of the Amazonas basin. As illustrated in Fig.19c, high-Ti sills are found only in the lower Paleozoic series. Let's look now at the saline geology of the region and then at the effect its assimilation had on sill geochemistry.


    Saline geology

    A significant, as yet poorly delineated, set of variable hydrated potash salts and sylvinites occur in bedded halite in the Amazon Basin, Brazil (Figures 19a, 20; Szatmari et al. 1979). The Amazon Basin is about 2,100 km long and 300 km wide, it is an intracratonic sag basin atop an aulacogen between the Guyana and Guaporé cratons (Figure 19b). The basin fill contains a number of stacked mega-sequence cycles (as defined by wireline interpretation) ranging in age from Lower Ordovician (Autaz Mirim Member of Trombetas Formation) to Lower Permian (Figure 19b; Andirá Formation; Gonzaga et al., 2000). The basin has a widespread Upper Cretaceous cover (Alter do Chão Formation) and was affected by widespread tholeiitic magmatic activity at the end-Triassic (e.g.Penatecaua dolerites of the CAMP), making seismic-based hydrocarbon exploration difficult, especially as much of the basin still lies beneath thick tropical jungle. Since the recognition of a widespread igneous overprint of the Palaeozoic sedimentary succession in the 1970s, hydrocarbon exploration efforts have been subdued (Thomaz-Filho et al., 2008). However, in the past few years, SRTM studies are proving useful, in front of seismic surveys and drilling, in the general identification of geological features in the Amazon Basin (Ibanez et al., 2016)

    Th Amazonas-Solimoes intracratonic sag basin is developed on the same scale as the Alberta basin of Canada and entrains the Carboniferous (Pennsylvanian, ≈305 Ma) saline Nova Olinda Formation. It is made up of a large laterally extensive set of cyclic evaporite beds, dominated by interbedded combinations of anhydrite, shale and halite (Figures 20c, 21). These evaporites occur within the Carboniferous-Permian megasequence, known as the Tapajós Group, which can be up to 1600m thick (Milani and Zalan, 1999). The lowest part of the megasequence is a blanket of eolian sandstones (Monte Alegre Formation), which is covered by marine-influenced carbonates and evaporites (Itaituba and Nova Olinda Formations, respectively), along with subordinate sandstones and shales (Figure 19c). The Tapajós megacycle is closed by a suite of Permian continental redbeds (Andirá Formation) of Permian age. Subsequent east-west regional extension facilitated a pervasive intrusion of magmatic bodies during the end-Triassic to Early Jurassic (Penatecaua dolerites and equivalents).

    Individual halite beds in the Nova Olinda evaporite cycles are 20-80 m thick, while the Nova Olinda Fm. has an average thickness of 900m. Because of the high levels of entrained anhydrite beds in the Nova Olinda Fm., evaporite layers are not halokinetic, but are subject to collapse and flow about the basin margin, especially in areas of intense meteoric dissolution (Figure 20).


    Early Petrobras drilling programs conducted in the Amazon Basin from 1953 to 1963, defined the presence of halite but did not appreciate that persistent sylvinite/carnallite beds cap a number of the beds of NaCl in The Nova Olinda Formation. During the late 1960s and 1970s, higher-resolution gamma-ray logging tools were used, along with better mud technology and associated narrower calliper measures. This work identified a number of (0.5 - 2m thick) layers of sylvinite, within the halites (Szatmari et al. 1979). For example, the fifth and seventh depositional cycles define isolated salt sub-basins that accumulated significant potash salts in Fazendinha and Arari regions (Figure 20). KCl contents of these beds are between 28-33% in beds some 2.47-2.65 m thick (Garrett, 1995). The average ore depth at Fazendinha, the larger of the known potash areas, is 1,050m (Figure 20). Much of the halite and potash distribution is controlled by the underlying rift-basin architecture (Figure 19b). Potential potash reserves poorly defined, but are interpreted to be large (Szatmari et al., 1979; Garrett, 1995).

    Based on its texture, structure and chemistry, the potash intersection in the Amazon Basin is divided into three distinct zones, called informally, lower (milky or white sylvinite), middle (sulphates) and upper zones (red sylvinite) (Figure 20). The lower zone (milky-white sylvinite zone) contains sylvinite, with halite and subordinate intercalated kieserite and anhydrite beds. The lower potash zone is persistent within the basin and so covers an extensive area, whereas the upper potash zone is patchier. The greater extent of the lower potash zone is perhaps because it is the best isolated from any dissolution driven by circulation of undersaturated pore fluids through the overburden.

    The middle zone is composed of a combination of sulphate and chloride salts and is informally termed the sulphate zone. It hosts a variety of K, Mg and sulphate minerals that include a number of hydrated salts. Typical mineral assemblages encompass sylvinite, sylvite, and langbeinite (K2SO4.2MgSO4) as well as the hydrated salts; polyhalite (K2SO4.2MgSO4.2CaSO4.2H2O), kainite (MgSO4.KCl.3H2O) and kieserite (MgSO4.H2O). The sulphate distribution in this unit changes from anhydrite and polyhalite in the west (Fazendinha) to langbeinite and kainite in the east (Faro area). Towards the basin centre, chloride beds replace marginal sulphate beds in the sulphate unit. A gradual increase in potash concentration from west to east is interpreted by Sad et al., 1982, as indicating the inflow direction was from the basin's western boundary.

    The upper potash zone consists of coarsely-crystalline red sylvinite, with thin halite and anhydrite laminations. This level includes the best K2O grades drilled so far, averaging 23% K2O (between 33% to 16%). Red sylvinite is interpreted as a second generation product formed diagenetically by incongruent leaching of primary carnallite, but, as yet no carnallite (KCl.MgCl2.6H2O) has been identified in the upper unit.

    The potash zone is overlain by impermeable coarsely-crystalline halite, with minor shale intercalations in a zone up to 25 m thick, in turn, overlain by impermeable shale beds some 20 m thick. It is underlain by an impervious, at times sparry, halite interval some 70m thick (Figure 20). At the time it was described (1970s-mid 1980s) little was known of the significance of halite crystal textures in terms of their primary versus diagenetic signatures. Such a study of the nature of the halites enclosing the potash zone in the Amazon basin would aid in the definition of an ore genesis model. We do know that a single potash zone does not extend across the basin. This is seen in a compilation of existing Petrobras wells in the Amazon Basin, which intersect the Nova Olinda Fm. Instead, potash salts accumulated in a series of sumps atop a persistent thick halite unit (Figure 20).

    Elevated sulphate content in the potash zone of the Amazon Basin reflects the MgSO4-enriched nature of the world ocean during the Carboniferous. Potentially high levels of sulphate in proximity to adjacent sylvinite ore targets will complicate the processing of potential ore (see Warren 2016, Chapter 11). But in terms of supplying high levels of volatiles during sill intrusions, it is highly likely the various hydrated sulphate salts in the potash zone focused sill emplacement and contributed to elevated levels of halocarbons and sulphurous gases escaping into the earth's atmosphere at the end Triassic. As yet, no phreato-magmatic pipes have been documented in the Nova Olinda, but as the sourcing evaporite unit lie a kilometer beneath the surface and the dense tropical Amazon Jungle, this is not surprising. Increasing future use of STRM data may help solve this (Ibanez et alo., 2016)

     

    Saline sediment-sill interaction

    Sills from the Amazonas Basin have previously been described as low-Ti tholeiitic basalts and andesitic basalts De Min et al., 2003), and sills from both basins are generally characterised by a mineral assemblage of clinopyroxene, plagioclase, Fe–Ti oxides, rare olivine and orthopyroxene and accessory quartz-feldspar intergrowths. Recent studies report the presence of high-Ti sills in the eastern part of the Amazonas Basin (Figures 18, 21; Davies et al., 2017; Heimdal et al., 2018, 2019; Marzoli et al., 2018), but no high-Ti occurrences have been observed in the Solimões Basin. 


    High-precision U–Pb dates from four dolerites from the Amazonas and Solimões basins overlap in age, with U–Pb ages for low-Ti dolerites of 201.525 ±0.065 (Amazonas Basin) and 201.470 ±0.089 (Solimões Basin), and for high-Ti dolerites in the Amazonas Basin of 201.477 ±0.062 and 201.364 ±0.023 Ma (Figure 18; Davies et al., 2017; Heimdal et al., 2018). This suggests that low-and high-Ti CAMP magmatism were active simultaneously, although low-Ti magmatism likely started earlier.

    Detailed studies of CAMP sill geochemistry showing likely assimilation of chloride salts from the Nova Olinda evaporites are published in Heimdal et al., 2019, and summarised in this section. They show the bulk of e dolerites as sampled in the wells, illustrated in Figure 22, are characterised by phenocrysts of clinopyroxene and plagioclase in subophitic to intergranular textures, Fe–Ti oxides, and rare olivine and orthopyroxene. A different mineralogical assemblage (microphenocrysts of alkali-feldspar, quartz, biotite and apatite) is found in small independent domains, localised within the framework of coarser plagioclase and clinopyroxene laths. These fine-grained evolved domains crystallised in late-stage, evolved melt pockets in the interstitial spaces between earlier crystallised coarser grained crystals.


    The majority of the studied dolerites are generally evolved tholeiitic basalts and basaltic andesites with low TiO2 concentrations (<2.0 wt.%). Four samples have high TiO2 concentrations (>2.0 wt.%), and are found in the eastern part of the Amazonas Basin (Figure 20a, c).

    Whole-rock major and trace element and Sr-Nd isotope geochemistry of both low- and high-Ti sills is similar to that of previously published CAMP rocks from the two magma types. Low-Ti sills show enriched isotopic signatures (143Nd/144Nd201Ma from 0.51215 to 0.51244; 87Sr/86Sr201Ma from 0.70568 to 0.70756), coupled with crustal-like characteristics in the incompatible element patterns (e.g. depletion in Nb and Ta). Unaltered high-Ti samples show more depleted isotopic signatures (143Nd/144Nd201Ma from 0.51260 to 0.51262; 87Sr/86Sr201Maf from 0.70363 to 0.70398).

    Low-Ti dolerites from both the Amazonas and Solimões basins contain biotite with extremely high Cl concentrations (up to 4.7 wt.%). They show that there is a strong correlation between host-rock lithology and Cl concentrations in biotite from the dolerites, and interpret this to reflect large-scale crustal contamination of the low-Ti magmas by halite-rich evaporites (Figure 21). The findings of Heimdal et al. (2019) support the hypothesis that sill-evaporite interactions increased volumes of volatile released during the emplacement of CAMP, and underlines the case for the active involvement of this LIP in the end-Triassic extinction event.


    End-Cretaceous extinction event - Saline interactions driven by a bolide impact)

    About 66 million years ago, at the end of the Cretaceous, one or possibly multiple large asteroids collided with the Earth. Paul Renne dated this impact at 66.043±0.011 million years ago on the Yucatan Peninsula, based on argon-argon dating (Renne, 2013). He went on to conclude that the main end-Cretaceous mass extinction event occurred within 32,000 years of this date. The bolide produced a crater some 150x180 km in diameter named the Chicxulub impact structure (Figure 23). Worldwide, a record of this event is evidenced by an iridium-enriched interval, in what is now called the Cretaceous-Tertiary Boundary Clay (KTBC) (Alvarez et al., 1980).

    Other authors favouring additional bolide impacts at the end of the Cretaceous, such as Lerbekmo (2014) and Chaterjee (1997), have argued that some 40,000 years later, a much larger meteorite struck the shelf of the India-Seychelles continent, which was drifting northward in the southern Indian Ocean, producing a crater, some 450x600 km across, named the Shiva impact (Lerbekmo, 2014; Chaterjee, 1997). If a bolide-related feature, the Shiva crater was split by subsequent plate tectonism and today is not widely recognised by the scientific community as a K-T impact site.

    As for any sound scientific hypothesis, there are ongoing arguments for the Chicxulub site being the "smoking gun" for the end-Cretaceous extinction event, many of these arguments and the supporting literature is discussed in (Kring, 2007). I shall focus on the saline geology of the Yucatan impact site, but recognise the arguments of some authors that the Shiva site is closely linked in time with the extrusion of the Deccan Traps. More importantly, voluminous Deccan Traps eruptions and intrusions had likely already degraded the end-Cretaceous atmosphere. A large bolide crashing into an anhydrite saltern in the palaeo-Gulf of Mexico was perhaps the coup de grâce for many already -stressed late Mesozoic communities (Wang et al., 2018)


    Saline Geology of the Yucatan site

    As it is covered by a Tertiary-age sediment carapace, there are no current evaporite outcrops on the Yucatan Peninsula. However, the region is underlain by thick Cretaceous anhydrite beds and has a nearby giant oil field, Cantarell, reservoired in a carbonate breccia trap possibly related to the impact (Grajales-Nishimura et al., 2000). Ongoing petroleum exploration means a number of exploration wells sample the Cretaceous geology of the Yucatan Peninsula (inset in Figure 24). Regionally, Cretaceous (Albian) saltern anhydrite beds extend from Guatemala, across the Yucatan Peninsula and north possibly to Veracruz. Depositionally similar, back-reef saltern beds typify the early Cretaceous (Albian) Ferry Lake Anhydrite, which extends across the onshore northern, and offshore eastern, Gulf of Mexico (Pittman, 1985; Petty, 1995; Loucks and Longman, 1982).

    Pemex wells drilled on the Yucatan Peninsula, penetrate some 1300 –3500 m of bedded Tertiary, Cretaceous, and Jurassic strata (Figure 24; Ward et al., 1995). Palaeozoic metamorphic rocks are intersected at 2418 m in well Y4 and at 3202 m in well Y1. ‘‘Volcanic rock/andesite,’’ now broadly interpreted as an ‘‘impact-melt rock’’ or suevite is intersected in the lower parts of wells Y6 and C1. Based on the well geology there are seven major biostratigraphic-lithostratigraphic units in the Mesozoic section overlying basement rocks in the vicinity of the Chixulub impact site (Units A-F; Ward et al., 1995 and references therein). The regional depositional setting is typical of a Cretaceous carbonate platform, which at times became sufficiently isolated to deposit stacked anhydrite saltern beds in a rudistid back-reef setting (Warren, 2016; Chapter 5).

    Unit A consists of red and grey sandstone, shale, and silty dolomite near the base of wells Y1, Y2, and Y4. This unit is Jurassic to Early Cretaceous in age (López Ramos, 1975).

    Unit B is predominantly dolomite in its lower part, becomes rich in intercalated anhydrite and dolomite upward. Rock salt was cored in this unit in T1 at 2378–2381 m. Nummoloculina sp. was identified in Y2, suggesting an Albian age.

    Unit C is predominantly shallow-water limestone in the lower part, becoming more dolomitic upward. At the base of unit C in wells Y1 and Y2 is a horizon with the large benthic orbitulinid foraminifer Dicyclina schlumbergeri? (Figure 23\4). Nummoloculina (N. heimi?) also occurs in the lower part of this unit in cores Y1 and Y4. Nearer the platform margin (Y4), the upper part of this unit contains a rudist limestone, but in other wells the rocks reflect more restricted depositional environments across the platform interior. Shallow-subtidal to intertidal dolomite makes up most of this section in Y5A, where anhydrite is interlayered with dolomite in the upper parts of the unit in Y1, Y2, and T1. The fossil assemblage indicates an Albian-Cenomanian age for unit C.

    Unit D is predominantly somewhat deeper-subtidal limestone and marl, with horizons containing abundant tiny, mainly trochospiral planktic foraminifers as seen in samples from Y1, Y2-Y4, and Y5A.

    Unit E consists of shallow-platform limestones with intervals containing abundant small planktic foraminifers. The unit contains rudist-bearing limestones considerd by López Ramos (1975) as Turonian, and a similar age is indicated by the presence of Marginotruncana pseudolinneiana and Dicarinella imbricata in samples from Y1, Y2, Y4, and Y5A.

    Unit F consists of dolomitized shallow-platform limestone with benthic foraminifers. Abundant textularid and miliolid foraminifers are at the top of unit F (Fig. 2). The presence of Marginotruncana schneegansi and Globotruncana fornicata in well Y5A suggests a Santonian age for that part of this unit.

    Unit G is a thick interval of breccia with abundant sand- to gravel-sized angular to subrounded fragments of dolostone, anhydrite, and minor limestones suspended in a dolomicrite matrix. The poorly-sorted fabric is similar to that of debris-flow deposits. López Ramos (1973) reported marl and limestone intercalations within the thick breccia from 1090 to 1270 m in well C1 (Figs. 1 and 2). In addition, Y4 and Y4 contain dolomite that may separate an upper breccia with rare or no planktic foraminifers from a lower breccia with abundant planktic foraminifers. Core in Y2 is composed of finely crystalline anhydrite, possibly also representing a less disturbed sedimentary layer or anhydrite block within the breccia interval.

    Clasts of carbonate rocks in these breccias are fragments of many different kinds of dolostone and limestone, with different diagenetic histories. Anhydrite fragments typically make up 15%–20% of the breccia; much of the anhydrite is composed of tiny angular cleavage splinters. Some breccia layers contain grey-green fragments of altered volcanic ‘‘glass’’ and spherules. Other minor but significant constituents of the breccia are fragments of melt rock and basement as seen in Y6 (1295.5–1299 m), Y6 (1377–1379.5 m), and C1 (1393–1394 m). In addition, Hildebrand et al. (1991) found shocked quartz from Y6 (1208 –1211 m), and Sharpton et al. (1994) reported shock-deformed quartz and feldspar grains and melt inclusions in the dolomite-anhydrite breccia.

    Planktic and benthic foraminifers are present in the breccia matrix and include Abathomphalus mayaroensis, Globotrun-canita conica, Rosita patelliformis, Pseudoguembelina palpebra, Racemiguembelina fructicosa, and Hedbergella monmouthensis, which indicate a late Maastrichtian (end-Cretaceous) age for formation of the breccia (Ward et al., 1995).

    Climatic outcomes of the Yucatan impact

    Widespread Jurassic anhydrites, hydrocarbon reservoirs and source rocks surround the Yucatan impact region; their vaporisation on bolide impact and rapid entry into the upper atmosphere added a good deal to the ensuing climatic mayhem (Figure 25) As we discussed for LIP emplacement, anhydrite decomposes at high temperatures, to form SO2 gas, CaO, and oxygen. Thermodynamic calculation and extrapolation using the free energy of formation of anhydrite and its reaction products as a function of temperature up to 1120°C (Robie et al., 1979), give an equilibrium pressure of 1 bar SO2 over the reaction:

    2CaSO4 = 2CaO + 2S02 + O2

    at a temperature around 1500°C (Brett, 1992; Yang and Ahrens, 1998). Experimental studies by Rowe et al. (1967) indicate that anhydrite decomposes in an open crucible above 1200°C. Temperatures higher than 1500°C are well in the range of temperatures of material subjected to strong shock in large bolide impacts, and at higher temperatures the equilibrium pressure would be considerably higher. Because the system is open, SO2 and oxygen would escape to the atmosphere as they did in the laboratory crucible of Rowe et al. (1967) and would continue to do so as long as post-impact temperatures were elevated.

    Published discussions of the impact site geology all consider anhydrite as the evaporite mineralogy, with minor volumes of halite (well T1 in Figure 24). This lower salinity end of the evaporite series is typical of mega-sulphate settings, worldwide (Warren, 2016; Chapter 5) In addition, there is no evidence for hydrated potash salts in the region and this too is typical of starn salterns in a meg-sulphate basin. There is, however, the additional possibility that not all of the saltern gypsum had converted to anhydrite at the time of the impact. If so, this would have further detabilised and volatised the various lithologies at the site of the impact.


    Intercalated carbonates, kerogens and other organic sediments at the collision site contributed additional CO2, CH4, H2O, and halocarbons to the atmosphere, as well as vast quantities of heat and particulates. The following discussion of the various contributors to climatic changes, driven by the Chicxulub impact, is taken mostly from Kring, 2007 (and contained references).

    Acid rain; Because the Chicxulub impact occurred in a region with anhydrite, sulphurous vapour was injected into the stratosphere, producing sulphate aerosols and eventually sulphuric acid rain. Estimates of the amount of S liberated vary, consensus ranges from 7.5 × 1016 to 6.0 × 1017 g S, which would have produced 7.7 × 1014 to 6.1 × 1015 mol of sulphuric acid rain. In addition, the earth’s atmosphere was shock-heated by the impact event, producing nitric acid rain as well. Independent of the geology of the impact siter, the earth's atmosphere is heated when pierced by a bolide as the vapour-rich plume expands out from an impact site, and ejected debris rains through the atmosphere. In a Chicxulub-sized impact event, the ejecta debris is, estimated to produce ≈1×1014 mol of NOx in the atmosphere and, thus, ≈1×1015 moles of nitric acid rain. Impact-generated wildfires may have produced an additional ≈3×1015 mol of nitric acid. Sulphuric and nitric acid rain fell over a few months to a few years (Figure 25a).

    Wildfires; Evidence of impact-generated fires is recovered from K/T boundary sequences worldwide in the form of fusinite pyrolitic polycyclic aromatic hydrocarbons, carbonised plant debris, and charcoal. The distribution of the fires is still poorly understood and may have had a restricted geographic distribution limited to the vicinity of the impact event, produced not by impact ejecta but by the direct radiation of the impact fireball which had a plasma core with temperatures over 10,000 °C. Several additional parameters influence the outcome (e.g., the trajectory of the impacting object, its speed, and mass of the ejecta). The amount of soot recovered from K/T boundary sediments (imply that the fires released ≈104 GT of CO2, ≈102 GT CH4 and 103 GT CO, which is equal to or larger than the amount of CO2 produced from vapourised target sediments. This likely had a severe effect on the global carbon cycle (Figure 25a).

    Dust and aerosols in the atmosphere; Calculations suggest that dust and sulphate aerosols from the impact event, and soot from post-impact wildfires, caused surface temperatures to fall by preventing sunlight from reaching the surface where it was needed for photosynthesis. The base of the marine food chain, composed of photosynthetic plankton, collapsed. Slight increases or decreases in average water temperatures cannot extinguish photosynthetic plankton, nor the presence or absence of organisms higher up the food chain. Photosynthesisers are primarily affected by the availability of their energy source, light. Consequently, the loss of photosynthetic plankton following the Chicxulub impact event is evidence that sunlight was significantly blocked, whether it was by dust, soot, aerosols, or some other agent.

    The timescale for particles settling through the atmosphere range from a few hours to approximately a year (Figure 25a, b). The time needed for the bulk of the dust to settle out of the atmosphere is ambiguous, however, because the size distribution of the dust is unclear. Some sites seem to be dominated by spherules ≈250 μm in diameter, which would have settled out of the atmosphere within hours to days. However, if there is a substantial amount of submicron material, then it may remain suspended in the atmosphere for many months. Soot, if it were able to rise into the stratosphere, would have taken similarly long times to settle. Soot that only rose into the troposphere, however, would have been flushed out of the atmosphere promptly by rain.

    The dust, aerosols, and soot caused surface cooling after the brief period of atmospheric heating that immediately followed the impact. The magnitude of that cooling is unclear, however, because the opacity generated by the three components is uncertain and their lifetime in the atmosphere is also uncertain. Nonetheless, significant decreases in temperature of several degrees to a few tens of degrees have been proposed for at least short periods. Short-term cooling likely had a severe effect on the global carbon cycle, in what is popularly termed a “nuclear winter’ scenario (Figure 25).

    Ozone destruction; Ozone-destroying Cl and Br is produced from the vaporised projectile, vaporised target lithologies, and biomass burning. Over five orders of magnitude more Cl than is needed to destroy today's ozone layer was injected into the stratosphere, compounded by the addition of Br and other reactants. The affect on the ozone layer may have lasted for several years, although it is uncertain how much of an effect it had on surface conditions. Initially, dust, soot, and NO2 may have absorbed ultraviolet radiation, and sulphate aerosols may have scattered the radiation. The settling time of dust was probably rapid relative to the time span of ozone loss, but it may have taken a few years for the aerosols to precipitate.

    Greenhouse gases; Water and CO2 were produced from Chicxulub's target lithologies and the projectile, which could have potentially caused greenhouse warming after the dust, aerosols, and soot settled to the ground. Significant CO2, CH, and H2O were added to the atmosphere. Some of these components came directly from target materials. These include carbonates, which liberate CO2 when vaporised, and also includes hydrocarbons, the remainder of which has subsequently migrated into cataclastic dykes beneath the crater and impact breccias deposited along the Campeche Bank (e.g. Cantarell field). Water was liberated from the saturated sedimentary sequence and the overlying ocean (the lesser of the two sources).

    The residence times of gases like CO2 are greater than those of dust and sulphate aerosols, so greenhouse warming may have occurred after a period of cooling. Estimates of the magnitude of the heating vary considerably, from an increase of global mean average temperature of 1 to 1.5 °C (based on estimates of CO2 added to the atmosphere by the impact) to ≈7.5 °C (based on measures of fossil leaf stomata).

    Local and regional effects; The local and regional effects of the impact were enormous. Tsunamis radiated across the Gulf of Mexico, crashing onto nearby coastlines, and also radiated farther across the proto-Caribbean and Atlantic basins. Tsunamis were 100 to 300 m high when they crashed onto the gulf coast and ripped up sea floor sediments down to water depths of 500 m. The Gulf of Mexico region was also affected by the high-energy deposition of impact ejecta, density currents, and seismically-induced slumping of coastal sediments following magnitude 10 earthquakes. Tsunamis may have penetrated more than 300 km inland. The local landscape (both continental and marine) was buried beneath a layer of impact ejecta that was several hundred meters thick near the impact site and decreased with radial distance. Peak thicknesses along the crater rim may have been 600 to 800 m. Along the Campeche bank, 350 to 600 km from Chicxulub, impact deposits of ≈50 to ≈300 m are logged in the Cantarell boreholes.

    Impact events also produce shock waves and air blasts that radiate across the landscape. Wind speeds over 1000 km/h are possible near the impact site, although they decrease with distance from the impact site. The pressure pulse and winds can scour soils and shred vegetation and any animals living in nearby ecosystems. Estimated radii of the area damaged by an air blast range from ≈900 to ≈1800 km.

    Significant heat would have been another critical regional effect. Core temperatures in the plume rising from the crater were over 10,000° C, possibly high enough to generate fires out to distances of 1500 to 4000 km. The intense thermal pulse would have been relatively short-lived (5 to 10 min). Additional heating and spontaneous wildfires were ignited when impact ejecta fell through the atmosphere (3 to 4 days; Figure 25a).

    The end-Cretaceous bolide impact had both short and long term effects on the Earth's climate and its atmospheric temperatures (Figure 25b). Over hours to days following impact, there was severe atmospheric heating as ejecta rained down through the atmosphere. This was following by a period of weeks to years of cooler temperatures as the atmosphere was polluted by SO2, NOx and soot from the impact preventing sunlight reaching the surface (nuclear winter scenario). Then, across time frame of decades to millennia, after the atmosphere cleared, increased CO2 levels drove a period of global warming. The legacy of the impact and the biotal recovery over the next few hundred thousand years is documented in a recent paper by Lowery et al., 2018. They showed that life reappeared in the basin just years after the impact and a high-productivity ecosystem was established within 30 kyr.

    Extinction events intensified by heating evaporites

    Evaporite salts are more chemically reactive at earth surface conditions than other sediments. Subsurface evaporites are prone to dissolution, alteration and reprecipition from the time they first precipitated and throughout their subsurface journeys in the diagenetic and metamorphic realms (Warren 2016). The same is true, but perhaps more so, if bedded salts are exposed to a heat source outside the normal geothermal gradient experienced in burial. Additional heat can come for the emplacement of igneous sills, magma bodies or the hot hydrothermal circulation it drives. Or it can come from near instantaneous heating to thousands of degrees associated with a bolide impact. Volatile products that result from this heating, as they enter the earth's atmosphere, can be inimical to life and include vast volumes of halocarbons, SO2, methane and CO2. Methane and CO2 come from kerogens and hydrocarbons stored in intercalated mudstones and limestones while volatilisation of carbonates can supply CO2.

    The reactivity of evaporites and the vast volumes of volatiles released explains the intimate association of saline giants, heating and the three most devastating of the five major Phanerozoic extinction events.

    Interestingly, two other events on the list of the "big five;" the Emeishan and late Devonian events (Figure 1) also have possible associations with heated evaporites. The Emeishan LIP intersects the edge of the anhydrite-rich Sichuan basin, while the 120km-diam., Late Devonian, Woodleigh bolide impacted the intracratonic Silurian Yaringa Fm. salts (including potash beds) on the coast of West Australia (SaltWork GIS database version 1.8 overlays, Chen et al., 2018; Glikson et al, 2005). But, before definitive conclusions can be made, more work is required to better tie down impact age, actual geographic extent of LIP emplacement, extent of evaporite breccias and evaporite volumes.

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    Evaporite interactions with magma Part 2 of 3: Nature of volatile exhalations from saline giants?

    John Warren - Saturday, March 16, 2019

     

    Introduction

    This article discusses general mechanisms of earth-scale volatile entry into the ancient atmosphere during events that involved rapid and widespread heating of saline giants. It develops this notion by looking at whether volumes of volatiles escaping to the atmosphere are enhanced by either the introduction of vast quantities of molten material to a saline giant or the thermal disturbance of that salt basin by bolide impacts. This begins a discussion of the contribution of heated evaporites in two (or three if the Captitanian is counted as a separate event) of the world's five most significant extinction events. It also looks at possible evaporite associations with a substantial bolide impact that marks the end of the Cretaceous. The next article presents the geological details and implications of the various magma-evaporite-volatile associations tied to major extinction events.

    As we have seen for evaporite interactions with giant and supergiant volumes of commodities in particular deposits, such as hydrocarbons, base metals (Cu, Pb-Zn and IOCG deposits) evaporites do not form a commodity accumulation. But if evaporites are involved in the accumulation and enrichment processes, the size and strength of the accumulation are much improved. Because of their high reactivity compared to the kinetic stability at and near  thelithosphere's surface across most other lithologies, evaporite act not as creators of enrichment but as facilitators of enrichment (Warren, 2016 Chapters 9, 10, 14, 15 and Salty Matters, March 31, 2017).


    End-Permian event

    The end-Permian extinction event, colloquially known as the Great Dying, occurred around 252 Ma (million years) ago, and defines the boundary between the Permian and Triassic geologic periods, as well as between the Palaeozoic and Mesozoic eras. It is the Earth's most severe extinction event, when up to 96% of all marine species, 70% of terrestrial vertebrate species disappeared (Table 1, Figure 1). It also involves the only known mass extinction of a number of insect species (≈25%). Some 57% of all biological families and 83% of all genera became extinct. The end-Cretaceous extinction, which marks the demise of dinosaurs, is less severe, although it probably has a stronger hold on the western zeitgeist, while on land, the end-Triassic event marks the ascendancy of the dinosaurs.


    Suggested mechanisms driving the end-Permian extinction event include; massive volcanism centred on the Emeishan and Siberian Traps and the ensuing coal or gas fires and explosions, along with a runaway greenhouse effect that was triggered by temperature increases in marine waters (Figure 2). It also may have involved one or more large meteor impact events and a rise in oceanic water temperatures that drove a sudden release of methane from the sea floor due to methane-clathrate dissociation.

    The end-Permian event follows on closely from the Capitanian (Emeishan) extinction event when in south China fusulinacean foraminifers and brachiopods lost 82% and 87% of species, respectively (Bond et al., 2015). Proximity in time of the two events may explain why the breadth of the end-Permian extinction event was so severe. The Earth's biota was still recovering from the Emeishan event when the vicissitudes of the End-Permian calamity further decimated the world's biota.

    Both the Emeishan and end-Permian extinction events tie to elevated mercury levels in sediments that encompass their respective boundaries (Grasby et al., 2016). Astride both boundaries, the mercury stratigraphy shows relatively constant background values of 0.005–0.010 μg g–1. However, there are notable spikes in Hg concentration over an order of magnitude above background associated with the two extinction events. The Hg/total organic carbon (TOC) ratio shows similar large spikes, indicating that they represent a real increase in Hg loading to the environment. These Hg loading events are associated with enhanced Hg emissions created by the outflows of the Emeishan and end-Permian large igneous province (LIP) magmas.

    Interestingly, there is indirect evidence for a synchronous antipodeal impact crater that some argue may have instigated the Siberian volcanism, in much the same way that the end-Cretaceous bolide impact on the Yucatan Peninsula is considered by some to be the antipodeal driver of the Deccan Trap volcanism (von Frese et al., 2009). Other contributing, but likely more gradual tiebacks to the Great Dying, include sea-level variations, increasing oceanic anoxia, increasing aridity tied to the accretion of the Pangean supercontinent, and shifts in ocean circulation driven by climate change (Figure 2).

    End-Triassic event

    The end-Triassic extinction event, some 201.3 Ma, defines the Triassic-Jurassic boundary. In the oceans, a whole class (conodonts) and 23-34% of marine genera disappeared. On land, all archosaurs other than crocodylomorphs (Sphenosuchia and Crocodyliformes) and Avemetatarsalia (pterosaurs and dinosaurs), some remaining therapsids, and many of the large amphibians became extinct. About 42% of all terrestrial tetrapods went extinct (Figure 3). This event vacated terrestrial ecological niches, allowing the dinosaurs to assume the dominant roles in the Jurassic period. It happened in less than 10,000 years and occurred just before the Pangaean supercontinent started to break apart (Tanner, 2018).


    The extinction event marks a floral turnover as well. About 60% of the diverse monosaccate and bisaccate pollen assemblages disappear at the T-J boundary, indicating a significant extinction of plant genera. Early Jurassic pollen assemblages are dominated by Corollina, a new genus that took advantage of the empty niches left by the extinction.

    Worldwide the end-Triassic extinction horizon is marked by perturbations in ocean and atmosphere geochemistry, including the global carbon cycle, as expressed by significant fluctuations in carbon isotope ratios (Korte et al., 2019). At this time the Central Atlantic Magmatic Province (CAMP) volcanism triggered environmental changes and likely played a crucial role in this biotic crisis (Schoene et al., 2010). Biostratigraphic and chronostratigraphic studies link the end-Triassic mass extinction with the early phases of CAMP volcanism, and notable mercury enrichments in geographically distributed marine and continental strata are shown to be coeval with the onset of the extrusive emplacement of CAMP (Percival et al. 2017; Marzoli et al., 2018). Sulphuric acid induced atmospheric aerosol clouds from subaerial CAMP volcanism can explain a brief, relatively cool seawater temperature pulse in the mid-paleolatitude Pan-European seaway across the T–J transition. The occurrence of CAMP-induced carbon degassing may explain the overall longterm shift toward much warmer conditions.

    End-Cretaceous event

    The end-Cretaceous extinction event defines Cretaceous-Tertiary (K–T) boundary, and was a sudden mass extinction event some 66 million years ago. Except for some ectothermic species, such as the leatherback sea turtle and crocodiles, no tetrapods weighing more than 25 kilograms survived. The K-T event marked the end of the Cretaceous period and with it, the entire Mesozoic Era, opening the Cenozoic Era.

    A wide range of species perished in the K–T extinction, the best-known being the non-avian dinosaurs. It also destroyed a plethora of other terrestrial organisms, including certain mammals, all pterosaurs, some birds, lizards, insects, and plants. In the oceans, the extinction event killed off plesiosaurs and the giant marine lizards (Mosasauridae) as well as devastating fish, sharks, molluscs (especially ammonites, which became extinct) populations, and many species of plankton. It is estimated that 75% or more of all species on Earth vanished in the end-Cretaceous event.

    In its wake, the same extinction event also provided evolutionary opportunities as many groups underwent remarkable adaptive radiation—sudden and prolific divergence into new forms and species within the disrupted and emptied ecological niches. Mammals in particular diversified in the Paleogene, evolving new forms such as horses, whales, bats, and primates. Birds, fish, and perhaps lizards also radiated in newly vacant niches.


    In the geologic record, the K–T event is marked by a thin layer of sediment called the K–Pg (Cretaceous - Paleogene) boundary, that is found throughout the world in both marine and terrestrial rocks. The boundary clay shows high levels of the metal iridium and is widely interpreted as indicating the impact of a massive comet or asteroid 10 to 15 km (6 to 9 mi) wide some 66 million years ago (Figure 4a,b). The impact devastated the global environment, mainly through a lingering impact winter, which halted photosynthesis in plants and plankton.

    The impact hypothesis, also known as the Alvarez hypothesis (Alvarez et al., 1980), was bolstered by the discovery of the 180-kilometer-wide (112 mi) Chicxulub crater in the Gulf of Mexico in the early 1990s, which provided conclusive evidence that the K–Pg boundary clay represented debris from an asteroid impact. In a 2013 paper, Paul Renne dated the impact at 66.043±0.011 million years ago, based on argon-argon dating (Renne, 2013). He went on to conclude that the main end-Cretaceous mass extinction event occurred within 32,000 years of this date. A 2016 drilling project into the Chicxulub peak ring, confirmed that the peak ring was comprised of granite, likely ejected within minutes from deep in the earth, but the well contained hardly any anhydrite/gypsum, the usual sulphate-containing seafloor rock across the region (Figure 4a, b). As we shall see in part 3, the missing CaSO4 was vaporised in the impact and dispersed as sulphurous aerosols into the atmosphere, causing longer-term deleterious effects on the climate and food chain. Another causal or contributing factors to the end-Cretaceous extinction event may have been the synchronous outflows of the Deccan Traps and other volcanic eruptions, so driving climate change, and possibly sea level change (von Frese et al., 2009).

    Volatiles released when cooking saline giants and associated organic-rich sediments

    Particular sets of assimilations and metamorphic alterations of evaporites occur within the explosive milieu associated with both igneous interactions and pressurised heating of salts tied to a bolide impact. Any carbonate and organic matter layers present in the saline sequence or adjacent strata generates additional volatiles that will quickly enter the earth's atmosphere. Figure 5 is a schematic of the estimated amount of volatiles released during contact metamorphism of different types of sedimentary rocks in contact with an igneous sill or magma body (after Ganino et al., 2009; Pang et al., 2013). More catastrophic volumes of similar volatile suites enter the atmosphere if a large bolide impacts a region underlain by a saline giant.


    Hence, salty interactions must be considered and quantified when attempting to understand earth-scale environmental changes whenever large evaporite masses are caught up in regions of LIP emplacement or bolide impact. In such areas:

  • Basalt and granitoids do not release large volumes of volatiles, as compared to the amounts of volatiles that are released by the heating or assimilation of saliniferous country rock (heat transfer and hydrothermal circulation).
  • Most porous sandstones and organic-lean shales caught up in a contact aureole or consumed in a magma, release water vapour; a release that has little effect on global climate.
  • During desulphation of a magma, gypsum or anhydrite masses are assimilated into a rising magma chamber or the emplacement of a thick sill. If anhydrite beds are consumed (melted and absorbed) by a magma batholith, the reaction releases abundant SO2 constituting up to 47 wt% of the bedded sulphate (Gorman et al., 1984). Direct melting requires high temperatures (≈ 1300- 1400 °C). Such widespread desulphation of thick Devonian anhydrite beds occurred during the emplacement of the supergiant Noril'sk nickel deposit in Siberia (Black et al., 2014; Warren, 2016, Chapter 16).
  • But such elevated temperatures (≈1400°C) are not typical of most contact aureoles where a sill or dyke intrudes anhydritic country rock. However, similar high-volume SO2 releases can proceed at temperatures as low as 615°C if the anhydrite is impure and contains interlayers rich in organics and hydrocarbons (e.g., West and Sutton, 1954; Pang et al., 2013). This is especially so if the interacting calcium sulphate is gypsum (hydrated salt) rather than anhydrite. Experiments by Newton and Manning (2005) demonstrated that the solubility of anhydrite increases enormously with NaCl activity (salinity) in hydrothermal solutions at ≈600 to 800°C (Figure 6).


  • Pure limestone contains large amounts of CO2, but like anhydrite the thermal decomposition of limestone or dolomite into CaO, MgO and CO2 takes place at high temperatures (>950 °C) that are typical when blocks of sedimentary carbonate are assimilated into a magma chamber, but less typical of contact aureoles tied to dykes and sills. Impure limestones can release large amounts of CO2 (up to 29 wt%) during the formation of calc-silicates in the contact aureole at moderate temperatures of 450–500 °C. As early as 1940, Bowen documented the release of CO2 by decarbonisation reactions during progressive metamorphism of siliceous dolomites (Bowen, 1940)
  • Likewise, devolatilization of fine-grained calcareous and saline sedimentary rocks during contact metamorphism directly generates fluids rich in CO2 (i.e., decarbonisation) and SO2 (i.e., desulphatation), which in theory can enter the magmatic system.
  • When heated at a relatively low temperature (<300-400 °C), contact metamorphism and hydrothermal leaching of bituminous halite and organic-carbon-rich saline mudstones releases large volumes of chlorohalogens and methane (Visscher et al., 2004; Beerling et al., 2007). Halocarbon compounds (aka halogenated hydrocarbons) are chemicals in which one or more carbon atoms are linked by covalent bonds with one or more halogen atoms (fluorine, chlorine, bromine or iodine). Methyl chloride (CH3Cl) and methyl bromide (CH3Br) are commonplace halocarbons when a halite-dominant saline giant interacts with igneous sill emplacement. When thermally-derived chlorohalogens enter the upper atmosphere, they tend to be reactive and will degrade ozone.
  • Buring coal and coal gas release abundant CO2. Depending on its grade, coal can ignite at temperatures between 400-530°C. Methane will auto-ignite at temperatures around 550-600°C and in an oxygenated setting produces large volumes of carbon dioxide and water vapour. Flashpoints are much lower than these ignition temperatures.
  • Sulphidic (pyritic) sediments release abundant SO2 when heated at lower temperatures (<400°C).
  • Heating of hydrated salts at moderate temperatures (90-250°C) can release pressurised pulses of hypersaline chloride or sulphate brine, with the dominant ionic proportions dependent on predominant hydrated salt; e.g., carnallite incongruently alters as it releases an MgCl2 brine, gypsum incongruently alters as it releases a Ca-SO4 brine (see part 1). Such pressurised pulses are essential in the generation of explosive breccia pipes sourced at the sill penetration level in the hydrated evaporite interval (discussed in detail for the Siberian Traps in part 3).
  • Getting volatiles into the atmosphere

    When a saline giant is heated during emplacement of a large igneous province (LIP) or during the impact of a large bolide, it and adjacent carbonates and organic-rich mudstones release large volumes of volatiles that can have short and long term harmful effects on the Earth's biosystems (Black et al., 2012, 2014; Jones et al., 2016; Part 3 this series). The volume of volatiles released to the atmosphere by these interactions, especially sulphurous products (SO2, H2S), thermogenic CH4, organohalogens and CO2, are considered primary contributors to three or four of the major extinction events outlined in Figure 1, and perhaps others, as discussed in part 3.

    Height and volume of various volatile injections into the layers of Earth's atmosphere controls the longevity and intensity of climatic effects and are tied to the chemistry of particular volatiles (Figure 7; Textor et al., 2003; Robock, 2000). The low concentration of water in typical modern volcanic plumes results in the formation of relatively dry aggregates entering the atmosphere. More than 99% of these aggregates are frozen because of their fast ascent to low-temperature regions of the atmosphere. With increased salinities, the salinity effect increases the amount of liquid water attaining the stratosphere by one order of magnitude, but the ice phase is still highly dominant. Consequently, the scavenging efficiency for HCl is very low, and only 1% is dissolved in liquid water.


    Scavenging by ice particles via direct gas incorporation during diffusional growth is a significant process for volatile transport. The salinity effect increases the total scavenging efficiency for HCl from about 50% to about 90%. The sulfur-containing gases SO2 and H2S are only slightly soluble in liquid water; however, these gases are incorporated into ice particles in the atmosphere with an efficiency of 10 to 30%. Despite scavenging, more than 25% of the HCl and 80% of the sulphur gases reach the stratosphere during a more intense modern explosive eruption because most of the particles containing these species are typically lifted there by the force of the eruption (Figure 7b).

    Sedimentation of the particles tends to remove the volcanic gases from the stratosphere. Hence, the final quantity of volcanic gases injected in a particular eruption depends on the fate of the particles containing them, which is in turn dependent on the volcanic eruption intensity and environmental conditions at the site of the eruption.

    Today, volcanically-derived SO2 and H2S are the dominant sources for sulphur species in the atmosphere (Jones et al., 2016; Robock, 2000). Conversion of SO2 to aerosols is one of the critical drivers of climatic cooling during recent eruptions (Figure 7a; Robock, 2000). For SO2 to be effective in causing cooling in the atmosphere, escaping hydrogen sulphide quickly oxidises to SO2. Over hours to weeks following its eruptive escape the ongoing reaction of SO2 with atmospheric H2O forms a H2SO4 (sulphuric acid) aerosol, and this is a major cause of the acid rains tied to volcanism (Figure 7a, b).

    Tropospheric sulphate aerosols have an atmospheric lifetime of a couple of weeks due to the rapid incorporation as precipitation into the hydrological cycle (Figure 7b; Robock, 2000). However, if the intensity of the escaping volatile plume is capable of injecting sulphurous material above the tropopause into the stratosphere, then due to the lack of removal by precipitation, the lifetimes of sulphurous aerosols and the associated cooling effects are considerably extended (years rather than weeks: Figure 7a versus 7b).

    Modern eruptions

    World-scale cooling has been observed following a number of modest (by large igneous province standards) volcanic eruptions over the past few centuries (Figure 8; Bond and Wignall, 2014; Sigurdsson, 1990; and references therein). A recent example is provided by the Mount Pinatubo eruption of 1991, which injected 20 megatons of SO2 more than 30 km into the stratosphere. The result was a global temperature decrease approaching 0.5 °C for three years (although this cooling was probably exacerbated contemporaneous Mount Hudson eruption in Chile). One of the largest historical eruptions occurred in 1783-1784 from the Laki fissure in Iceland when a ≈15 km3 volume of basaltic magma was extruded, releasing ≈122 Mt of SO2, 15 Mt of HF, and 7 Mt of HCl. Laki’s eruption columns extended vertically up to 13 km, injecting sulfate aerosols into the upper troposphere and lower stratosphere, where they reacted with atmospheric moisture to produce ≈200 Mt of H2SO4. This aerosol-rich fog hung over the Northern Hemisphere for five months, leading to short-term cooling, and harmful acid rain in both Europe and North America. Additionally, HCl and HF emissions damaged terrestrial life in Iceland and mainland Europe, as this low-level fluorine-rich haze stunted plant growth and acidified soils.

    By causing or aiding in the collapse of food chains during the more intense sulphurous releases involved in the heating of large volumes of anhydrite held in ancient saline giants, vast quantities of acid rain may have killed much of the vegetation on land and photosynthetic organisms in the oceans during the three extinction events discussed in part 3.


    Halocarbons

    For halocarbons to form in a volcanic eruption requires the combination halogens with organic matter/methane or other hydrocarbons. We shall consider the levels and origins of two of the more common halocarbons in today's atmosphere; methyl chloride (CH3Cl) and methyl bromide (CH3Br) although many other species of halogenated hydrocarbons are present both naturally and anthropogenically (Schwandner, 2002; Visscher et al., 2004).

    The average Cl concentration of the Earth has been estimated to be 17 ppm (Worden, 2018 and references therein). Chlorine is the dominant anion in seawater, most modern and ancient evaporite beds and associated brines. Chlorine is present in most igneous rocks at low concentrations with little difference in level shown between granite and basic igneous rocks (both have a Cl- concentration of about 0.02%). However, igneous glass typically has higher Cl concentrations (≈0.08%). Chlorine is concentrated within any residual vapour phase during volcanic eruptions so can be independent of the volatiles created by heating of saline giants. Without the latter, the contribution of volcanically-erupted Cl to the atmosphere is still considerable. For example, the estimated current global volcanic emission of Cl is between 0.4 and 170 mt/year, while individual eruptions can produce hundreds of kilotons of Cl. For example, in 1980, St Helens emitted 670 kt of Cl into the atmosphere.

    In crystalline igneous rocks Br is found at low concentrations, typically <1 ppm in mid-ocean ridge basalts (MORB) (Worden, 2008 and references)). The average Br concentration of the Earth has been estimated to be 0.05 ppm. Chlorine/Bromine ratios are typically between 200 and 1000 in igneous rocks. Bromine is, however, found at relatively high concentrations (up to 300 ppm) in melt inclusions and matrix glass in acid igneous rocks since it is a highly incompatible element that does not easily sit within silicate, oxide or sulphide minerals. Bromine is concentrated within any residual vapour phase during volcanic eruptions. Based on experimentally-derived fractionation factors for halogens in volcanic materials, crustal average halogen concentrations, and measured amounts of Cl emitted from volcanoes, it can be concluded that the contribution of volcanically-erupted Br to the atmosphere is considerable. For example, the estimated current global volcanic emission of Br is between 2.6 and 78 kt while individual eruptions (e.g., St Helens in 1980) can emit 2.4–5.6 kt.

    The hinterlands of sedimentary basins that predominantly enriched in primary igneous rocks will provide only small quantities of Br into the sediment supply but rocks enriched in glass-bearing igneous rocks may supply relatively greater amounts of Br (Worden, 2018). Bromine is found in sedimentary basins as dissolved Br-, in solid solution in halite (NaClxBr1−x), or in less common salts resulting from potash-facies evaporites, such as sylvite. Bromine is also associated with organic-rich sediments, especially in marine settings, including organic-rich mudstone and coal. At a concentration of 65 mg/L, Br- is the second most abundant halogen in modern seawater.

    Organic matter and its more evolved forms –kerogen and hydrocarbons– are typical of most large evaporite basins. Mesohaline carbonates interlayered with anhydrite and halite beds can entrain high levels of organic matter to form high-yield source rocks, while the brine inclusions in some halites contain high amounts of volatile hydrocarbons and pyrobitumens. Evaporite beds composed of anhydrite or halite make excellent seals holding back large volumes of hydrocarbons (for literature documentation of these observations see Warren, 2016, Chapters 9 and 10). In combination, saline giants and their heat-responsive lithologies will contain vast volumes of potential volatiles, including halocarbons.

    Ozone (O3) destruction

    When halocarbons enter the stratosphere, they decimate the ozone layer, allowing harmful levels of ultraviolet (UV) radiation to reach the earth's surface (Figures 7a, 9a). Ozone is destroyed by the entry of a number of free radical catalysts into the stratosphere; today the most important catalysts are the hydroxyl radical (OH), nitric oxide radical (NO), chlorine radical (Cl) and the bromine radical (Br). Each radical is characterised by an unpaired electron in its molecular structure and is thus extremely reactive. All of these radicals have both natural and man-made sources; at present, most of the OH and NO in the stratosphere is naturally occurring, but human activity has drastically increased the levels of chlorine and bromine.

    The elements that form radicals in the stratosphere are found in stable organic compounds, especially halocarbons, which reach the stratosphere without being destroyed in the troposphere due to their low reactivity. Once in the stratosphere, the Cl and Br atoms are released from the parent halocarbon by the action of ultraviolet light.


    Ozone (O3) is a highly reactive molecule that quickly reduces to the more stable oxygen (O2) form with the assistance of a catalyst (radical). Cl and Br atoms destroy ozone molecules through a variety of catalytic cycles. The simplest example of such a reaction is when a chlorine atom reacts with an ozone molecule, taking an oxygen atom to form chlorine monoxide (ClO) and leaving behind an oxygen molecule (O2) (Figure 9b). The ClO can then react with another molecule of ozone, once more releasing the chlorine atom as ClO, so far yielding two molecules of oxygen. This ClO reaction can be repeated until the ClO is flushed from the stratosphere (Figure 9b, Fahey, 2007)

    Thus the overall effect of halocarbons entering the stratosphere is a decrease in the amount of ozone. A single chlorine radical can continuously destroy ozone for up to two years (this the time scale for its transport back down into the troposphere; Figure 7a). But there are other stratopheric reactions that remove CLO from this catalytic cycle by forming reservoir species such as hydrogen chloride (HCl) and chlorine nitrate (ClONO).

    Bromine radicals are even more efficient than chlorine at destroying ozone on a per-atom basis, but at present there is much less bromine than chlorine in the atmosphere. Laboratory studies have shown that fluorine and iodine atoms can participate in similar catalytic cycles. However, fluorine atoms react rapidly with water and methane to form strongly bound HF in the Earth's stratosphere, while organic molecules containing iodine react so quickly in the lower atmosphere that they do not reach the stratosphere in significant quantities.

    Halocarbon concentrations below the tropopause are always higher by several orders of magnitude than in the stratosphere, which contains the seasonally and locally variable ozone layer responsible for absorption of incident solar UV radiation (Schwandner, 2002). Penetration of the tropopause allows the ascent of long-lived halocarbons and today occurs primarily as a result of rising tropical air masses in a Hadley cell, rare turnover events, or large Plinian volcanic eruptions.

    Over the two to three years a chlorine or bromine radical can remain in the stratosphere, it reacts with ozone and converts it to oxygen. It has been estimated that a single chlorine atom can react with an average of 100,000 ozone molecules before it is removed from the catalytic cycle (Figure 8b. Other halocarbon-enabled reactions drive ozone destruction (these catalysts are derived from anthropogenic CFCs and other industrial halocarbons). Over the past half-century, our anthropogenic focus on ozone destruction from industrial chemicals has driven the public's understanding into to the much-needed legislated prevention of the entry of additional industrial halocarbons (especially CFCs) into the stratosphere.


    Implications

    However, there are additional deep-time implications for the health of the Earth's biota when natural events of the past drastically increased the amount of halocarbons entering the stratosphere, along with increased levels of sulphurous volatiles and greenhouse gases. We know modern volcanic exhalations containing relatively high levels of chlorine and bromine. But times of intense magmatic/volcanogenic or bolide heating of evaporites in a saline giant will contribute even greater volumes of halocarbons to the stratospheric levels of the atmosphere (Figure 10). If coals and peats are also present (typically not in the saline portion of the basin's sediment fill), then the heating of these additional organic-rich sediments will contribute even more carbon to the vast volumes of the halocarbons created by heating of the evaporites. Heating reactions in the saline giant and associated deposits can also supply elevated levels of the greenhouse gases CO2 and CH4. Explosive volcanism tied to the emplacement of LIPs in the region of a saline giant or the atmosphere-scale disturbance linked to the impact of a large bolide in an area underlain by a saline giant are efficient mechanisms to move large volumes of halocarbons, sulphurous volatiles and greenhouse gasses to the troposphere. The third article in this series will document the specific evaporite geology that contributed to four of the five major Phanerozoic extinction events (Figure 10).

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    Volatile evaporite interactions with magma Part 1 of 3: Indications of hydrated salts?

    John Warren - Sunday, February 10, 2019

    Introduction

    Direct and indirect interactions between magma and evaporites at a regional scale are neither well documented, nor well understood. Mostly, this is because little or no salt remains once the high-temperature interactions have run their course; instead, there is a suite of indirect geochemical and indicator-mineral assemblages (Warren, 2016). Aside from the presence of what can be ambiguous indicator suites, some hard-rock geologists with a career working in igneous and metamorphic terranes may not be well versed in textures indicative of the former presence of sedimentary evaporites, nor their varying volatility, nor their meta-evaporitic and meta-igneous siblings.

    The term pyrometasomatic encompasses some, but not all, of the types of salt-magma interaction and reactions that occur when evaporites and molten magmas of different types are nearby. Styles of evaporites interactions with magma are a spectrum, with two endmember situations; 1) orthomagmatic (salt-assimilative and internal to the magma), and 2) paramagmatic (salt-interactive and external to the magma). Both encompass outcomes that can include a variety of substantial ore deposits (Warren, 2016; Chapter 16). Only in situations where igneous sills and dykes have intruded salt masses, with contacts preserved, can direct effects of magma-salt interaction be documented. Even then, determining the timing of the evaporite igneous interaction can be problematic; one must ask if the chemistry and texture indicate, 1) syn-igneous emplacement, or 2) post-emplacement alteration and deeply circulating groundwater flushing, or 3) a combination.

    Orthomagmatic and paramagmatic evaporite associations are distinct from occurrences of primary igneous/magmatic anhydrites, which precipitate from sulphate-saturated melts. Igneous anhydrite forms independently of any sedimentary evaporite assimilation, as seen, for example, in anhydrite crystals crystallised in trachyandesitic pumice erupted from El Chichón Volcano in 1982, or in dacitic pumices erupted from Mount Pinatubo in 1991 and in acidic lavas in the Yanacocha district of northern Peru (Luhr et al., 2008; Chambefort et al., 2008). These evaporite assimilations are also distinct from fumarolic anhydrite, which precipitates where groundwaters and sulphur-bearing magmatic fluids interact, as on Usu Volcano, Hokkaido, and many central American and Andean volcanoes such as El Laco (Zimbelman et al., 2005). Likewise, they are distinct from the anhydrite precipitates (white smokers) in and below submarine vents across numerous mid-oceanic ridges (Humphris et al., 1995). See Warren 2016 (Chapter 16) for more geological detail on these non-evaporite-igneous anhydrite occurrences.

    Cooking with salt (thermal decomposition of hydrated versus non-hydrated salts)

    Perhaps the most critical factor controlling the local intensity of magmatic interaction with an evaporitic country rock is whether or not the sedimentary evaporite assemblage, in proximity to an igneous heat source, contains abundant hydrated salts, such as gypsum, polyhalite or carnallite. Hydrated evaporite salts, when interacting with the igneous realm, are highly volatile and likely to decompose. They tend to release their water of crystallisation at temperatures many hundreds of degrees below the melting points of their anhydrous counterparts (Table 1).


    In contrast, anhydrous salts, such as halite beds intruded by igneous dykes or sills, are much less reactive. At a local scale (measured in metres) with respect to an intrasalt-igneous interaction, there are a number of documented thermally-driven alteration styles, typically created by the intrusion of dolerite dykes and sills into cooler halite, or the outflow of extrusive igneous flows over cooler halite beds (Knipping and Herrmann, 1985; Knipping, 1989; Grishina et al, 1992, 1998; Gutsche, 1988; Steinmann et al., 1999; Wall et al., 2010). Hot igneous material interacts with somewhat cooler anhydrous salt masses, typically halite or anhydrite, to create narrow but distinct heat and mobile fluid-release envelopes(Figure 1), also reflected in the resulting recrystallised inclusion-modified salt textures (Figure 2).


    Based on studies of inclusion chemistry and homogenization temperatures in fluid inclusions in bedded halite near intrusives, it seems that the extent of the influence of a dolerite sill or dyke in bedded salt is marked by fluid (brine)-inclusion migration. This is evidenced by the disappearance of chevron structures and consequent formation of clear (sparry) recrystallised halite, with a new set of higher-temperature brine inclusions located at intercrystal or polyhedral intersections. Such a migration envelope is documented in bedded Cambrian halites intruded by end-Permian dolerite dykes in the Tunguska region of Siberia (Figure 2; Grishina et al., 1992). There, as a rule of thumb, an alteration halo extends up to twice the thickness of the dolerite sill above the sill and almost the thickness of the sill below (Figure 1).

    Four inclusion type associations were found in halite as a function of the ratio of the distance of the sample from the intrusion contact (d) to the thickness of the intrusion (h), i.e. d/h (Figure 2). Chevron structures with aqueous inclusions progressively disappear as d/h decreases; the disappearance of chevrons occurs at greater distances above than below the intrusive sill. At d/h < 5 above the sill, a low-density CO2 vapour phase appears in brine inclusions, at d/h < 2 H2S-bearing liquid-CO2 inclusions appear, sometimes associated with carbonaceous material and orthorhombic S8, and for d/h < 0.9, CaCl2, CaCl2.KCl and nCaCl2.n MgCl2 solids occur in association with free water and liquid CO2 inclusions, with H2S, SCO, and Sg. The d/h values marking the transitions outlined above are lower below the sills than above. The water content of the inclusions progressively decreases on approaching the sills, whereas their CO2 content and density increase. Carnallite, sylvite and calcium chloride can occur as solid inclusions in the two associations nearest to the sill for d/h<2. Carnallite and sylvite occur as daughter minerals in brine inclusions. The presence of carbon dioxide is taken to indicate fluid circulation and dissolution/recrystallisation phenomena induced by the basalt intrusions. The origin of carbon dioxide is likely related to carbonate dissolution during magmatism (see Salty Matters, Oct 31, 2016).


    In some shallow locations, relatively rapid magma emplacement can lead to linear breakout trends outlined by phreatomagmatic or phreatic explosion craters. Such phreatic explosion craters have been imaged on the Tertiary seafloor horizons in parts of the North Sea (Figure 3; Wall et al., 2010). The dykes were emplaced into Paleozoic and Mesozoic sediments and have a common upper termination in Early Tertiary sediments. The dykes are part of the British Tertiary volcanic province emplaced some 58 Ma. These dykes are characterised by a narrow 0.5–2 km wide vertical disturbance of seismic reflections that have linear plan view geometry. Negative magnetic anomalies directly align with the vertical seismic disturbance zones and indicate the presence of underlying igneous material. Linear coalesced collapse craters are found above the dykes. The collapse craters formed above the dyke due to the release of volatiles at the dyke tip and resulting gaseous expansion and subsequent volume loss. According to Wall et al. (2010), the larger craters likely formed due to explosive phreatomagmatic interaction between magma and pore water. The linearly aligned collapse craters can be considered an Earth analogue to Martian pit chain craters.

    A phreatic eruption, also called a phreatic explosion, ultravulcanian eruption or steam-blast eruption, occurs when magma heats ground or surface water and is a separate but related occurrence to a phreatomagmatic eruption. A phreatomagmatic deposit typically contains solid inclusions of magmatic (igneous) material, whereas debris tied to a phreatic deposit does not, but ties to the effects of juvenile and deeply circulated hydrothermal waters. Extreme temperatures associated with an emplaced magma (anywhere from 500 to 1,170 °C) can cause a near-instantaneous phase change to steam, so forming a phreatomagmatic deposit. That is, rapid heating results in an intense explosion made up of steam, water, ash, rock, and volcanic bombs. During the eruption of Mount St. Helens, hundreds of steam explosions preceded the1980 Plinian eruption of the volcano core. Many authors argue a less intense geothermal event results in a mud volcano, but there are many other active mud volcanoes worldwide that tie to compactional overpressure unrelated to any magma emplacement (Warren et al., 2011). As the published interpretation of aligned phreatic breakout structures illustrated in Figure 3 is based on seismic without well control, the explosion mechanism may be solely phreatic heating or phreatomagmatic.


    Deposits of phreatic eruptions (as contrasted with a phreatomagmatic eruption) typically include steam and rock fragments without the inclusion of fragments derived from liquid magma, lava or volcanic ash. The temperature of the phreatic fragments can range from cool to incandescent. So if molten magma is present, the resulting explosive debris deposit is typically classified as a phreatomagmatic eruption. These eruptions can create broad, low-relief craters called maars. In contrast, phreatic explosions lack debris derived from molten (igneous) material, but emplacement can be accompanied by carbon dioxide or hydrogen sulfide gas emissions. CO2 can asphyxiate at sufficient concentration; H2S is a broad spectrum poison. A 1979 phreatic eruption on the island of Java killed 140 people, most of whom were overcome by poisonous gases. Phreatic eruptions, even if the deposit lacks igneous rock fragments, are typically classed as a type of volcanic eruptions because a phreatic eruption can force juvenile fluids to the surface. But when a phreatic explosion is related to an igneous feature intersecting an evaporite bed, the resultant textures show a contrast between heating of anhydrous and hydrous salts


    Hydrous salt interactions in Germany

    Textures created by an igneous intrusion into a variably-hydrated evaporite succession can be studied in the dyke-and sill-intruded halite levels exposed in the walls of potash mines of the Werra-Fulda district of Germany (Figure 4; Steinmann et al., 1999; Schofield et al., 2014). There, the Permian Zechstein salt series contains two important potash salt horizons (2-10m thick), which are mined at a depths ≈ 800 m, from within a 400m thick halite host (Figure 4a). In the later Tertiary, basaltic melts intruded these Zechstein evaporites, but it seems only a few dykes reached the Miocene landsurface. The basaltic melt ties to regional volcanic activity, some 10 to 25 Ma. Basalts exposed in the halite-dominant portions of the mine walls are typically subvertical dykes, rather than sills. The basaltic intervals intersect the salt over zones up to several kilometres wide (Figure 4b). However, correlations of individual dyke swarms, either between different mines, or between surface and subsurface outcrops is difficult.


    From a paleogeographic perspective, the Werra-Fulda Basin is situated in a southern embayment of the European Zechstein Basin. It contains cyclic evaporites of the Werra Formation (Z1). In the Neuhof area, the evaporites of the Zechstein are underlain by siliciclastic rocks of the Permian Rotliegend interval. The higher Zechstein-cycles (Z2 – Z7), on top of the Werra Formation, consist of a siliciclastic succession with intercalated limestone and anhydrite layers (Strauch et al., 2018; Beer and Barnasch, 2018). The Werra Formation is dominated by rock salt with a thickness up to 300 m.

    Two potash seams (Seam Hessen and Seam Thüringen) separate the rock salt of the Werra Formation into three distinct units (Figure 4b). Lower, Middle and Upper Werra rock salt). Seam Hessen mainly consists of hard salt (kieserite, sylvite, halite and anhydrite). It is overlain by several, potash mineral-bearing horizons which show a strong vertical and lateral heterogeneity and consist of kieserite, sylvite, carnallite, halite and anhydrite. Internally, three separate units are identified within the potash Seam Hessen (Figure 4). The “Wurmsalz”, a hard salt with up to four strongly folded anhydritic clay bands represents the lower part of Seam Hessen. The middle part consists of massive, kieserite-rich hard salt with abundant sylvite lenses (“Flockensalz”). The “Bändersalz”, a banded hard salt which is typically intercalated with brownish, halitic layers occurs in the upper part of Seam Hessen. Potash Seam Thüringen usually occurs around 50 m below Seam Hessen. Its lower part is dominated by a well-bedded hard salt with intercalated rock salt. Its upper part consists of a variety of rock types including carnallite, sylvite and hard salt.

    In the Fulda region the thermally-driven release of water of crystallisation within particular Zechstein salt beds intersecting igneous dykes creates thixotropic or subsurface “peperite” textures in hydrated carnallitite ore layers, where heated water of crystallisation escaped from the hydrated-salt lattice. Dehydration-driven loss of mechanical strength focuses zones of magma entry into particular horizons in the salt mass, wherever hydrated salt layers were intersected (Figure 5b verses 5c). In contrast, dyke and sill margins are much sharper and narrower in zones of contact with anhydrous salt intervals (Figure 5a, d; Schofield et al., 2014).

    Accordingly, away from the immediate vicinity of the direct thermal aureole, heated and overpressured dehydration waters can enter a former Zechstein carnallite halite bed, and drive the creation of extensive soft sediment deformation and peperite textures in the previous hydrated layer (Figure 5c, d). Mineralogically, sylvite and coarse recrystallised halite dominate the salt fraction in the peperite intervals/beds. These deformed beds formed within a hydrated salt bed and so differ from the conventional notion of volcanic peperites indicating water-saturated sediment interactions with very shallow dyke or sill emplacements.

    Sylvite in these altered zones is a form of dehydrated carnallite, not a primary-textured salt. In the Fulda region, such altered zones and deformed units can extend along former carnallite layers to tens or even a hundred or more metres from the dyke feeder. Ultimately, the deformed potash bed passes laterally out into the unaltered bed, which can retain abundant inclusion-rich primary chevron halite and carnallite (Figure 5c versus 5d). That is, nearer the basalt dyke, the carnallite is transformed mainly into inclusion-poor halite and sylvite, the result of recrystallisation combined with incongruent flushing of warm saline fluids mobilised from the hydrated carnallite crystal lattice as it was heated and decomposed in response to nearby dyke emplacement. During such Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-rich diagenetic and juvenile fluids were mixed with fluids originating from thermally-mobilised crystallisation water in the carnallite as it converted to sylvite.

    Nearer the basalt dyke, the carnallite is largely transformed into inclusion-poor halite and sylvite, the result of incongruent flushing of warm saline fluids mobilised from the hydrated carnallite crystal lattice as it was heated by dyke emplacement. During Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-fluids were mixed with fluids originating from thermally-mobilised crystallisation water in the carnallite, as it converted to sylvite. This brine mixture altered the basalts during post-intrusive cooling, an event which numerical models suggest was quite rapid (Knipping, 1989): a dyke of less than 0.5 m thickness probably cooled to temperatures less than 200°C within 14 days of dyke emplacement.

    Worldwide, igneous dykes intersecting salt beds tend to widen to become sills in two zones: 1) along evaporite units within the halite mass that contain hydrated salts, such as carnallite or gypsum (Figure 5b, c) and, 2) where rising magma has ponded and so created laccoliths at the upper or lower halite contact with the adjacent nonsalt strata, or against a salt wall (Warren, 2016). The first alteration of the hydrated salt layer is a form of mineral alteration and recrystallisation in response to a pulse of released water/steam as dyke-driven heating forces the dehydration of hydrated salt layers. The second alteration is often folding and fluid-like disaggregation of the former, now dehydrated, layer in response to the mechanical strength contrast at a hydrated-nonhydrated salt-bed contact (Warren, 2016).

    Surface expression of hydrated bedded salts interacting with magma in Dallol, Ethiopia

    Local potash ores typify thermal sump depressions in the Dallol and Musley areas (Figure 6a, b, c, 7) where a similar set of subsurface destabilisation processes occurred when rising magma reached the levels of hydrated salts (kainite and carnallite beds) in the Houston Formation of the Danakhil depression fill (see Warren 2016 and Bastow et al. 2018 for more detailed discussion of the potash stratigraphy). To attain these hydrated salt levels the rising dyke swarm had passed relatively passively through the Lower Rocksalt Formation (Salty Matters, April 29, 2015). Emplacement of the magma/dykes into hydrated evaporites below the vicinity of what is now the Dallol Mound would have mobilised and deformed the hydrated potash salt level, converting carnallite to sylvite, kainite to bischofite and lesser kieserite, as well as creating widespread cavities filled with rising pressured volatiles carried by MgCl and KCl brines. Pressurisacreation of a cavernous network filled with volatiles at the level of the Houston Formation would have aided in forming the four-way dip closure now seen on the exposed and eroding salt beds that make up much of the Dallol Mound surface.


    Once these hydrothermal cavities dissolved and breached the way to surface, the feeder brines cool and precipitate prograde salts such as halite, sylvite and bischofite. Such destabilisation has likely accommodated the emplacement of a basaltic sill at the level of the potash salts, in turn driving the uplift of the lake beds above this region outlined by the centripetal dips of the Dallol Mound. Mound-related uplift and hydrothermal activity then stimulate the formation of natural areas of ground collapse, sulphurous and acidic springs and fumaroles, along with the creation of water-filled chimneys and doline sags, filling with various hydrothermal salts, in the vicinity of the volcanic mound  (Figure 6).

    That is this type of potash in the Dallol Mound region is hydrothermally reworked from the uplifted equivalents of the Houston Formation. Even today this hydrology is precipitating carnallitite (associated with bischofite and minor kieserite) in various hydrothermal brine pools atop and around the Dallol Mound, such as the carnallite-dominant Crescent deposit (Figure 7). These hydrothermal salts owe their origins to daylighting of pressurised fluid systems and cavities.


    The last pressurised phreatic explosion crater formed in 1926. They were created by the volatile products of hydrated salt layers (Houston Fm) where these salts had come into contact with thermal aureoles or actual lithologies of newly emplaced dykes that had penetrated the underlying halite section. Volcanic rock fragments and other igneous debris have yet to make it to the surface in the Dallol Mound region, although active volcanic mounds and flows do cover the saltflat surface tens of kilometres to the south (Erte Alle ) and north. Based on the analogy exposed within the Zechstein-hosted potash mines of the Fulda region of Germany, it is likely that as well as creating at-surface brine pools, this hydrothermal dyke-related hydrology locally converts most subsurface carnallitite to a disturbed sylvinite bed at the level of contact with the Houston Fm.

    Implications

    It seems a "one-size-fits-all" model does not characterise magmatic interactions with massively bedded evaporites. Instead, there is a mineralogical control to the intensity of the interaction and the depth of thermal influence of recrystallisation and mobilisation textures. When a dyke-swarm intersects halite or anhydrite the thermally-driven recrystallisation and fluid migration halo is more limited, as outlined in Figure 1 and Figure 5a, d.

    In contrast, when a dyke swarm intersects an interval containing hydrous salts such kainite, carnallite or gypsum, the heating drives the expulsion of the bound-water at decomposition temperatures much lower than the salts melting point (Table 1). Such hydrous-salt intervals devolatise, fluidise and flow, with the effects of the heating halo extending much further away from the heat source, driven in part by steam-driven hydrofracturing. On cooling, the resulting mineralogy in the highly-deformed bed is dominated by the anhydrous form of the devolatised salt, as in the sylvite unit after carnallite as seen in potash seams adjacent to dykes in the Fulda Region (Figures 5b, c).

    Closer in to the heat source, the basalt that has moved in along the hydrous potash beds show abundant peperite textures (Figure 5c; Schofield et al., 2014). Actually, this is a unique form of peperite that is tied to beds of hydrous evaporite. It forms outside the usual scenario envisaged for peperite whereby molten igneous material interacts with wet sediment, with the water in the wet sediment held in interparticle pores.

    The classic definition of a peperite is that it is a "genetic term applied to a rock formed essentially in situ by disintegration of magma intruding and mingling with unconsolidated or poorly consolidated, typically wet sediments. The term also refers to similar mixtures generated by the same processes operating at the contacts of lavas and other hot volcaniclastic deposits with such sediments" (Skilling et al. 2002).

    In the case of the bedded hydrous salt intervals, before the intrusion of the igneous heat source, there was little to no free water, other than occasional brine inclusions in associated halite chevrons. What makes these hydrous-salt peperites interesting is that it is the igneous heating drives a mineralogic transformation in the hydrous salts that makes the formerly "dry" salt bed become "wet" sediment.

    Before our work in the Fulda region (Schofield et al., 2014), the nature of igneous interactions with evaporites was understood to be mainly that documented by studies in areas with intrusives interacting with thick anhydrous halite and anhydrite beds. The heating haloes were seen as driving recrystallisation and brine migration over limited lateral distances of a few metres. However, the potash seam interactions in the Fulda region show this alteration distance can be much greater (hundreds of metres) id hydrous salt layers are heated.

    The surface geology in the Dallol Mound region of Ethiopia shows an even more impressive set of igneous dyke hydrated salt interactions (Warren, 2016). There the potash interval known as the Houston Formation is a tens-of-metres thick section of hydrated salts below the upper halite unit and atop the lower halite. When the rising igneous dyke swarm rose to the level of Houston Formation, it drove a broad linear devolatisation zone in the dyke-heated alteration halo. This, in turn, forced the formation of the closed anticlinal uplift structure that is the Dallol mound. The release of MgCl2 during volatisation also explains phreatic breakout features that are outlined by at-surface collapse dolines with their hot (104-108°C) brine lakes and unusual bischofite (MgCl2) precipitates. Likewise, the same set of processes explains the occurrences of metres to tens of metres thick bischofite intervals that are intersected in cores in some of the potash exploration wells in the vicinity of Dallol Mound (pers. obs). These are likely cavity fill deposits formed as a byproduct of kainite and carnallite devolatisation sourced at the level of Houston Formation.

    This set of more mobile brine fluid escape features has implications for nuclear waste storage in halite successions where a storage cavity may be in proximity to an interval of hydrous evaporite salts. Halite-hosted purpose-built caverns in thick evaporite intervals are one of the safest places in the world to store waste but perhaps not in parts of the salt succession that entrain beds of hydrous salts such as carnallite or kainite (Warren, 2017).

    References

    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.

    Beer, W., and L. Barnasch, in press, Werra-Fulda-Becken, SDGG- Monography.

    Chambefort, I., J. H. Dilles, and A. J. R. Kent, 2008, Anhydrite-bearing andesite and dacite as a source for sulfur in magmatic-hydrothermal mineral deposits: Geology, v. 36, p. 719-722.

    Grishina, S., J. Dubessy, A. Kontorovich, and J. Pironon, 1992, Inclusions in salt beds resulting from thermal metamorphism by dolerite sills (eastern Siberia, Russia): European Journal of Mineralogy, v. 4, p. 1187-1202.

    Grishina, S., J. Pironon, M. Mazurov, S. Goryainov, A. Pustilnikov, G. Fonderflaas, and A. Guerci, 1998, Organic inclusions in salt - Part 3 - Oil and gas inclusions in Cambrian evaporite deposits from east Siberia - A contribution to the understanding of nitrogen generation in evaporite: Organic Geochemistry, v. 28, p. 297-310.

    Gutsche, A., 1988, Mineralreaktionen und Stotransporte an einem Kontakt Basalt-Hartsalz in der Werra-Folge des Werkes Hattorf: Unpubl. diploma thesis, thesis, Georg-August-Universita, Gottingen.

    Humphris, S. E., P. M. Herzig, D. J. Miller, J. C. Alt, K. Becker, D. Brown, G. Brugmann, H. Chiba, Y. Fouquet, J. B. Gemmell, G. G., M. D. Hannington, N. G. Holm, J. J. Honnorez, G. J. Iturrino, R. Knott, R. Ludwig, K. Nakamura, S. Petersen, A. L. Reysenbach, P. A. Rona, S. Smith, A. A. Sturz, M. K. Tivey, and X. Zhao, 1995, The internal structure of an active sea-floor massive sulphide deposit: Nature, v. 377, p. 713-716.

    Knipping, B., 1989, Basalt intrusions in evaporites: Lecture Notes in Earth Sciences (Springer-Verlag), v. 24, p. 132 pp.

    Knipping, B., and A. G. Hermann, 1985, Mineralreaktionen und Stoff transporte an einem Kontakt Basalt-Carnallitit im Kalisalzhorizont Thüringen der Werra-Serie des Zechsteins: Kali und Steinsalz, v. 9, p. 111-124.

    Luhr, J. F., 2008, Primary igneous anhydrite: Progress since its recognition in the 1982 El ChichÛn trachyandesite: Journal of Volcanology and Geothermal Research, v. 175, p. 394-407.

    Schofield, N., I. Alsop, J. Warren, J. R. Underhill, R. Lehné, W. Beer, and V. Lukas, 2014, Mobilizing salt: Magma-salt interactions: Geology, v. 42, p. 599-602.

    Skilling, I. P., J. D. L. White, and J. McPhie, 2002, Peperite: a review of magma–sediment mingling: Journal of Volcanology and Geothermal Research, v. 114, p. 1-17.

    Steinmann, M., P. Stille, W. Bernotat, and B. Knipping, 1999, The corrosion of basaltic dykes in evaporites: Ar-Sr-Nd isotope and rare earth elements evidence: Chemical Geology, v. 153, p. 259-279.

    Strauch, B., M. Zimmer, A. Zirkler, S. Höntzsch, and A. M. Schleicher, 2018, The influence of gas and humidity on the mineralogy of various salt compositions – implications for natural and technical caverns: Advances in Geoscience, v. 45, p. 227-233.

    Wall, M., J. Cartwright, R. Davies, and A. McGrandle, 2010, 3D seismic imaging of a Tertiary Dyke Swarm in the Southern North Sea, UK: Basin Research, v. 22, p. 181-194.

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

    Warren, J. K., 2017, Salt usually seals, but sometimes leaks: Implications for mine and cavern stabilities in the short and long term: Earth-Science Reviews, v. 165, p. 302-341.

    Warren, J. K., A. Cheung, and I. Cartwright, 2011, Organic Geochemical, Isotopic and Seismic Indicators of Fluid Flow in Pressurized Growth Anticlines and Mud Volcanoes in Modern Deepwater Slope and Rise Sediments of Offshore Brunei Darussalam; Implications for hydrocarbon exploration in other mud and salt diapir provinces (Chapter 10), in L. J. Wood, ed., Shale Tectonics, v. 93: Tulsa OK, AAPG Memoir 93 (Proceedings of Hedberg Conference), p. 163-196.

    Zimbelman, D. R., R. O. Rye, and G. N. Breit, 2005, Origin of secondary sulfate minerals on active andesitic stratovolcanoes: Chemical Geology, v. 215, p. 37-60.

     

    Brine evolution and origins of potash - primary or secondary. Ancient potash ores: Part 3 of 3

    John Warren - Monday, December 31, 2018

    Introduction

    In the previous two articles in this series on potash exploitation, we looked at the production of either MOP or SOP from anthropogenic brine pans in modern saline lake settings. Crystals of interest formed in solar evaporation pans and came out of solution as: 1) Rafts at the air-brine interface, 2) Bottom nucleates or, 3) Syndepositional cements precipitated within a few centimetres of the depositional surface. In most cases, periods of more intense precipitation tended to occur during times of brine cooling, either diurnally or seasonally (sylvite, carnallite and halite are prograde salts). All anthropogenic saline pan deposits examples can be considered as primary precipitates with chemistries tied to surface or very nearsurface brine chemistry.

    In contrast, this article discusses ancient potash deposits where the chemistries and ore textures are responding to ongoing alteration processes in the diagenetic realm. Unlike the modern brine pans where brines chemistries and harvested mineralogies are controllable, at least in part, these ancient deposits show ore purities and distributions related to ongoing natural-process overprints.



    Table 1 lists some modern and ancient potash deposits and prospects by dividing them into Neogene and Pre-Neogene deposits (listing is extracted and compiled from SaltWork® database Version 1.7). The Neogene deposits are associated with a time of MgSO4-enriched seawaters while a majority of the Pre-Neogene deposits straddle times of MgSO4 enrichment and depletion in the ocean waters.

    Incongruent dissolution in burial

    Many primary evaporite salts dissolve congruently in the diagenetic realm; i.e., the composition of the solid and the dissolved solute stoichiometrically match, and the dissolving salt goes entirely into solution (Figure 1a). This situation describes the typical subsurface dissolution of anhydrous evaporite salts such as halite or sylvite. However, some evaporite salts, typically hydrated salts, such as gypsum or carnallite, dissolve incongruently in the diagenetic realm, whereby the composition of the solute in solution does not match that of the solid (Figure 1b). This solubilisation or mineralogical alteration is defined by the transformation of the "primary solid" into a secondary solid phase, typically an anhydrous salt, and the loss of water formerly held in the lattice structure. The resulting solution generally carries ions away in solution.


    More than a century ago, van't Hoff (1912) suggested that much subsurface sylvite is the result of incongruent solution of carnallite yielding sylvite and a Mg-rich solution. According to Braitsch (1971, p. 120), the incongruent alteration (dissolution) of carnallite is perhaps the most crucial process in the alteration of subsurface potash salts and the formation of diagenetic (secondary) sylvite.

    Widespread burial-driven incongruent evaporite reactions in the diagenetic realm include the burial transition of gypsum to anhydrite (reaction 1)

    CaSO4.2H2O --> CaSO4 + 2 H2O ... (1)

    and the in-situ conversion of carnallite to sylvite via the loss of magnesium chloride in solution (reaction 2)

    KMgCl3.6H2O --> KCl + Mg++ + 2Cl- + 6H2O ...(2)

    Typically, a new solid mineral remains, and the related complex solubility equilibrium creates a saline pore water that may, in turn, drive further alteration or dissolution as it leaves the reaction site (Warren, Chapters 2 and 8). Specifically for ancient potash, reaction 2 generates magnesium and chloride in solution and has been used to explain why diagenetic bischofite and dolomite can be found in proximity to newly formed subsurface sylvite. Bischofite is a highly soluble salt and so is metastable in many subsurface settings where incongruent dissolution is deemed to have occurred, including bischofite thermal pool deposits in the Dallol sump in the Danakhil of Ethiopia (Salty Matters, May 1, 2015). In many hydrologically active systems, solid-state bischofite is flushed by ongoing brine crossflow and so help drive the formation of various burial dolomites. Only at high concentrations of MgCl2 can carnallite dissolve without decomposition.

    Laboratory determinations

    In the lab, the decomposition of carnallite in an undersaturated aqueous solution is a well-documented example of incongruent dissolution (Emons and Voigt, 1981; Xia et al., 1993; Hong et al., 1994; Liu et al., 2007; Cheng et al., 2015, 2016). When undersaturated water comes into contact with carnallite, the rhombic carnallite crystals dissolve and, because of the common ion effect, small cubic KCl crystals form in the vicinity of the dissolving carnallite. As time passes, the KCl crystals grow into larger sparry subhedral forms and the carnallite disappears.

    Carnallite’s crystal structure is built of Mg(H2O)6 octahedra, with the K+ ions are situated in the holes of chloride ion packing meshworks, with a structural configuration similar to perovskite lattice types (Voigt, 2015). Potassium in the carnallite lattice can be substituted by other large single-valence ions like NH4+, Rb+, Cs+ or Li(H2O)+, (H3O)+ and Cl- by Br- and I-. These substitutions change the lattice symmetry from orthorhombic in the original carnallite to monoclinic.

    When interpreting the genesis of ancient potash deposits and solutions, the elemental segregation in the lattice means trace element contents of bromide, rubidium and caesium in primary carnallite versus sylvite daughter crystals from incongruent dissolution can provide valuable information. For example, in a study by Wardlaw, (1968), a trace element model was developed for sylvite derived from carnallite that gave for Br and Rb concentration ranges of 0.10–0.90 mg/g and 0.01–0.18 mg/g, respectively. In a later study of sylvite derived by fresh-water leaching of magnesium chloride under isothermal conditions at 25 °C. Cheng et al. (2016), defined a model whereby primary sylvites precipitated from MgSO4-deficient sea water, gave Br and Rb concentration ranges of 2.89–3.54 mg/g and 0.017–0.02 mg/g, respectively (no evaporation occurred at saturation with KCl). In general, they concluded sylvite derived incongruently from carnallite would contain less Br and more Rb than primary sylvite (Figure 2; Cheng et al., 2016).


    Subsurface examples

    The burial-driven mechanism widely cited to explain the incongruent formation of sylvite from carnallite is illustrated in Figure 3 (Koehler et al., 1990). Carnallite precipitating from evaporating seawater at time 1 forms from a solution at 30°C and atmospheric (1 bar) pressure, and so plots as point A, which lies within the carnallite stability field (that is, it sits above the dashed light brown line). With subsequent burial, the pressure increases so that the line defining carnallite-sylvite boundary (solid dark brown line) moves to higher values of K. By time 2, when the pressure is at 1 Kbar (corresponds to a lithostatic load equivalent to 2-3 km depth), the buried carnallite is thermodynamically unstable and so is converting to sylvite + solution (as the plot field now lies in the sylvite + solution field (Figure 3). If equilibrium is maintained the carnallite reacts incongruently to form further sylvite and MgCl2-solution. Thus, provided the temperature does not rise substantially, increasing pressure as a result of burial will favour the breakdown of carnallite to sylvite. However, as burial proceeds, the temperature may become high enough to favour once again the formation of carnallite from sylvite + solution (that is the solution plot point from A moves toward the right-hand side of the figure and back into the carnallite stability field).


    Sylvite, interpreted to have formed from incongruent dissolution of primary carnallite, is reported from the Late Permian Zechstein Formation of Germany (Borchert and Muir, 1964), Late Permian Salado Formation of New Mexico (Adams, 1970), Early Mississippian Windsor Group of Nova Scotia (Evans, 1970), Early Cretaceous Muribeca Formation of Brazil and its equivalents in the Gabon Basin, West Africa (Wardlaw, 1972a, b; Wardlaw and Nicholls, 1972; Szatmari et al., 1979; de Ruiter, 1979), Late Cretaceous of the Maha Sarakham Formation, Khorat Plateau, Thailand and Laos (Hite and Japakasetr, 1979), Pleistocene Houston Formation, Danakil Depression, Ethiopia (Holwerda and Hutchinson, 1968), and Middle Devonian Prairie Formation of western Canada (Schwerdtner, 1964; Wardlaw, 1968) (See Table 1).

    This well-documented literature base supports a long-held notion that there is a problem with sylvite as a primary (first precipitate) marine bittern salt, especially if the mother seawater had ionic proportions similar to those present in modern seawater (see Lowenstein and Spencer, 1990 for an excellent, if 30-year-old, review). We know from numerous evaporation experiments, that sylvite does not crystallise during the evaporation of modern seawater at 25°C, except under metastable equilibrium conditions (Braitsch, 1971; Valyashko, 1972; Hardie, 1984). The sequence of bitterns crystallising from modern seawater bitterns was illustrated in the previous Salty Matters article in this series (see Figure 1 in October 31 2018).

    Across the literature documenting sylvite-carnallite associations in ancient evaporites, the dilemma of primary versus secondary sylvite is generally solved in one of three ways. Historically, many workers interpreted widespread sylvite as a diagenetic mineral formed by the incongruent dissolution of carnallite (Explanation 1). Then there is the interpretation that some sylvite beds, perhaps associated with tachyhydrite, were precipitated in the evaporite bittern part of a basin hydrology that was fed by CaCl2-rich basinal hydrothermal waters (Explanation 2: see Hardie, 1990 for a good discussion    of this mechanism). Then there is the third, and increasingly popular explanation of primary or syndepositional sylvite at particular times in the chemical evolution of the world oceans (MgSO4-depleted oceans).


    Changes in the relative proportions of magnesium, sulphur and calcium in the world’s oceans are well supported by brine inclusion chemistry of co-associated chevron halite (Figure 4). Clearly, there are vast swathes of times in the earth’s past when the chemistry of seawater changed so that MgSO4 levels were lower than today and it was possible that sylvite was a primary marine bittern precipitate (see Lowenstein et al., 2014 for an excellent summary).

    In my opinion, there is good evidence that all three explanations are valid within their relevant geological contexts but, if used exclusively to explain the presence of ancient sylvite, the argument becomes somewhat dogmatic. I would say that that, owing to its high solubility, the various textures and mineralogical associations of carnallite/sylvite and sulphate bitterns found in ancient potash ore beds reflect various and evolving origins. Ambient textures and mineralogies are dependent on how many times and how pervasively in a potash sequence’s geological burial history an evolving and reactive pore brine chemistry came into contact with parts or all of the extent of highly reactive potash beds (Warren, 2000; 2010; 2016).

    In my experience, very few ancient examples of economic potash show layered textures indicating primary precipitation on a brine lake floor, instead, most ancient sylvite ores show evidence of at least one episode of alteration. That is, various forms and textures in potash may dissolve, recrystallise and backreact with each other from the time a potash salt is first precipitated until it is extracted. The observed textural and mineralogical evolution of a potash ore association depends on how open was the hydrology of the potash system at various stages during its burial evolution. The alteration can occur syndepositionally, in brine reflux, or later during flushing by compactional or thermobaric subsurface waters or during re-equilibration tied to uplift and telogenesis. Tectonism (extensional and compactional) during the various stages of a basin’s burial evolution acts as a bellows driving fluid flow within a basin, so forcing and speeding up the focused circulation of potash-altering waters.

     

    A similar, but somewhat less intense, textural evolution tied to incongruent alteration is seen in the burial history of other variably hydrated evaporite salts. For example, CaSO4 can flip-flop from gypsum to anhydrite and back again depending on temperature, pore fluid salinity and the state of uplift/burial. Likewise, with the more complicated double salt polyhalite, there are mineralogical changes related to whether it formed in a MgSO4 enriched or depleted world ocean and the associated chemistry of the syndepositional reflux brines across extensive evaporite platforms (for a more detailed discussion of polyhalite see Salty Matters, July 31, 2018). Kainite-kieserite-carnallite also show evidence of ongoing incongruent interactions. This means that, as in gypsum/anhydrite/polyhalite or kain ite/kieserite sequences, there will be primary and secondary forms of both carnallite and sylvite that can alternate during deposition, during burial and any deep meteoric flushing and then again with uplift. In Quaternary brine factories these same incongruent chemical relationships are what facilitate the production of MOP (sylvite) from a carnallitite feed or SOP from kainite/kieserite/schoenite feed (see articles 1 and 2 in this series).

     

    To document the three end-members of ancient sylvite-carnallite decomposition/precipitation we will look at three examples; 1) Oligocene potash in the Mulhouse Basin where primary sylvite textures are commonplace, 2) Devonian potash ores in western Canada, where multiple secondary stages of alteration are seen, and 3) Igneous-dyke associated sylvite in east Germany where thermally-driven volatisation (incongruent melting) forms sylvite from dehydrated carnallite.


    Oligocene Potash, Mulhouse Basin France

    Moving backwards into deep time, this 34 Ma deposit contains some of the first indications of well preserved primary marine-fed sylvinite (MOP) textures exemplified by laterally-continuous mm-scale alternations of potash and halite layers and lamina (Figure 5a-c). Interestingly, all solid-state potash deposits laid down in the post-Oligocene period contain increased proportions of MgSO4 salts, making them much more difficult to economically mine and process (see Table 1 and Salty Matters, May 12, 2015)

    From 1904 until 2002, potash was conventionally mined in France from the Mulhouse Basin (near Alsace, France). With an area of 400 km2, the Mulhouse Basin is the southernmost of a number of Lower Oligocene evaporite basins that occupied the upper Rhine Graben, which at that time was a narrow adiabatic-arid rift valley (Figure 6a). The graben was a consequence of the collision between European and African plates during the Paleogene. It is part of a larger intracontinental rift system across Western Europe that extended from the North Sea to the Mediterranean Sea, stretching some 300 km from Frankfurt (Germany) in the north, to Basel (Switzerland) in the south, with an average width of 35 km (Cendon et al., 2008). The southern extent of the graben is limited by a system of faults that place Hercynian massifs and Triassic materials into contact with the Paleogene filling. Across the north, a complex system of structures (including salt diapirs) put the basin edges in contact with Triassic, Jurassic and Permian materials. In the region of the evaporite basins, the Paleogene fill of the graben lies directly on the Jurassic basement. The sedimentary filling of this rift sequence is asymmetrical with the deeper parts located at the southwestern and northeastern sides of the Graben (Rouchy, 1997).


    Palaeogeographical reconstructions place the potential marine seaway seepage feed to the north or perhaps also southeast of the Mulhouse Basin, while marginal continental conglomerates tend to preclude any contemporaneous hydrographic connection with Oligocene ocean water (Blanc-Valleron, 1991; Hinsken et al., 2007; Cendon et al., 2008). At the time of its hydrographic isolation, some 34 Ma, the basin was located 40° north of the equator. Total fill of Oligocene lacustrine/marine-fed sediments in the graben is some 1,700m thick. The saline stage is dominated by anhydrite, halite and mudstone. The main saline sequence is underlain by non-evaporitic Eocene continental mudstones, with lacustrine fossils and local anhydrite beds. Evaporite bed continuity in the northern part of the basin is disturbed by (Permian-salt cored) diapiric and or erosional/fault movement. Consequently, these northern basins are not considered suitable for conventional potash mining (Figure 6a).

    The Paleogene fill of the basin is divided into 6 units; a pre-evaporitic series, Lower Salt Group (LSG), Middle Salt Group (MSG), an Upper Salt Group (USG - with potash), Grey Marls Fm., and the Niederroedern Fm (Figure 7; Cendon et al., 2008). The LSG and lower section of the MSG are interpreted as lacustrine in origin, based on the limited palaeontologic and geochemical data. However, based on the presence of Cenozoic marine nannoplankton, shallow water benthic foraminifera, and well-diversified dinocyst assemblages in the fossiliferous zone below Salt IV, Blanc-Valleron (1991) favours a marine influence near the top of the MSG, while recognising the ambiguity of marine proportions with brackish faunas. Many marine-seepage fed brine systems have salinities that allow halotolerant species to flourish in marine-fed basins with no ongoing marine hydrographic connection (Warren, 2011). According to Blanc-Valleron and Schuler (1997), the region experienced a Mediterranean climate with long dry seasons during Salt IV member deposition.


    In detail, the Salt IV member is made up of some 210 m of evaporitic sequence, with two relatively thin potash levels (Ci and Cs). The stratigraphy associated with this potash zone is, from base to top (Figure 7):

    S2 Unit: 11.5 m thick with distinct layers of organic-rich marls, often dolomitic, with dispersed anhydrite layers.

    S1 Unit: 19 m thick, evenly-bedded and made up of alternating metre-scale milky (inclusion-rich) halite layers, with much thinner marls and anhydrite layers. Marls show a sub-millimetric lamination formed by micritic carbonate laminae alternating with clay, quartz, and organic matter-rich laminae. Hofmann et al. (1993a, b) interpreted these couplets as reflecting seasonal variations. Anhydrite occasionally displays remnant swallowtail ghost textures, which suggest that at least part of the anhydrite first precipitated as subaqueous gypsum. Halite shows an abundance of growth-aligned primary chevron textures, along with fluid-inclusion banding suggesting halite was subaqueous and deposited beneath shallow brine sheets (Lowenstein and Spencer, 1990).

    S Unit: Is 3.7 m thick and consists of thin marl layers and anhydrite, similar to the S2 Unit, with a few thin millimetric layers of halite.

    Mi Unit: With a thickness of 6 m, it is mostly halite with similar characteristics to the S1 Unit. Sylvite was detected in one sample, but its presence is probably related to the evolution of interstitial brines (Cendon et al., 2008).

    Ci Unit (“Couche inférieure”): Is formed by 4 m of alternating marls/anhydrite, halite, and sylvite beds (Figure 7).

    The Ti unit consists of alternating beds of halite, marl and anhydrite. The top of the interval is the T unit, which is similar to the S unit and consists of alternating beds of marl and anhydrite. Above this is the Ms or upper Marl, near identical to the lower marl Mi. The Mi is overlain by the upper potash bed (Cs), a thinner, but texturally equivalent, bed compared to the sylvinitic Ci unit.

    Thus, the Oligocene halite section includes two thin, but mined, potash zones: the Couche inferieure (Ci; 3.9m thick), and Couche superieure (Cs; 1.6m thick), both occur within Salt IV of the Upper Salt group (Figures 5, 7).

    Both potash beds are made up of stacked, thin, parallel-sided cm-dm-thick beds (averaging 8 cm thickness), which are in turn constructed of couplets composed of grey-coloured halite overlain by red-coloured sylvite (Figure 5b). Each couplet has a sharp base that separates the basal halite from the sylvite cap of the underlying bed. In some cases, the separation is also marked by bituminous partings. The bottom-most halite in each dm-thick bed consists of halite aggregates with cumulate textures that pass upward into large, but delicate, primary chevrons and cornets. Clusters of this chevron halite swell upward to create a cm-scale hummocky boundary with the overlying sylvite (Figure 5c; Lowenstein and Spencer, 1990).

    The sylvite member of a sylvinite couplet consists of granular aggregates of small transparent halite cubes and rounded grains of red sylvite (with some euhedral sylvite hoppers) infilling the swales in the underlying hummocky halite (Figure 5b,c). The sylvite layer is usually thick enough to bury the highest protuberances of the halite, so that the top of each sylvite layer, and the top of the couplet, is flat. Dissolution pipes and intercrystalline cavities are noticeably absent, although some chevrons show rounded coigns. Intercalated marker beds, formed during times of brine pool freshening, are composed of a finely laminated bituminous shale, with dolomite and anhydrite.

    The sylvite-halite couplets record combinations of unaltered settle-out and bottom-nucleated growth features, indicating primary chemical sediment accumulating in shallow perennial brine pools (Lowenstein and Spencer, 1990). Based on the crystal size, the close association of halites with sylvite layers, their lateral continuity and the manner in which sylvite mantles overlie chevron halites, the sylvites are interpreted as primary precipitates. Sylvite first formed at the air-brine surface or within the uppermost brine mass and then sank to the bottom to form well-sorted accumulations. As sylvite is a prograde salt it, like halite, probably grew during times of cooling of the brine mass (Figure 8a). These times of cooling could have been diurnal (day/night) or weather-front induced changes in the above-brine air temperatures. Similar cumulate sylvite deposits form as ephemeral bottom accumulations on the floor of modern Lake Dabuxum in China during its more saline phases.


    The subsequent mosaic textural overprint seen in many of the Mulhouse sylvite layers was probably produced by postdepositional modification of the crystal boundaries, much in the same way as mosaic halite is formed by recrystallisation of raft and cumulate halite during shallow burial. Temperature-based inclusion studies in both the sylvite and the halites average 63°C, suggesting solar heating of surface brines as precipitation took place (Figure 8b; Lowenstein and Spencer, 1990). Similar high at-surface brine temperatures are not unusual in many modern brine pools, especially those subject to periodic density stratification and heliothermometry (Warren 2016; Chapter 2).

    Mineralogically, potash evaporites in the Mulhouse Basin in the Rhine Graben (also known as the Alsatian (Alsace) or Wittelsheim Potash district) contain sylvite with subordinate carnallite, but lack the abundant MgSO4 salts characteristic of the evaporation of modern seawater. The Rhine graben formed during the Oligocene, via crustal extension, related to mantle upwelling. It was, and is, a continental graben typified by high geothermal gradients along its rift axis. In depositional setting, it is not dissimilar to pree-120,000-year potash fill stage in the Quaternary Danakil Basin or the Dead Sea during deposition of potash salts in the Pliocene Sedom Fm. The role of a high-temperature geothermal inflow in defining the CaCl2 nature of the potash-precipitating brines, versus a derivation from a MgSO4-depleted marine feed, is considered significant in the Rhine Graben deposits, but is poorly understood and still not resolved (Hardie, 1990; Cendón et al., 2008). World ocean chemistry in the Oligocene is on a shoulder between the MgSO4-depleted CaCl2-rich oceans of the Cretaceous and the MgSO4-enriched oceans of the Neogene (Figure 4).


    Cendón et al. (2008) conclude brine reaction processes were the most important factors controlling the major-ion (Mg, Ca, Na, K, SO4, and Cl) evolution of Mulhouse brines (Figure 9a-d). A combined analysis of fluid inclusions in primary textures by Cryo-SEM-EDS with sulphate- d34S, d18O and 87Sr/86Sr isotope ratios revealed likely hydrothermal inputs and recycling of Permian evaporites, particularly during the more advanced stages of evaporation that laid down the Salt IV member. Bromine levels imply an increasingly concentrated brine at that time (Figure 9a). The lower part of the Salt IV (S2 and S1) likely evolved from an initial marine input (Figure 9b-d).

    Throughout, the basin was disconnected from direct marine hydrographic connection and was one of a series of sub-basins formed in an active rift setting, where tectonic variations influenced sub-basin interconnections and chemical signatures of input waters. Sulphate-d34S shows Oligocene marine-like signatures at the base of the Salt IV member (Figure 9c, d). However, enriched sulphate-d18O reveals the importance of synchronous re-oxidation processes.

    As evaporation progressed, other non-marine or marine-modified inputs from neighbouring basins became more important. This is demonstrated by increases in K concentrations in brine inclusions and Br in halite, sulphate isotopes trends, and 87Sr/86Sr ratios (Figure 9b, c). The recycling (dissolution) of previously precipitated evaporites of Permian age was increasingly important with ongoing evaporation. In combination, this chemistry supports the notion of a connection of the Mulhouse Basin with basins situated north of Mulhouse. The brine evolution eventually reached sylvite precipitation. Hence, the chemical signature of the resulting brines is not 100% compatible with global seawater chemistry changes. Instead, the potash phase is tied to a hybrid inflow, with significant but decreasing marine input.

    There was likely an initial marine source, but this occurred within a series of rift-valley basin depressions for which there was no direct hydrographic connection to the open ocean, even at the time the Middle Salt Member (potash-entraining) was first deposited (Cendon et al., 2008). That is, the general hydrological evolution of the primary textured evaporites in the Mulhouse basin sump is better explained as a restricted sub-basin with an initial marine-seepage stage. This gradually changed to ≈ 40% marine source near the beginning of evaporite precipitation, with the rest of hydrological inputs being non-marine. There was a significant contribution of solutes from recycled, in part diapiric, Permian evaporites, likely remobilised by the tectonics driving the formation of the rift valley (Hinsken et al., 2007; Cendon et al., 2008). The general proportion of solutes did not change substantially over the time of evaporite precipitation. However, as the basin restriction increased, the formerly marine inputs changed to continental, diapiric or marine-modified inputs, perhaps fed from neighbouring basins north of Mulhouse basin. As in the Ethiopian Danakhil potash-rift, it is likely brine interactions occurred both during initial and early post-depositional reflux overprinting of the original potash salt beds.


    West Canadian potash (Devonian)

    The Middle Devonian (Givetian) Prairie Evaporite Formation is a widespread potash-entraining halite sequence deposited in the Elk Point Basin, an early intracratonic phase of the Western Canada Sedimentary Basin (WCSB; Chipley and Kyser, 1989). Today, it is the world’s predominant source of MOP fertiliser (Warren, 2016). The flexure that formed the basin and its subsealevel accommodation space was a distal downwarp to, and driven by, the early stages of the Antler Orogeny (Root, 2001). Texturally and geochemically the potash layers in the basin show the effects of multiple alterations and replacements of its potash minerals, especially interactions between sylvite and carnallite in a variably recrystallised halite host.

    Regionally halite constitutes a large portion of the four formations that make up the Devonian Elk Point Group (Figure 10): 1) the Lotsberg (Lower and Upper Lotsberg Salt), 2) the Cold Lake (Cold Lake Salt), 3) the Prairie Evaporite (Whitkow and Leofnard Salt), and 4) the Dawson Bay (Hubbard Evaporite). Today the remnants of the Middle Devonian Prairie Evaporite Formation constitute a bedded unit some 220 metres thick, which lies atop the irregular topography of the platform carbonates of the Winnipegosis Fm. Extensive solutioning of the various salts has given rise to an irregular thickness to the formation and the local absence of salt (Figure 11a).


    The Elk Point Group was deposited within what is termed the Middle Devonian “Elk Point Seaway,” a broad intracratonic sag basin extending from North Dakota and northeastern Montana at its southern extent north through southwestern Manitoba, southern and central Saskatchewan, and eastern to northern Alberta (Figure 11a). Its Pacific coast was near the present Alberta-British Columbia border, and the basin was centred at approximately 10°S latitude. To the north and west the basin was bound by a series of tectonic ridges and arches; but, due to subsequent erosion, the true eastern extent is unknown (Mossop and Shetsen, 1994). In northern Alberta, the Prairie Evaporite is correlated with the Muskeg and Presqu’ile formations (Rogers and Pratt, 2017).

    Hydrographic isolation of the intracratonic basin from its marine connection resulted in the deposition of a drawndown sequence of basinwide (platform-dominant) evaporites with what is a uniquely high volume of preserved potash salts deposited within a clayey halite host. The potash resource in this basin far exceeds that of any other known potash basin in the world.


    Potash geology

    Potash deposits mined in Saskatchewan are all found within the upper 60-70 m of the Prairie Evaporite Formation, at depths of more than 400 to 2750 metres beneath the surface of the Saskatchewan Plains. Within the Prairie Evaporite, there are four main potash-bearing members, in ascending stratigraphic order they are: Esterhazy, White Bear, Belle Plaine and Patience Lake members (Figure 11b). Each member is composed of various combinations of halite, sylvite, sylvinite, and carnallitite, with occurrences of sylvite versus carnallite reliably definable using wireline signatures (once the wireline is calibrated to core or mine control - Figure 12; Fuzesy, 1982).

    The Patience Lake Member is the uppermost Prairie Evaporite member and is separated from the Belle Plaine by 3-12 m of barren halite (Holter, 1972). Its thickness ranges from 0-21 m and averages 12 m, its top 7-14 m is made up halite with clay bands and stratiform sylvite. This is the targeted ore unit in conventional mines in the Saskatoon and Lanigan areas and is the solution-mined target, along with the underlying Belle Plaine Member, at the Mosaic Belle Plaine potash facility. The Belle Plaine member is separated from the Esterhazy by the White Bear Marker beds made up of some 15 m of low-grade halite, clay seams and sylvinite. The Belle Plaine Member is more carnallite-prone than the Patience Lake member (Figure 12). It is the ore unit in the conventional mines at Rocanville and Esterhazy (Figure 11b) where its thickness ranges from 0-18 m and averages around 9 m. In total, the Prairie Evaporite Formation does not contain any significant MgSO4 minerals (kieserite, polyhalite etc.) although some members do contain abundant carnallite. This mineralogy indicates precipitation from a Devonian seawater/brine chemistry somewhat different from today’s, with inherently lower relative proportions of sulphate and lower Mg/Ca ratios (Figure 4).

    The Prairie Evaporite Fm. is nonhalokinetic throughout the basin, it is more than 200 m thick in the potash mining district in Saskatoon and 140 m thick in the Rocanville area to the southeast (Figure 11a; Yang et al., 2009). The Patience Lake member is the main target for conventional mining near Saskatoon. The Esterhazy potash member rises close to the surface in the southeastern part of Saskatchewan near Rocanville and on into Manitoba. This is a region where the Patience Lake Member is thinner or completely dissolved (Figure 11b). Over the area of mineable interest in the Patience Lake Member, centred on Saskatoon, the ore bed currently slopes downward only slightly in a westerly direction, but deepens more strongly to the south at a rate of 3-9 m/km. Mines near Saskatoon are at depths approaching a kilometre and so are nearing the limits of currently economic shaft mining.

    The main shaft for the Colonsay Mine, which took IMC Global Inc. more than five years to complete through a water-saturated sediment column, finally reached the target ore body at a depth of 960 metres. Such depths and a southerly dip to the ore means that the conventional shaft mines near Saskatoon define a narrow WNW-ESE band of conventional mining activity (Figure 11c). To the south potash is recovered from greater depths by solution mining; for example, the Belle Plaine operation leaches potash from the Belle Plaine member at a depth of 1800m.

    The Prairie Evaporite typically thins southwards in the basin; although local thickening occurs where carnallite, not sylvite, is the dominant potash mineral (Worsley and Fuzesy, 1979). The Patience Lake member is mined at the Cory, Allan and Lanigan mines, and the Esterhazy Member is mined in the Rocanville area (Figure 11c). Ore mined from the 2.4 m thick Esterhazy Member in eastern Saskatchewan contain minimal amounts of insolubles (≈1%), but considerable quantities of carnallite (typically 1%, but up to 10%) and this reduces the average KCl grade value to an average of 25% K2O. The converse is true for ore mined from the Patience Lake potash member in western Saskatchewan near Saskatoon, where carnallite is uncommon in the Cory and Allan mines. The mined ore thickness is a 2.74-3.35 metre cut off near the top of the 3.66-4.57metre Prairie Lake potash member. Ore grade is 20-26% K2O and inversely related to thickness (Figure 12). The insoluble content is 4-7%, mostly clay and markedly higher than in the Rocanville mines.


    A typical sylvinite ore zone in the Patience Lake member can be divided into four to six units, based on potash rock-types and clay seams (Figures 12, 13a; M1-M6 of Boys, 1990). Units are mappable and have been correlated throughout the PCS Cory Mine with varying degrees of success, dependent on partial or complete loss of section from dissolution. Potash deposition appears to have been early and related to short-term brine seaway cooling and syndepositional brine reflux. So the potash layering (M1-M6) is cyclic, expressed in the repetitive distribution of hematite and other insoluble minerals (Figure 13). Desiccation polygons, desiccation cracks, subvertical microkarst pits and chevron halite crystals indicate that the Patience Lake member that encompasses the potash ore was deposited in and just beneath a shallow-brine, salt-pan environment (Figure 13b; Boys, 1990; Lowenstein and Spencer, 1990; Brodlyo and Spencer, 1987; pers. obs).

    Clay seams form characteristic thin stratigraphic segregations throughout the potash ore zone(s) of the Prairie Evaporite, as well as disseminated intervals, and constitute about 6% of the ore as mined. For example, the insoluble minerals found in the PCS Cory samples are, in approximate order of decreasing abundance: dolomite, clay [illite, chlorite (including swelling-chlorite/chlorite), and septechlorite, quartz, anhydrite, hematite, and goethite. Clay minerals make up about one-third of the total insolubles: other minor components include: potassium feldspar, hydrocarbons, and sporadic non-diagnostic palynomorphs (Figure 13; Boys, 1990).

    In all mines, the clays tend to occur as long continuous seams or marker layers between the potash zones and are mainly composed of detrital chlorite and illite, along withauthigenic septechlorite, montmorillonite and sepiolite (Mossman et al., 1982; Boys, 1990). Of the two chlorite minerals, septechlorite is the more thermally stable. The septechlorite, sepiolite and vermiculite very likely originated as direct products of settle-out, syndepositional dissolution or early diagenesis under hypersaline conditions from a precursor that was initially eolian dust settling to the bottom of a vast brine seaway. The absence of the otherwise ubiquitous septechlorite from Second Red Beds west of the zero-edge of the evaporite basin supports this concept (Figure 9, 10).


    Potash Textures

    Texturally, at the cm-scale, potash salt beds in the Prairie Evaporite (both carnallitite and sylvinite) lack the lateral continuity seen in primary potash textures in the Oligocene of the Mulhouse Basin (Figure 14). Prairie potash probably first formed as syndepositional secondary precipitates and alteration products at very shallow depths just beneath the sediment surface. These early prograde precipitates were then modified to varying degrees by ongoing fluid flushing in the shallow burial environment. The cyclic depositional distribution of disseminated insolubles as the clay marker beds was possibly due to a combination of source proximity, periodic enrichment during times of brine freshening and the strengthening of the winds blowing detritals out over the brine seaway. Possible intra-potash disconformities, created by dissolution of overlying potash-bearing salt beds, are indicated by an abundance of residual hematite in clay seams with some cutting subvertically into the potash bed. Except in, and near, dissolution levels and collapse features, the subsequent redistribution of insolubles, other than iron oxides, is not significant.

    In general, halite-sylvite (sylvinite) rocks in the Prairie Evaporite ore zones generally show two end member textures; 1) the most common is a recrystallised polygonal mosaic texture with individual crystals ranging from millimetres to centimetres and sylvite grain boundaries outlined by concentrations of blood-red halite (Figure 14a). 2) The other end member texture is a framework of euhedral and subhedral halite cubes enclosed by anhedral crystals of sylvite (Figure 14b). This is very similar to ore textures in the Salado Formation of New Mexico interpreted as early passive precipitates in karstic voids.

    Petrographically, the halite-carnallite (carnallitite) rocks display three distinct textures. Most halite-carnallite rocks contain isolated centimetre-sized cube mosaics of halite enclosed by poikilitic carnallite crystals (Figure 14c); 1) Individual halite cubes are typically clear, with occasional cloudy crystal cores that retain patches of syndepositional growth textures (Lowenstein and Spencer, 1990). 2) The second texture is coarsely crystalline halite-carnallite with equigranular, polygonal mosaic textures. In zones where halite overlies bedded anhydrite, most of the halite is clear with only the occasional crystal showing fluid inclusion banding.

    Bedded halite away from the ore zones generally retains a higher proportion of primary depositional textures typical of halite precipitation in shallow ephemeral saline pans (Figure 14d; Brodylo and Spencer, 1987). Crystalline growth fabrics, mainly remnants of vertically-elongate halite chevrons, are found in 50-90% of the halite from many intervals in the Prairie Evaporite. Many of the chevrons are truncated by irregular patches of clear halite that formed as early diagenetic cements in syndepositional karst.

    In contrast, the halite hosting the potash ore layers lacks well-defined primary textures but is dominated by intergrown mosaics. From the regional petrology and the lower than expected Br levels in halite in the Prairie Evaporite Formation, Schwerdtner (1964), Wardlaw and Watson (1966) and Wardlaw (1968) postulated a series of recrystallisation events forming sylvite after carnallite as a result of periodic flushing by hypersaline solutions. This origin as a secondary precipitate (via incongruent dissolution) is supported by observations of intergrowth and overgrowth textures (McIntosh and Wardlaw, 1968), collapse and dissolution features at various scales and timings (Gendzwill, 1978; Warren 2017), radiometric ages (Baadsgaard, 1987) and palaeomagnetic orientations of the diagenetic hematite linings associated with the emplacement of the potash (Koehler, 1997; Koehler et al., 1997).

    Dating of clear halite crystals in void fills within the ore levels shows that some of the exceptionally coarse and pure secondary halites forming pods in the mined potash horizons likely precipitated during early burial, while other sparry halite void fills formed as late as Pliocene-Pleistocene (Baadsgard, 1987). Even today, alteration and remobilisation of the sylvite and carnallite and the local precipitation of bischofite are ongoing processes, related to the encroachment of the contemporary dissolution edge or the ongoing stoping of chimneys fed by deep artesian circulation (pers obs.).


    Fluid inclusion studies support the notion of primary textures (low formation temperatures in chevron halite in the Prairie evaporite and an associated thermal separation of non-sylvite and sylvite associated halite (Figure 15; Chipley et al., 1990). Most fluid inclusions found in primary, fluid inclusion-banded halite associated with the Prairie potash salts contain sylvite daughter crystals at room temperature or nucleate them on cooling (e.g. halite at 915 and 945 m depth in the Winsal Osler well; Lowenstein and Spencer, 1990). In contrast, no sylvite daughter crystals have been observed in fluid inclusions outlining primary growth textures from chevron halites away from the potash deposits.

    The data illustrated as Figure 15 clearly show that inclusion temperatures in primary halite chevrons are cooler than those in halites collected in intervals nearer the potash levels. Sylvite daughter crystal dissolution temperatures from fluid inclusions in the cloudy centres of halite crystals associated with potash salts are generally warmer (Brodlyo and Spencer, 1987; Lowenstein and Spencer, 1990). Sylvite and carnallite daughter crystal dissolution temperatures from fluid inclusions in fluid inclusion banded halite from bedded halite-carnallite are the hottest. This mineralogically-related temperature schism establishes that potash salts occur in stratigraphic intervals in the halite where syndepositional surface brines were warmer. In the 50° - 70°C temperature range there could be overlap with heliothermal brine lake waters. Even so, these warmer potash temperatures imply parent brines would likely be moving via a shallow reflux drive and are not the result of primary bottom nucleation (in contrast to primary sylvite in the Mulhouse Basin). Whether the initial Prairie reflux potash precipitate was sylvite or carnallite is open to interpretation (Lowenstein and Hardie, 1990).


    Fluid evolution from mineral and isotope chemistry

    Analysis of subsurface waters from various Canadian potash mines and collapse anomalies in the Prairie Evaporite suggest that, after initial potash precipitation, a series of recrystallising fluids accessed the evaporite levels at multiple times throughout the burial history of the Prairie Formation (Chipley, 1995; Koehler, 1997; Koehler et al., 1997). Likewise, the isotope systematics and K-Ar ages of sylvite in both halite and sylvite layers indicate that the Prairie Evaporite was variably recrystallised during fluid overprint events (Table 2; Figure 16). These event ages are all younger than original deposition (≈390Ma) and likely correspond to ages of various tectonic events that influenced subsurface hydrology along the western margin of North America.


    Chemical compositions of inclusion fluids in the Prairie Evaporite, as determined by their thermometric properties, reveal at least two distinct waters played a role in potash formation: a Na-K-Mg-Ca-Cl brine, variably saturated with respect to sylvite and carnallite; and a Na-K-Cl brine (Horita et al., 1996). That is, contemporary inclusion water chemistry is a result in part of ongoing fluid-rock interaction. The ionic proportions in some halite samples are not the result of simple evaporation of seawater to the sylvite bittern stage (Figure 17a; Horita et al., 1996). There is a clear separation of values from chevron halites in samples from the Lanigan and Bredenbury (K-2 area) mines, which plot closer to the concentration trend seen in halite from modern seawater and values from clear or sparry halite. The latter encompass much lower K and higher Br related to fractionation tied to recrystallisation. Likewise, the influence of ongoing halite and potash salt dissolution is evident in the chemistry of shaft and mine waters with mine level waters showing elevated Mg and K values, (Figure 17b; Wittrup and Kyser, 1990; Chipley, 1995). What is more the mine waters of today show  substantial overlap with waters collected more than thirty years ago (Jensen et al., 2006)


    This notion of ongoing fluid-rock interaction controlling the chemistry of mine waters is supported by dD and d18O values of inclusion fluids in both halite and sylvite, which range from -146 to 0‰ and from -17.6 to -3.0‰, respectively (Figure 18). Most of the various preserved isotope values are different from those of evaporated seawater, which should have dD and d18O values near 0‰.

    Furthermore, the dD and d18O values of inclusion fluids are probably not the result of precipitation of the evaporite minerals from a brine that was a mixture of seawater and meteoric water. The low latitude position of the basin during the Middle Devonian (10-15° from the equator), the required lack of meteoric water to precipitate basinwide evaporites, and the expected dD and d18O values of any meteoric water in such a setting, make this an unlikely explanation. Rather, the dD and d18O values of inclusion fluids in the halites reflect ambient and evolving brine chemistries as the fluids in inclusions in the various growth layers were intermittently trapped during the subsurface evolution of the Prairie Formation in the Western Canada Sedimentary Basin. They also suggest that periodic migration of nonmarine subsurface water was a significant component of the crossflowing basinal brines throughout much of the recrystallisation history (Chipley, 1995).

    Prairie carnallite-sylvite alteration over time

    Ongoing alteration of carnallite to sylvite and the reverse reaction means a sylvite-carnallite bed must be capable of gaining or losing fluid at the time of alteration. That is, any reacting potash beds must be permeable at the time of the alteration. By definition, there must be fluid egress to drive incongruent alteration of carnallite to sylvite or fluid ingress to drive the alteration of sylvite to carnallite. There can also be situations in the subsurface where the volume of undersaturated fluid crossflow was sufficient to remove (dissolve) significant quantities of the more soluble evaporite salts. Many authors looking at the Prairie evaporite argue that the fluid access events during the alteration of carnallite to sylvite or the reverse, or the complete leaching of the soluble potash salts was driven by various tectonic events (Figure 16). In the early stages of burial alteration (few tens of metres from the landsurface) the same alteration processes can be driven by varying combinations of brine reflux, prograde precipitation and syndepositional karstification, all driven by changes in brine level and climate, which in turn may not be related to tectonism (Warren, 2016; Chapters 2 and 8 for details).

    In the potash areas of the Western Canada Sedimentary Basin, the notion of 10-100 km lateral continuity is a commonly stated precept for both sylvinite and carnallitite units across the extent of the Prairie Evaporite. But when the actual distribution and scale of units are mapped based on mined intercepts, there are numerous 10-100 metre-scale discontinuities (anomalies) present indicating fluid ingress or egress (Warren, 2017).


    Sometimes ore beds thin and alteration degrades the ore level (Section A-A1-A2), other times these discontinuities can locally enrich sylvite ore grade (B-B1; Figure 19). Discontinuities or salt anomalies are much more widespread in the Prairie evaporite than mentioned in much of the potash literature (Figure 19). Mining for maintenance of ore grade shows that unexpectedly intersecting an anomaly in a sylvite ore zone can have a range of outcomes ranging from the inconsequential to the catastrophic, in part because there is more than one type of salt anomaly or “salt horse" (Warren, 2017).


    Figure 20 summarises what are considered the three most common styles of salt anomaly in the sylvite ore beds of the Prairie Evaporite, namely 1) Washouts, 2) Leach anomalies, 3) Collapse anomalies. These ore bed disturbances and their occurrence styles are in part time-related. Washouts are typically early (eogenetic) and defined as... “salt-filled V- or U-shaped structures, which transect the normal bedded sequence and obliterate the stratigraphy” (Figure 20a; Mackintosh and McVittie, 1983, p. 60). They are typically enriched in, or filled by, insoluble materials in their lower one-third and medium-coarse-grained sparry halite in the upper two thirds. Up to several metres across, when traced laterally they typically pass into halite-cemented paleo-sinks and cavern networks (e.g. Figure 20b). Most washouts likely formed penecontemporaneous to the potash beds they transect, that is, they are preserved examples of synkarst, with infilling of the karst void by a slightly later halite cement. They indicate watertable lowering in a potash-rich saline sump. This leaching was followed soon after by a period of higher watertables and brine saturations, when halite cements occluded the washouts and palaeocaverns. Modern examples of this process typify the edges of subcropping and contemporary evaporite beds, as about the recently exposed edges of the modern Dead Sea. As such, “washouts” tend to indicate relatively early interactions of the potash interval with undersaturated waters, they may even be a part of the syndepositional remobilisation hydrology that focused, and locally enriched, potash ore levels.

    In a leach anomaly zone, the stratabound sylvinite ore zone has been wholly or partially replaced by barren halite, without significantly disturbing the normal stratigraphic sequence (clay marker beds) which tend to continue across the anomaly (Figure 20b). Some loss of volume or local thinning of the stratigraphy is typical in this type of salt anomaly. Typically saucer-shaped, they have diameters ranging from a few metres up to 400m. Less often, they can be linear features that are up to 20 m wide and 1600m long. Leach zones can form penecontemporaneously due to depressions and back-reactions in the ore beds, or by later low-energy infiltration of Na-saturated, K-undersaturated brines. The latter method of formation is also likely on the margins of collapse zones, creating a hybrid situation typically classified as a leach-collapse anomaly (Mackintosh and McVittie, 1983; McIntosh and Wardlaw, 1968).

    Of the three types of salt anomaly illustrated, leach zone processes are the least understood. Historically, when incongruent dissolution was the widely accepted interpretation for loss of unit thickness in the Prairie Evaporite, many leach anomalies were considered metasomatic. Much of the original metasomatic interpretation was based on decades of detailed work in the various salt mines of the German Zechstein Basin. There, in an endemic halokinetic terrane, evaporite textures were considered more akin to metamorphic rocks, and the term metasomatic alteration was commonly used when explaining leach anomalies (Bochert and Muir, 1964, Braitsch, 1971). In the past two decades, general observations of the preservation of primary chevron halite in most bedded evaporites away from the potash layers in the Prairie Evaporite have led to reduced use of notions of widespread metamorphic-like metasomatic or solid-state alteration processes in bedded evaporites. There is just too much preserved primary texture in the bedded salt units adjacent to potash beds to invoke pervasive burial metasomatism of the Prairie Evaporite.


    So how do leach anomalies, as illustrated in Figure 20b, occur in nonhalokinetic settings? One possible explanation is given by the depositional textures documented in anomalies in the Navarra Potash Province (Figure 21). There, the underlying and overlying salt stratigraphy is contiguous, while the intervening sylvite passed laterally into a syndepositional anomaly or “salt horse” created by an irregular topography on the salt pan floor prior to the deposition of onlapping primary sylvinite layers (see Warren 2016, 2017 for detailed discussion)

    On the other hand, in halokinetic situations (which characterises much Zechstein salt) solid-state alteration via inclusion related migration in flowing salt beds is a well-documented set of texture-altering processes (diffusion metasomatism). Most workers in such halokinetic systems would agree that there must have been an original stratiform potassium segregation present during or soon after deposition related to initial precipitation, fractional dissolution and karst-cooling precipitation. But what is controlling potassium distribution now in the Zechstein salts is a recrystallised and remobilised set of textures, which preserve little or no crystal-scale evidence of primary conditions (Warren, 2016; Chapter 6). The complex layering in such deposits may preserve a broad depositional stratigraphy, but the decimetre to metre scale mineral distributions are indications of complex interactions of folds, overfolds, and disaggregation with local flow thickening. We shall return to this discussion of Zechstein potash textures in the next section dealing with devolatisation of hydrated salts such as carnallite. in zones of local heating

    Collapse zones in the Prairie Evaporite are characterized by a loss of recognizable sylvinite ore strata, which is replaced by less saline brecciated, recemented and recrystallized material, with the breccia blocks typically made of the intrasalt or roof lithologies (Figure 20c), so angular fragments of the Second Red beds and dolostones of the Dawson Bay Formation are the most conspicuous components of the collapse features in the Western Canada Sedimentary Basin. When ore dissolution is well developed, all the halite can dissolve, along with the potash salts, and the overlying strata collapse into the cavity (these are classic solution collapse features). Transitional leached zones typically separate the collapsed core from normal bedded potash. Such collapse structures indicate a breach of the ore layers by unsaturated waters, fed either from below or above. For example, in the Western Canada Sedimentary Basin, well-developed collapse structures tend to occur over the edges and top Devonian mud mounds, while in the New Mexico potash zone the collapse zones are related to highs in the underlying Capitan reef trend (Warren, 2017). Leaching fluids may have come from below or above to form collapse structures at any time after initial deposition. When connected to a water source, these are the subsurface features that when intersected can quickly move the mining operation out of the salt into an adjacent aquifer system, a transition that led to flooding in most of the mine-lost operations listed earlier.

    Identifying at the mine scale the set of processes that created a salt anomaly in a sylvite bed also has implications in terms of its likely influence on mine stability whatever decision is made on how to deal with it as part of the ongoing mine operation (Warren, 2016, 2017). Syndepositional karst fills and leach anomalies are least likely to be problematic if penetrated during mining, as the aquifer system that formed them is likely no longer active. In contrast, penetration or removal of the region around a salt-depleted collapse breccia may lead to uncontrollable water inflows and ultimately to the loss of the mine.

    Unfortunately, in terms of production planning, the features of the periphery of a leach anomaly can be similar if not identical to those in the alteration halo that typically forms about the leached edge a collapse zone. The processes of sylvite recrystallisation that define the edge of collapse anomaly can lead to local enrichment in sylvite levels, making these zones surrounding the collapse core attractive extraction targets (Boys 1990, 1993). Boundaries of any alteration halo about a collapse centre are not concentric, but irregular, making the prediction of a feature’s geometry challenging, if not impossible. The safest course of action is to avoid mining salt anomalies, but longwall techniques make this difficult and so they must be identified and dealt with (see Warren 2017).


    Cooking sylvite: Dykes & sills in potash salts

     

    In addition to; 1) primary sylvite and 2) sylvite/carnallite alteration via incongruent transformation in burial, there is a third mode of sylvite formation related to 3) igneous heating driving devolatisation of carnallitite, which can perhaps be considered a form of incongruent melting (Warren, 2016). And so, at a local scale (measured in metres to tens of metres) in potash beds cut by igneous intrusions, there are a number of documented thermally-driven alteration styles and thermal haloes. Most are created by the intrusion of hot doleritic or basaltic dykes and sills into cooler salt masses, or the outflow of extrusive igneous flows over cooler salt beds (Knipping, 1989; Grishina et al, 1992, 1998; Gutsche, 1988; Steinmann et al., 1999; Wall et al., 2010). Hot igneous material interacts with somewhat cooler anhydrous salt masses to create narrow, but distinct, heat and mobile fluid-release envelopes, also reflected in the resulting salt textures. At times, relatively rapid magma emplacement can lead to linear breakout trends outlined by phreatomagmatic explosion craters, as imaged in portions of the North Sea (Wall et al., 2010) and the Danakhil/Dallol potash beds in Ethiopia (Salty Matters, May 1, 2015).

    Based on studies of inclusion chemistry and homogenization temperatures in fluid inclusions in bedded halite near intrusives, it seems that the extent of the influence of a dolerite sill or dyke in bedded salt is marked by fluid-inclusion migration, evidenced by the disappearance of chevron structures and consequent formation of clear halite with a different set of higher-temperature inclusions. Such a migration envelope is well documented in bedded Cambrian halites intruded by end-Permian dolerite dykes in the Tunguska region of Siberia (Grishina et al., 1992).

    Defining h as the thickness of the dolerite intrusion in these salt beds, and d as the distance of the halite from the edge of the intrusion, then the disappearance of chevrons occurs at greater distances above than below the intrusive sill. For d/h < 0.9 below the intrusion, CaCl2, CaCl2, KCl and nCaCl2, mMgCl2 solids occur in association with water-free and liquid-CO2 inclusions, with H2S, SCO and orthorhombic or glassy S8. For a d/h of 0.2-2 above the intrusion, H2S-bearing liquid-CO2 inclusions are typical, with various amounts of water. Thus, as a rule of thumb, an alteration halo extends up to twice the thickness of the dolerite sill above the sill and almost the thickness of the sill below (Figure 22).

    In a series of autoclavation laboratory experiments, Fabricius and Rose-Hampton (1988) found that; 1) at atmospheiric pressure carnallite melts incongruently to sylvite and hydrated MgCl2 at a temperature of 167.5°C. 2) the melting/transformation temperature increase to values in excess of 180°C as the pressure increases (Figure 23).


    A similar situation occurs in the dyke-intruded halite levels exposed in the mines of the Werra-Fulda district of Germany (Steinmann et al., 1999; Schofield, et al., 2014). There the Herfa-Neurode potash mine is located in the Werra-Fulda Basin in the Hessian district of central Germany (Figure 24a). The targeted ore levels consist of the carnallite-rich Kaliflöz Hessen (K1H) and Kaliflöz Thüringen (K1Th) intervals, which form part of the Zechstein 1 (Z1) bedded Werra salt succession (Warren, 2016). In the mine the K1H and K1Th units range in thickness from 2 m to 10 m, are generally subhorizontal and occur at a depth of 650–710 m below the present-day surface.


    In the later Tertiary, basaltic melts intruded these Zechstein evaporites as numerous sub-vertical dykes, but only a few dykes attained the Miocene landsurface. Basaltic melt production was related to regional volcanic activity some 10 to 25 Ma. Basalts exposed in the mine walls, where it cuts non-hydrous units of halite or anhydrite, are typically subvertical dykes, rather than subhorizontal sills. The basalts are phonolitic tephrites, limburgites, basanites and olivine nephelinites. Dyke margins are usually vitrified, forming a microlitic limburgite glass along dyke edges in contact with halite (Figure 24b; Knipping, 1989). At the contact on the evaporite side of the glassy rim, there is a cm-wide carapace of high-temperature salts (mostly anhydrite and ferroan carbonates). Further out, the effect of the high-temperature envelope is denoted by transitions to clear halite, with higher temperature fluid inclusions (Knipping 1989). All of this metre-scale alteration is an anhydrous alteration halo, the halite did not melt (melting temperature of 804°C), rather than migrating, the fluid driving recrystallisation was mostly from entrained brine/gas inclusions. The dolerite/basalt interior of the basaltic dyke is likewise altered and salt soaked, with clear, largely inclusion-free halite typically filling vesicles in the basalt.

    Heating of hydrated (carnallitic) salt layers, adjacent to a dyke or sill, tends to drive off the water of crystallisation (chemical or hydration thixotropy) at much lower temperatures than that at which anhydrous salts, such as halite or anhydrite, thermally melt (Figure 24c; Table 3). For example, in the Fulda region, the thermally-driven release of water of crystallisation within carnallitic beds creates thixotropic or subsurface “peperite” textures as carnallitite alters to sylvinite layers. These are layers where heated water of crystallisation escaped from the hydrated-salt lattice. Dehydration-driven loss of mechanical strength focuses zones of magma entry into particular subhorizontal horizons in the salt mass, wherever hydrated salt layers were present. In contrast, dyke and sill margins are much sharper and narrower in zones of contact with anhydrous salt intervals and the intrusive is sub-vertical to steeply dipping (Figure 24b versus 24c).

    Accordingly, away from the immediate vicinity of the direct thermal aureole, heated and overpressured dehydration waters can enter carnallite halite bed, and drive the creation of extensive soft sediment deformation and peperite textures in hydrated layer (Figure 24c). Mineralogically, sylvite and coarse recrystallised halite dominate the salt fraction in the peperite intervals of the Herfa-Neurode mine. Sylvite in the altered zone is a form of dehydrated carnallite, not a primary-textured salt. Across the Fulda region, such altered zones and deformed units can extend along former carnallite layer to tens or even a hundred or more metres from the dyke feeder. Ultimately, the deformed potash bed passes back out into the unaltered bed, which retains abundant inclusion-rich halite and carnallite (Schofield et al., 2014).

    That is, nearer the basalt dyke, the carnallite is largely transformed into inclusion-poor halite and sylvite, the result of incongruent flushing of warm saline fluids mobilised from the hydrated carnallite crystal lattice as it was heated by dyke emplacement. During Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-fluids were mixed with fluids and gases originating from thermally-mobilised crystallisation water in the carnallite, as it converted to sylvite. This brine/gas mixture altered the basalts during post-intrusive cooling, an event which numerical models suggest was quite rapid (Knipping, 1989): a dyke of less than 0.5 m thickness probably cooled to temperatures less than 200°C within 14 days of dyke emplacement.

    The contrast in alteration extent between anhydrous and hydrous salt layers shows alteration effects are minimal wherever the emplacement temperature of the magma is below that of the anhydrous salt body as it is next to a basalt dyke. If this is the mechanism driving entry of igneous-related volatiles (gases and liquids) into a salt body, then the distribution of products (including CO2) will be highly inhomogeneous and related to the minerally of the salt unit adjacent to the intrusive. Worldwide, dykes intersecting salt beds tend to widen to become sills in two zones: 1) along evaporite units within the halite mass that contain hydrated salts, such as carnallite or gypsum (Figure 24c) and, 2) where rising magma has ponded and so created laccoliths at the upper or lower halite contact with the adjacent nonsalt strata or against a salt wall (Figure 22 vs 24). The first is a response to a pulse of released water as dyke-driven heating forces the dehydration of hydrated salt layers. The second is a response to the mechanical strength contrast at the salt-nonsalt contact.

    In summary, sylvite formed from a carnallite precursor during Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-fluids were mixed with fluids originating from thermally-mobilised crystallisation water in the carnallite, as it converted to sylvite. This brine mixture altered the basalts during post-intrusive cooling, an event which numerical models suggest was quite rapid (Knipping, 1989): a dyke of less than 0.5 m thickness probably cooled to temperatures less than 200°C within 14 days of dyke emplacement.

    How do we produce potash salts?

     

    Over this series of three articles focused on current examples of potash production, we have seen there are two main groups of potash minerals currently utilised to make fertiliser, namely, muriate of potash (MOP) and sulphate of potash (SOP). MOP is both mined (generally from a Pre-Neogene sylvinite ore) or produced from brine pans (usually via processing of a carnallitite slurry). In contrast, large volumes of SOP are today produced from brine pans in China and the USA but with only minor production for solid-state ore targets. Historically, SOP was produced from solid-state ores in Sicily, the Ukraine, and Germany but today there are no conventional mines with SOP as the prime output in commercial operation (See Salty Matters, May 12, 2015).

    The MgSO4-enriched chemistry of modern seawater makes the economic production of potash bitterns from a seawater-feed highly challenging. Today, there is no marine-fed plant anywhere in the world producing primary sylvite precipitates. However, sylvite is precipitating from a continental brine feed in salt pans on the Bonneville salt flat, Utah. There, a brine field, drawing shallow pore waters from saltflat sediments, supplies suitably low-MgSO4 inflow chemistry to the concentrator pans. Sylvite also precipitates in solar evaporator pans in Utah that are fed brine circulated through the abandoned workings of the Cane Creek potash mine (Table 1).

    Large-scale production of MOP fertiliser from potash precipitates created in solar evaporation pans is taking place in perennially subaqueous saline pans of the southern Dead Sea and the Qaidam Basin. In the Dead Sea, the feed brine is pumped from the waters of the northern Dead Sea basin, while in the Qaidam sump the feed is from a brine field drawing pore brines with an appropriate mix of river and basinal brine inputs. In both cases, the resulting feed brine to the final concentrators is relatively depleted in magnesium and sulphate. These source bitterns have ionic proportions not unlike seawater in times of ancient MgSO4-depleted oceans. Carnallitite slurries, not sylvinite, are the MOP precipitates in pans in both regions. When feed chemistry of the slurry is low in halite, then the process to recover sylvite is a cold crystallisation technique. When halite impurity levels in the slurry are higher, sylvite is manufactures using a more energy intensive, and hence more expensive, hot crystallisation technique. Similar sulphate-depleted brine chemistry is used in Salar Atacama, where MOP and SOP are recovered as byproducts of the production of lithium carbonate brines.

    Significant volumes of SOP are recovered from a combination of evaporation and cryogenic modification of sulfate-enriched continental brines in pans on the edge of the Great Salt Lake, Utah, and Lop Nur, China. When concentrated and processed, SOP is recovered from the processing of a complex series of Mg-K-SO4 double salts (schoenitic) in the Odgden pans fed brines from the Great Salt Lake. The Lop Nur plant draws and concentrates pore waters from a brine field drawing waters from glauberite-polyhalite-entraining saline lake sediments.

    All the Quaternary saline lake factories supply less than 20% of the world's potash; the majority comes from the conventional mining of sylvinite ores. The world's largest reserves are held in Devonian evaporites of the Prairie Evaporite in the Western Canada Sedimentary Basin. Textures and mineral chemistry show that the greater volume of bedded potash salts in this region is not a primary sylvite precipitate. Rather the ore distribution, although stratiform and defined by a series of clay marker beds, actually preserves the effects of multiple modifications and alterations tied to periodic egress and ingress of basinal waters. Driving mechanisms for episodes of fluid crossflow range from syndepositional leaching and reflux through to tectonic pumping and uplift (telogenesis). Ore distribution and texturing reflect local-scale (10-100 metres) discontinuities and anomalies created by this evolving fluid chemistry. Some alteration episodes are relatively benign in terms of mineralogical modification and bed continuity. Others, generally tied to younger incidents (post early Cretaceous) of undersaturated crossflow and karstification, can have substantial effects on ore continuity and susceptibility to unwanted fluid entry. In contrast, ore textures and bed continuity in the smaller-scale sylvinite ores in the Oligocene Mulhouse Basin, France, indicate a primary ore genesis.

    What makes it economic?

    Across the Quaternary, we need a saline lake brine systems with the appropriate brine proportions, volumes and climate to precipitate the right association of processable potash salts. So far, the price of potash, either MOP or SOP, and the co-associated MgSO4 bitterns, precludes industrial marine-fed brine factories.

    In contrast, to the markedly nonmarine locations of potash recovery from the Quaternary sources, almost all pre-Quaternary potash operations extract product from marine-fed basinwide ore hosts during times of MgSO4-depleted and MgSO4-enriched oceans (Warren, 2016; Chapter 11). This time-based dichotomy in potash ore sources with nonmarine hosts in the Quaternary deposits and marine evaporite hosted ore zones in Miocene deposits and older, reflects a simple lack of basinwide marine deposits and appropriate marine chemistry across the Neogene (Warren, 2010). As for all ancient marine evaporites, the depositional system that deposited ancient marine-fed potash deposits was one to two orders of magnitude larger and the resultant deposits were typically thicker stacks than any Quaternary potash settings. The last such “saline giant” potash system was the Solfifera series in the Sicilian basin, deposited as part of the Mediterranean “salinity crisis,” but these potential ore beds are of the less economically attractive MgSO4- enriched marine potash series.

    So, what are the factors that favour the formation of, and hence exploration for, additional deposits of exploitable ancient potash? First, large MOP solid-state ore sources are all basinwide, not lacustrine deposits. Within the basinwide association, it seems that intracratonic basins host significantly larger reserves of ore, compared to systems that formed in the more tectonically-active plate-edge rift and suture association. This is a reflection of: 1) accessibility – near the shallow current edge of a salt basin, 2) a lack of a halokinetic overprint and, 3) the setup of longterm, stable, edge-dissolution brine hydrologies that typify many intracratonic basins. Known reserves of potash in the Devonian Prairie evaporite in West Canadian Sedimentary Basin (WCSB) are of the order of 50 times that of next largest known deposit, the Permian of the Upper Kama basin, and more than two orders of magnitude larger than any other of the other known exploited deposits (Table 1).

    Part of this difference in the volume of recoverable reserves lies in the fact that the various Canadian potash members in the WCSB are still bedded and flat-lying. Beds have not been broken up or steepened, by any subsequent halokinesis. The only set of processes overprinting and remobilising the various potash salts in the WCSB are related to multiple styles and timings of aquifer encroachment on the potash units, and this probably took place at various times since the potash was first deposited, driven mainly by a combination of hinterland uplift and subrosion. In contrast, most of the other significant potash basins listed in Table 10 have been subjected to ongoing combinations of halokinesis and groundwater encroachments, making these beds much less laterally predictable. In their formative stages, the WCSB potash beds were located a substantial distance from the orogenic belt that drove flexural downwarp and creation of the subsealevel sag depression. Like many other intracratonic basins, the WCSB did not experience significant syndepositional compression or rift-related loading.

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    Brine evolution and origins of potash: primary or secondary? SOP in Quaternary saline lakes: Part 2 of 2

    John Warren - Friday, November 30, 2018

     


    Introduction

    This, the second in this series of articles on potash brine evolution deals with production of sulphate of potash in plants that exploit saline hydrologies hosted in Quaternary saline sumps. There are two settings where significant volumes of sulphate of potash salts are economically produced at the current time; the Ogden salt pans on the northeast shore of the Great Salt Lake in Utah and Lop Nur in China. Although potassium sulphate salts precipitate if modern seawater is evaporated to the bittern stage, as yet there is no operational SOP plant utilising seawater. This is due to concurrent elevated levels of magnesium and chlorine in the bittern, a combination that favours the precipitation of carnallite concurrently with the precipitation of double sulphate salts, such as kainite (Figure 1). Until now, this makes the processing of the multi-mineralogic precipitate for a pure SOP product too expensive when utilising a marine brine feed.

     

    Potash in Great Salt Lake, USA (SOP evolution with backreactions)

    Today sulphate of potash fertiliser is produced via a combination of solar evaporation and brine processing, using current waters of the Great Salt Lake, Utah, as the brine feed into the Ogden salt pans, which are located at the northeastern end of the Great Salt Lake depression in Utah (Figure 2a). A simpler anthropogenic muriate of potash (MOP) brine evolution occurs in the nearby Wendover salt pans on the Bonneville salt flats. There, MOP precipitates as sylvinite in concentrator pans (after halite). The Bonneville region has a bittern hydrochemistry not unlike like the evolved Na-Cl brines of Salar de Atacama, as documented in the previous article, but it is a brinefield feed without the elevated levels of lithium seen in the Andean playa (Figure 3b).

    Great Salt Lake brine contains abundant sulphate with levels sufficiently above calcium that sulphate continues to concentrate after most of the Ca has been used up in the precipitation of both aragonite and gypsum. Thus, as the brines in the anthropogenic pans at Ogden approach the bittern (post halite) stage, a series of sulphate double salts precipitate (Figure 4), along with carnallite and sylvite.


    Great Salt Lake brines

    The ionic proportions in the primary brine feed that is the endorheic Great Salt Lake water depends on a combination of; 1) the inflow volumes from three major rivers draining the ranges to the east, 2) groundwater inflow, 3) basin evaporation, and 4) precipitation (rainfall/snowfall) directly on the lake (Jones et al., 2009). Major solute inputs can be attributed to calcium bicarbonate-type river waters mixing with sodium chloride-type springs, which are in part hydrothermal and part peripheral recycling agents for NaCl held in the lake sediments. Spencer et al. (1985a) noted that prior to 1930, the lake concentration inversely tracked lake volume, which reflected climatic variation in the drainage. However, since that time, salt precipitation, primarily halite and mirabilite, and dissolution have periodically modified lake brine chemistry and led to density stratification and the formation of brine pockets of different composition.

    Complicating these processes is repeated fractional crystallisation and re-solution (backreaction) of lake mineral precipitates. The construction of a railway causeway has restricted circulation, nearly isolating the northern from the southern part of the lake, which receives over 95% of the inflow. Given that Great Salt Lake waters are dominated by Na and Cl, this has led to halite precipitation in the north (Figures 2a, 3a; Gwynn, 2002). Widespread halite precipitation also occurred before 1959, especially in the southern area of the lake, associated with the most severe droughts (Jones et al., 2009; Spencer et al., 1985a). Spencer et al. (1985a) also described the presence of a sublacustrine ridge, which probably separated the lake into two basins at very low lake stands in the Quaternary. Fluctuating conditions emphasise brine differentiation, mixing, and fractional precipitation of salts as significant factors in solute evolution, especially as sinks for CaCO3, Mg, and K in the lake waters and sediments. The evolution of these brine/rock system depends on the concentration gradient and types of suspended and bottom clays, especially in relatively shallow systems.


    Brine evolution across the Ogden pans

    Figure 3a plots the known hydrochemistry of the inflow waters to the Great Salt Lake and their subsequent concentration. Evolving lake waters are always Na-Cl dominant, with sulphate in excess of magnesium in excess of potassium, throughout. Any post-halite evaporite minerals from this set of chemical proportions will contain post-halite potash bittern salts with elevated proportions of sulphate and magnesium and so will likely produce SOP rather than MOP associations. Contrast these hydrochemical proportions with the inflow and evolution chemistry in the pore brines of the Bonneville salt flat (Figure 3b) the Dead Sea and normal marine waters. Across all examples, sodium and chlorine are dominant and so halite will be the predominant salt deposited after aragonite and gypsum (Figure 4). Specifically, there are changes in sulphate levels with solar concentration (Figure 3b). In brines recovered from feeder wells in the Bonneville saltflat, unlike the nearby bajada well waters, the Bonneville salt flat brines show potassium in excess of sulphate and magnesium. In such a hydrochemical system, sylvite, as well as carnallite, are likely potassium bitterns in post-halite pans. The Wendover brine pans on the Bonneville saltflat produce MOP, not SOP, along with a MgCl2 brine, and have done so for more than 50 years (Bingham, 1980; Warren, 2016).


    The mineral series in the Ogden pans

    Figure 5 illustrates a laboratory-based construction of the idealised evolution of a Great Salt Lake feed brine as it passes through the various concentration pans. Figure 5 is a portion of the theoretical 25°C sulphate-potassium-magnesium phase diagram for the Great Salt Lake brine system and shows precipitates that are in equilibrium with brine at a particular concentration. Figures 4 and 5 represent typical brine concentration paths at summertime temperatures (Butts, 2002). Importantly, these figures do not describe the entire brine concentration story and local variations; mineralogical complexities in the predicted brine stream are related to thermal stratification, retention times and pond leakage. Effects on the chemistry of the brine due to the specific day-by-day and season-by-season variations of concentration and temperature, which arise in any solar ponding operation, require onsite monitoring and rectification. Such ongoing monitoring is of fundamental import when a pilot plant is constructed to test the reality of a future brine plant and its likely products.

     

    Figure 6 illustrates the idealised phase evolution of pan brines at Ogden in terms of a K2SO4 phase diagram (no NaCl or KCl co-precipitates are shown; Felton et al., 2010). Great Salt Lake brine is pumped into the first set of solar ponds where evaporation initially proceeds along the line shown as Evap 1 until halite reaches saturation and is precipitated. Liquors discharged from the halite ponds are transferred to the potash precipitation ponds where solar evaporation continues as line Evap 2 on the phase diagram and potassium begins to reach saturation after about 75% of the water is removed. Potassium, sodium levels rise with further evaporation and schoenite precipitates in the schoenite crystalliser. After some schoenite precipitation occurs, the liquor continues to evaporate along the Evap 3 line to the point that schoenite, sylvite, and additional halite precipitate. Evaporation continues as shown by line Evap 3 to the point that kainite, sylvite and halite become saturated and precipitate. From this plot, the importance of the relative levels of extraction/precipitation of sulphate double salts versus chloride double salts is evident, as the evaporation plot point moves right with increasing chloride concentrations. That is, plot point follows the arrows from left to right as concentration of chloride (dominant ion in all pans) increases and moves the plot position right.


    Production of SOP in the Great Salt Lake

    To recover sulphate of potash commercially from pan bitterns fed from the waters of Great Salt Lake, the double salts kainite and schoenite are first precipitated and recovered in post-halite solar ponds (Figures 6, 7). The first salt to saturate and crystallise in the concentrator pans is halite. This is successively followed by epsomite, schoenite, kainite, carnallite, and finally bischofite. To produce a desirable SOP product requires ongoing in-pan monitoring and an on-site industrial plant whereby kainite is converted to schoenite. The complete salt evolution and processing plant outcome in the Ogden facility is multiproduct and can produce halite, salt cake and sulphate of potash and a MgCl2 brine product. Historically, sodium sulphate was recovered from the Great Salt Lake brines as a byproduct of the halite and potash production process, but ongoing low prices mean Na2SO4 has not been economically harvested for the last decade or so.

    The complete production and processing procedure is as follows (Figure 7; Butts, 2002, 2007; Felton et al., 2010): 1) Brine is pumped from the Great Salt Lake into solar evaporation ponds where sodium chloride precipitates in the summer. 2) When winter weather cools the residual (post-halite) brine in the pans to -1 to -4°C, sodium sulphate crystals precipitate as mirabilite in a relatively pure state. Mirabilite crystals can be picked up by large earth-moving machinery and stored outdoors each winter until further processing takes place. 3) The harvested mirabilite can be added to hot water, and anhydrous sodium sulphate precipitated by the addition of sodium chloride to the heated mix to reduce sodium sulphate solubility through the common ion effect. The final salt cake product is 99.5% pure Na2SO4. 4) To produce SOP, Great Salt Lake brines are allowed to evaporate in a set of halite ponds, until approaching saturation with potassium salts. The residual brine is then transferred to a mixing pond, where it mixes with a second brine (from higher up the evaporation series, that contains a higher molar ratio of magnesium to potassium. 5) This adjusted brine is then allowed to evaporate to precipitate sodium chloride once more, until it is again saturated with respect to potassium salts. 6) The saturated brine is then transferred to another pond, is further evaporated and precipitates kainite (Figures 5, 6). Kainite precipitation continues until carnallite begins to form, at which time the brine is moved to another pond and is allowed to evaporate further to precipitate carnallite. 6) Some of the kainite-depleted brine is recycled to the downstream mixing pond to maintain the required molar ratio of magnesium to calcium in this earlier mixing pond (step 4). 7) Once carnallite has precipitated, the residual brine is transferred to deep storage and subjected to winter cooling to precipitate additional carnallite as it is a prograde salt. 8) Cryogenically precipitated carnallite can be processed to precipitate additional kainite by mixing it with a kainite-saturated brine. 9) MgCl2-rich end-brines in the post-carnallite bittern pans are then further processed to produce either MgCl2 flakes or a 32% MgCl2 brine. These end-bitterns are then used as a feedstock to make magnesium metal, bischofite flakes, dust suppressants, freeze preventers, fertiliser sprays, and used to refresh flush in ion exchange resins.

    Some complexities in the observed mineral precipitation series in the Ogden Pans

    Under natural solar pond conditions in the Ogden Pans, the brine temperature fluctuates with the air temperature across day-night and seasonal temperature cycles, and there is a lag time for temperature response in waters any brine pan, especially if the pan is heliothermic. Atmosphere-driven fluctuations in temperature results in changes in ion saturations, which can drive selective precipitation or dissolution of salts in the brine body. Air temperature in the Ogden pans may be 35°C during the day and 15°C at night. Brine at point A in figure 5 may favour the formation of kainite during the daytime and schoenite at night. The result of the diurnal temperature oscillation is a mixture of both salts in a single pond from the same brine. In terms of extracted product, this complicates ore processing as a single pan will contain both minerals, produced at the same curing stage, at the same time, yet one double salt entrains KCl, the other K2SO4, so additional processing is necessary to purify the product stream (Butts, 2002).

    The sulphate ion in the pan waters is particularity temperature sensitive, and salts containing it in GSL pans tend to precipitate at cooler temperatures. Surficial cooling during the summer nights can cause prograde salts to precipitate, but the next day's heat generally provides sufficient activation energy to cause total dissolution of those salts precipitated just a few hours before (Butts, 2002). It is not unusual to find a 0.5 cm layer of hexahydrite (MgSO4.6H2O) at the bottom of a solar pond in the morning, but redissolved by late afternoon.

    Under controlled laboratory conditions, brine collected from the hypersaline north arm of the Great Salt Lake will not crystallise mirabilite  until the brine temperature reaches 2°C or lower. Yet, in the anthropogenic solar ponds, mirabilite has been observed to crystallise at brine temperatures above 7°C. During the winter, as the surface temperature of the GSL pan brine at night becomes very cold (2°C or lower), especially on clear nights, and mirabilite rafts will form on and just below the brine surface and subsequently sink into the somewhat warmer brine at the floor of the pond. Because there is insufficient activation energy in this brine to completely redissolve the mirabilite, it remains on the pan floor, until warmer day/night temperatures are attained. However, it is also possible for salts precipitated by cooling to be later covered by salts precipitated by evaporation, which effectively prevents dissolution of those more temperature-sensitive salts that would otherwise redissolve (Butts, 2002).

    There are also longer terms seasonal influences on mineralogy. Some salts deposited in June, July, and August (summer) will convert to other salts, with a possible total change in chemistry, when they are exposed to colder winter temperatures and rainfall. Kainite, for example, may convert to sylvite and epsomite and become a hardened mass on the pond floor; or if it is in contact with a sulphate-rich brine, it can convert to schoenite. Conversely, mirabilite will precipitate in the winter but redissolve during the hot summer months.

    The depth of a solar pond also controls the size of the crystals produced. For example, if halite (NaCl) is precipitated in a GSL pond that is either less than 8 cm or more than 30cm deep, it will have a smaller crystal size than when precipitated in a pond between 8 and 30 cm deep. Smaller crystals of halite are undesirable in a de-icing product since a premium price is paid for larger crystals.

    In terms of residence time, some salts require more time than others to crystallise in a pan. Brine that is not given sufficient time for crystallisation before it is moved into another pond, which contains brine at a different concentration, will produce a different suite of salts. For example, if a brine supersaturated in ions that will produce kainite, epsomite, and halite (reaction I), is transferred to another pond, the resulting brine mixture can favour carnallite (reaction 2), while kainite salts are eliminated.

    Reaction 1: 9.75H2O + Na+ + 2Cl- + 2Mg2+ + K+ + 2SO42+ —> MgSO4.KCl +2.75H2O + MgSO4.7H2O + NaCl

    Reaction 2: 12H2O + Na+ + 4Cl- + 2Mg2+ + K+ + 2SO42+ —> MgCl2.6H2O +MgSO4.6H2O + NaCl

    Reaction 1 retains more magnesium as MgCl2 in the brine; reaction 2 retains more sulphate. In reaction 2, it is also interesting to note the effect of waters-of-hydration on crystallization; forcing out salts with high waters of crystallization results in higher rates of crystallization. The hydrated salts remove waters from the brine and further concentrate the brine in much the same way as does evaporation.

    Pond leakage and brine capture (entrainment) in and below the pan floor are additional influences on mineralogy, regardless of brine depth or ponding area. As mentioned earlier, to precipitate bischofite and allow for MgCl2 manufacture, around ninety-eight percent of the water from present North Arm brine feed must evaporate. If pond leakage causes the level of the ponding area to drop too quickly, it becomes near impossible to reach saturation for bischofite (due to brine reflux). Control of pond leakage in the planning and construction phases is essential to assure that the precipitated salts contain the optimal quantity of the desired minerals for successful pond operation.

    The opposite of leakage is brine retention in a precipitated layer; it can also alter brine chemistry and recovery economics. Brine entrained (or trapped) in the voids between salt crystals in the pond floor is effectively removed from salt production and so affects the chemistry of salts that will be precipitated as concentration proceeds and can also drive unwanted backreactions. The time required to evaporate nearly ninety percent of the water from the present north arm Great Salt Lake brine in the Ogden solar pond complex, under natural steady state conditions, is approximately eighteen months.

    Summary of SOP production procedures in Great Salt Lake

    Sulphate of potash cannot be obtained from the waters of the Great Salt Lake by simple solar evaporation (Behrens, 2002). As the lake water is evaporated, first halite precipitates in a relatively pure form and is harvested. By the time evaporative concentration has proceeded to the point that saturation in a potash-entraining salt occurs, most of the NaCl has precipitated. It does, however, continue to precipitate and becomes the primary contaminant in the potassium-bearing salt beds in the higher-end pans.

    Brine phase chemistry from the point of potassium saturation in the evaporation series is complicated, and an array of potassium double salts are possible, depending on brine concentration, temperature and other factors. Among the variety of potash minerals precipitated in the potash harvester pans, the majority are double salts that contain atoms of both potassium and magnesium in the same molecule, They are dominated by kainite, schoenite, and carnallite. All are highly hydrated; that is, they contain high levels of water of crystallisation that must be removed during processing. SOP purification also involves removal of the considerable quantities of sodium chloride that are co-precipitated, after this the salts must be chemically converted into potassium sulphate.

    Controlling the exact mineralogy of the precipitated salts and their composition mixtures is not possible in the pans, which are subject to the vagaries of climate and associated temperature variations. Many of the complex double salts precipitating in the pans are stable only under fixed physiochemical conditions, so that transitions of composition may take place in the ponds and even in the stockpile and early processing plant steps.

    While weathering, draining, temperature and other factors can be controlled to a degree, it is essential that the Great Salt Lake plant be able to handle and effectively accommodate a widely variable feed mix (Behrens 2002). To do this, the plant operator has developed a basic process comprising a counter-current leach procedure for converting the potassium-bearing minerals through known mineral transition stages to a final potassium sulfate product (Figure 7). This set of processing steps is sensitive to sodium chloride content, so a supplemental flotation circuit is used to handle those harvested salts high in halite. It aims to remove the halite (in solution) and upgrade the feed stream to the point where it can be handled by the basic plant process.

    Solids harvested from the potash ponds with elevated halite levels are treated with anionic flotation to remove remaining halite (Felton et al., 2010). To convert kainite into schoenite, it is necessary to mix the upgraded flotation product with a prepared brine. The conversion of schoenite to SOP at the Great Salt Lake plant requires that new MOP is added, over the amount produced from the lake brines. This additional MOP is purchased from the open market. The schoenite solids are mixed with potash in a draft tube baffle reactor to produce SOP and byproduct magnesium chloride.

    The potassium sulfate processing stream defining the basic treatment process in the Great Salt Lake plant is summarised as Figure 7, whereby once obtaining the appropriate chemistry the SOP product is ultimately filtered, dried, sized and stored. Final SOP output may then be compacted, graded, and provided with additives as desired, then distributed in bulk or bagged, by rail or truck.


    Lop Nur, Tarim Basin, China (SOP operation)

    Sulphate of potash (SOP) via brine processing (solution mining) of lake sediments and subsequent solar concentration of brines is currently underway in the fault-bound Luobei Hollow region of the Lop Nur playa, in the southeastern part of Xinjiang Province, Western China (Liu et al., 2006; Sun et al., 2018). The recoverable sulphate of potash resource is estimated to be 36 million tonnes from lake brine (Dong et al., 2012). Lop Nur lies in the eastern part of the Taklimakan Desert (Figure 8a), China’s largest and driest desert, and is in the drainage sump of the basin, some 780 meters above sea level in a BSk climate belt. The Lop Nur depression first formed in the early Quaternary, due to the extensional collapse of the eastern Tarim Platform and is surrounded and typically in fault contact with the Kuruktagh (to the north), Bei Shan (to east) and Altun (to the south) mountains (Figure 8b).

    The resulting Lop Nur (Lop Nor) sump is a large groundwater discharge playa that is the terminal point of China’s largest endorheic drainage system, the Tarim Basin, which occupies an area of more than 530,000 km2 (Ma et al., 2010). The Lop Nur sump is the hydrographic base level to local and regional groundwater and surface water flow systems, and thus collectively captures all river and subsurface flow originating in the surrounding mountainous regions. The area has been subject to ongoing Quaternary climate and water supply oscillations, which over the last few hundred years has driven concentric strandzone contractions on the playa surface, to form what is sometimes called the “Great Ear Lake" of the Lop Nur sump (Liu et al., 2016a).

    Longer term widespread climate oscillations (thousands of years) drove precipitation of saline glauberite-polyhalite deposits, alternating with more humid lacustrine mudstones especially in fault defined grabens with the sump. For example, Liu et al. (2016b) conducted high-resolution multi-proxy analyses using materials from a well-dated pit section (YKD0301) in the centre of Lop Nur and south of the Luobei depression. They showed that Lop Nur experienced a progression through a brackish lake, saline lake, slightly brackish lake, saline lake, brackish lake, and playa in response to climatic changes over the past 9,000 years.

    Presently, the Lop Nur playa lacks perennial long-term surface inflow and so is characterised by desiccated saline mudflats and polygonal salt crusts. The upward capillary flux from the shallow groundwater helps to maintain a high rate of evaporation in the depression and drives the formation of a metre-thick ephemeral halite crust that covers much of the depression (Liu et al., 2016a).

    Historically, before construction of extensive irrigation systems in the upstream portion of the various riverine feeds to the depression and the diversion of water into the Tarim-Kongqi-Qargan canal, brackish floodwaters periodically accumulated in the Lop Nur depression. After the diversion of inflows, terminal desiccation led to the formation of the concentric shrinkage shorelines, that today outline the “Great Ear Lake” region of the Tarim Basin (Figure 8b; Huntington, 1907; Chao et al., 2009; Liu et al., 2016a).

    The current climate is cool and extremely arid (Koeppen BSk); average annual rainfall is less than 20 mm and the average potential evaporation rates ≈3500 mm/yr (Ma et al., 2008, 2010). The mean annual air temperature is 11.6°C; higher temperatures occur during July (>40°C), and the lower temperatures occur during January (<20°C). Primary wind direction is northeast. The Lop Nor Basin experiences severe and frequent sandstorms; the region is well known for its wind-eroded features, including many layered yardangs along the northern, western and eastern margins of the Lop Nur salt plain (Lin et al., 2018).


    Salinity and chemical composition of modern groundwater brine varies little in the ‘‘Great Ear” area and appears not to have changed significantly over the last decade (Ma et al., 2010). Dominant river inflows to the Lop Nor Basin are Na-Mg-Ca-SO4-Cl-HCO3 waters (Figure 9). In contrast, the sump region is characterised by highly concentrated groundwater brines (≈350 mg/l) that are rich in Na and Cl, poor in Ca and HCO3+CO3, and contain considerable amounts of Mg, SO4 and K, with pH ranging from 6.6 to 7.2 (Figure 9). When concentrated, the Luobei/Lop Nur pore brines is saturated with respect to halite, glauberite, thenardite, polyhalite and bloedite (Ma et al., 2010; Sun et al., 2018).

    Groundwater brines, pooled in the northern sub-depression, mostly in the Luobei depression, are pumped into a series of pans to the immediate south, where sulphate of potash is produced via a set of solar concentrator pans. Brines in the Luobei depression and adjacent Xingqing and Tenglong platforms are similar in chemistry and salinity to the Great Ear Lake area but with a concentrated saline reserve due to the presence of a series of buried glauberite-rich beds (Figure 9; Hu and Wang, 2001; Ma et al., 2010; Sun et al., 2018).

    K-rich mother brines in the Luobei hollow also contain significant MgSO4 levels and fill open phreatic pores in a widespread subsurface glauberite bed, with a potassium content of 1.4% (Liu et al., 2008; Sun et al., 2018). Feed brines are pumped from these evaporitic sediment hosts in the Luobei sump into a large field of concentrator pans to ultimately produce sulphate of potash (Figure 8a).

    Brine chemical models, using current inflow water and groundwater brine chemistries and assuming open-system hydrology, show good agreement between theoretically predicted and observed minerals in upper parts of the Lop Nor Basin succession (Ma et al., 2010). However, such shallow sediment modelling does not explain the massive amounts of glauberite (Na2SO4.CaSO4) and polyhalite (K2SO4MgSO4.2CaSO4.2H2O) recovered in a 230 m deep core (ZK1200B well) from the Lop Nor Basin (Figure 9a).


    Hydrochemical simulations assuming a closed system at depth and allowing brine reactions with previously formed minerals imply that widespread glauberite in the basin formed via back reactions between brine, gypsum and anhydrite and that polyhalite formed via a diagenetic reaction between brine and glauberite. Diagenetic textures related to recrystallisation and secondary replacement are seen in the ZK1200B core; they include gypsum-cored glauberite crystals and gypsum replacing glauberite. Such textures indicate significant mineral-brine interaction and backreaction during crystallisation of glauberite and polyhalite (Liu et al., 2008). Much of the glauberite dissolves to create characteristic mouldic porosity throughout the glauberite reservoir intervals (Figure 10, 11b)


    Mineral assemblages predicted from the evaporation of Tarim river water match closely with natural assemblages and abundances and, in combination with a model that allows widespread backreactions, can explain the extensive glauberite deposits in the Lop Nor basin (Ma et al., 2008, 2010). It seems that the Tarim river inflows, not fault-controlled upwelling hydrothermal brines, were the dominant ion source throughout the lake history. The layered distribution of minerals in the more deeply cored sediments documents the evolving history of inflow water response to wet and dry periods in the Lop Nor basin. The occurrence of abundant glauberite and gypsum below 40 m depth, and the absence of halite, polyhalite and bloedite in the same sediment suggests that the brine underwent incomplete concentration in the wetter periods 10b).

    In contrast, the increasing abundance of halite, polyhalite and bloedite in the top 40 m of core from the ZK1200B well indicate relatively dry periods (Figure 10a), where halite precipitated at lower evaporative concentrations (log Concentration factor = 3.15), while polyhalite and bloedite precipitated at higher evaporative concentrations (log = 3.31 and 3.48 respectively). Following deposition of the more saline minerals, the lake system once again became more humid in the later Holocene, until the anthropogenically-induced changes in the hydrology over the last few decades, driven by upstream water damming and extraction for agriculture (Ma et al., 2008). These changes have returned the sump hydrology to the more saline character that it had earlier in the Pleistocene.

    The Lop Nur potash recovery plant/factory and pan system, located adjacent to the LuoBei depression (Figures 8, 11a), utilises a brine-well source aquifer where the potash brine is reservoired in intercrystalline and vuggy porosity in a thick stacked series of porous glauberite beds/aquifers.

    Currently, 200 boreholes have been drilled in the Lop Nor brine field area showing the Late-Middle Pleistocene to Late Pleistocene strata are distributed as massive, continuous, thick layers of glauberite with well-developed intercrystal and mouldic porosity, forming storage space for potassium-rich brine (Figure 11b; Sun et al., 2018). However, buried faults and different rates of creation of fault-bound accommodation space, means there are differences in the brine storage capacity among the three brinefield extraction areas; termed the Luobei depression, the Xingqing platform and the Tenglong platform areas (Figures 9a, 11a).

    In total, there are seven glauberitic brine beds defined by drill holes in the Luobei depression, including a phreatic aquifer, W1L, and six artesian aquifers, W2L, W3L, W4L, W5L, W6L, and W7 (Figure 10b; Sun et al., 2018). At present, only W1L, W2L, W3L, and W4L glauberite seams are used as brine sources. There are two artesian brine aquifers, W2X and W3X, exposed by drill holes in the Xinqing platform and there are three beds in the Tenglong extraction area, including a phreatic aquifer, W1T, and two artesian aquifers, W2T and W3T (Figure 10b).

    W1L is a phreatic aquifer with layered distribution across the whole Luobei depression, with an average thickness of 17.54 m, water table depths of 1.7 to 2.3 m, porosities of 6.98% to 38.45%, and specific yields of 4.57% to 25.89%. Water yield is the highest in the central and northeast of the depression, with unit brine overflows of more than 5000 cubic meters per day per meter of water table depth (m3/dm). In the rest of the aquifer, the unit brine overflows range from 1000 to 5000 m3/dm (Sun et al., 2018). The W2L artesian aquifer is confined, nearly horizontal with a stratified distribution, and has an average thickness of 10.18 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 20 to 40 m, porosities of 4.34% to 37.8%, and specific yields of 1.08% to 21.04%. The W3L artesian aquifer is confined, with stratified distribution and an average thickness of 8.50 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 40 to 70 m, porosities of 2.85% to 19.97%, and specific yields of 1.10% to 13.37%. The W3L aquifer is also confined with stratified distribution, with an average thickness of 7.28 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 70 to 100 m, porosities of 5.22% to 24.72%, and specific yields of 1.03% to 9.91%. The lithologies of the four brine storage layers are dominated by glauberite, and occasional lacustrine sedimentary clastic rocks, such as gypsum (Figure 10a).

    The Xinqing platform consists of two confined potassium-bearing brine aquifers (Figure 10b). Confined brines have layered or stratified distributions. The average thicknesses of the aquifers are 4.38 to 7.52 m. Due to the F1 fault, there is no phreatic aquifer in the Xinqing platform, but this does not affect the continuity of the brine storage layer between the extraction areas. The W2X aquifer is confined, stratified, and distributed in the eastern part of this ore district with a north-south length of 77.78 km, east-west width of 16.82 km, and total area of 1100 km2. Unit brine overflows are 2.25 to 541.51 m3/dm, water table depths are 10 to 20 m, porosities are 3.89% to 40.69%, and specific yields are 2.01% to 21.15%. The W3X aquifer is also confined and stratified, with a north-south length of 76.10 km, east-west width of 18.81 km, and total area of 1444 km2. Unit brine overflows are 1.67 to 293.99 m3/d m, water table depths are 11.3 to 38 m, porosities are 4.16% to 26.43%, and specific yields are 2.11% to 14.19%23.

    The Tenglong platform consists of a phreatic aquifer and two confined aquifers. W1T is a phreatic, stratified aquifer and is the main ore body, and is bound by the F3 fault (Figure 10b). It is distributed across the northern part of the Tenglong extraction area, with a north-south length of about 33 km, east-west width of about 20 km, and total area of 610 km2. Water table depths are 3.26 to 4.6 m, porosities are 2.03% to 38.81%, and specific yields are 22.48% to 1.22%. On the other side of the F3 fault, in the southern part of the mining area, is the W2T confined aquifer (Figure 10b). Water table depths are 16.91 to 22 m, porosities are 3.58% to 37.64%, and specific yields are 1.35% to 18.69%. W3T is also a confined aquifer, with a stratified orebody distributed in the southern part of the mining area, with a north-south length of about 29 km, east-west width of about 21 km, and total area of 546 km2. Water table depths are 17.13 to 47 m, porosities are 2.69% to 38.71%, and specific yields are 1.26% to 17.64%.

    Lop Nur is an unusual potash source

    The glauberite-hosted brinefield in the Luobei depression and the adjacent platforms makes the Lop Nur SOP system unique in that it is the world's first large-scale example of brine commercialisation for potash recovery in a Quaternary continental playa aquifer system with a non-MOP brinefield target. Elsewhere, such as in the Dead Sea and the Qarhan sump, Salar de Atacama and the Bonneville salt flats, the brines derived from Quaternary lacustrine beds and water bodies are concentrated via solar evaporation in semi-arid desert scenarios. Potash plants utilising these Quaternary evaporite-hosted lacustrine brine systems do not target potassium sulphate, but process either carnallitite or sylvinite into a commercial MOP product

     Glauberite is found in a range of other continental Quaternary evaporite deposits around the world but, as yet,                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                       outside of Lop Nur is not economically exploited to produce sulphate of potash. For example, glauberite is a significant component in Quaternary cryogenic beds in Karabogazgol on the eastern shore of the Caspian Sea, in Quaternary evaporite beds in Laguna del Rey in Mexico, in saline lacustrine beds in the Miocene of Spain and Turkey, and in pedogenic beds in hyperarid nitrate-rich soils of the Atacama Desert of South America (Warren, 2016; Chapter 12).

    In most cases, the deposits are commercially exploited as a source of sodium sulphate (salt cake). In a saline Quaternary lake in Canada, SOP is produced by processing saline lake waters. This takes place in Quill Lake, where small volumes of SOP are produced via mixing a sylvite feed (trucked into the site) with a cryogenic NaSO4 lake brine.

    The Lop Nur deposit is mined by the SDIC Xinjiang Luobupo Hoevellite Co. Ltd, and the main product is potassium sulfate, with a current annual production capacity of 1.3 million tons. Pan construction began in 2000, and the plant moved in full -cale operation in 2004 when it produced ≈50,000 tons. The parent company, State Development and Investment Corporation (SDIC), is China’s largest state-owned investment holding company. The company estimates a potash reserve ≈ 12.2 billion tons in the sump. This makes Lop Nur deposit the largest SOP facility in the world, and it is now a significant supplier of high premium fertiliser to the Chinese domestic market.

    Implications

    A study of a few of the Quaternary pans worldwide manufacturing economic levels of potash via solar evaporation shows tha,t independent of whether SOP or MOP salts are the main product, all retain abundant evidence that salt precipitates continue to evolve as the temperature and the encasing brine chemistry change. As we shall see in many ancient examples discussed in the next article, ongoing postdepositional mineralogical alteration dominates the textural and mineralogical story in most ancient potash deposits.

    As we saw in the previous article, which focused on MOP in solar concentrator plants with brine feeds from Quaternary saline lakes, SOP production from brine feeds in Quaternary saline lakes is also related strongly to cooler desert climates (Figure 12). The Koeppen climate at Lop Nur is cool arid desert (BWk), while the Great Salt Lake straddles cool arid steppe desert and a temperate climate zone, with hot dry summer zones (BSk and Csa)


    Outside of these two examples, there are a number of other Quaternary potash mineral occurrences with the potential for SOP production, if a suitable brine processing stream can be devised (Warren, 2010, 2016). These sites include intermontane depressions in the high Andes in what is a high altitude polar tundra setting (Koeppen ET), none of which are commercial (Figure 12b).

    Similarly, there a number of non-commercial potash (SOP) mineral and brine occurrences in various hot arid desert regions in Australia, northern Africa and the Middle East (Koeppen BWh). Today, SOP in Salar de Atacama is currently produced as a byproduct of lithium carbonate production, along with MOP, as discussed in the previous article in this series.

    As for MOP, climatically, commercial potash brine SOP systems are hosted in Quaternary-age lacustrine sediments are located in cooler endorheic intermontane depressions (BWk, BSk). The association with somewhat cooler desert and less arid cool steppe climates underlines the need for greater volumes of brine to reside in the landscape in order to facilitate the production of significant volumes of potash bittern.

    Put simply, in the case of both MOP and SOP production in Quaternary settings, hot arid continental deserts simply do not have enough flowable water to produce economic volumes of a chemically-suitable mother brine. That is, currently economic Quaternary MOP and SOP operations produce by pumping nonmarine pore or saline lake brines into a set of concentrator pans. Mother waters reside in hypersaline perennial lakes in steep-sided valleys or in pores in salt-entraining aquifers with dissolving salt compositions supplying  a suitable ionic proportions in the mother brine. In terms of annual volume of product sold into the world market, Quaternary brine systems supply less than 15% The remainder comes from the mining of a variety of ancient solid-state potash sources. In the third and final article in this series, we shall discuss how and why the chemistry and hydrogeology of these ancient potash sources is mostly marine-fed and somewhat different from the continental hydrologies addressed so far.

    References

    Behrens, P., 2002, Industrial processing of Great Salt lake Brines by Great Salt Lake Minerals and Chemical Corporation, in D. T. Gywnne, ed., Great Salt Lake: A scientific, historical and economic overview, Utah Geological and Mineral Survey, Bulletin 116, p. 223-228.

    Bingham, C. P., 1980, Solar production of potash from brines of the Bonneville Salt Flats, in J. W. Gwynn, ed., Great Salt Lake; a scientific, history and economic overview. , v. 116, Bulletin Utah Geological and Mineral Survey, p. 229-242.

    Butts, D., 2002, Chemistry of Great Salt Lake Brines in Solar Ponds, in D. T. Gywnne, ed., Great Salt Lake: A scientific, historical and economic overview, Utah Geological and Mineral Survey, Bulletin 116, p. 170-174.

    Butts, D., 2007, Chemicals from Brines, Kirk-Othmer Encyclopedia of Chemical Technology, John Wiley & Sons, Inc., p. 784-803.

    Chao, L., P. Zicheng, Y. Dong, L. Weiguo, Z. Zhaofeng, H. Jianfeng, and C. Chenlin, 2009, A lacustrine record from Lop Nur, Xinjiang, China: Implications for paleoclimate change during Late Pleistocene: Journal of Asian Earth Sciences, v. 34, p. 38-45.

    Dong, Z., P. Lv, G. Qian, X. Xia, Y. Zhao, and G. Mu, 2012, Research progress in China's Lop Nur: Earth-Science Reviews, v. 111, p. 142-153.

    Felton, D., J. Waters, R. Moritz, D., and T. A. Lane, 2010, Producing Sulfate of Potash from Polyhalite with Cost Estimates, Gustavson Associates, p. 19.

    Hu, G., and N.-a. Wang, 2001, The sand wedge and mirabilite of the last ice age and their paleoclimatic significance in Hexi Corridor: Chinese Geographical Science, v. 11, p. 80-86.

    Huntington, E., 1907, Lop-Nor. A Chinese Lake. Part 1. The Unexplored Salt Desert of Lop: Bulletin of the American Geographical Society, v. 39, p. 65-77.

    Jones, B., D. Naftz, R. Spencer, and C. Oviatt, 2009, Geochemical Evolution of Great Salt Lake, Utah, USA: Aquatic Geochemistry, v. 15, p. 95-121.

    Lin, Y., L. Xu, and G. Mu, 2018, Differential erosion and the formation of layered yardangs in the Loulan region (Lop Nur), eastern Tarim Basin: Aeolian Research, v. 30, p. 41-47.

    Liu, C., W. Mili, J. Pengcheng, L. I. Shude, and C. Yongzhi, 2006, Features and Formation Mechanism of Faults and Potash-forming Effect in the Lop Nur Salt Lake, Xinjiang, China: Acta Geologica Sinica - English Edition, v. 80, p. 936-943.

    Liu, C.-A., H. Gong, Y. Shao, Z. Yang, L. Liu, and Y. Geng, 2016a, Recognition of salt crust types by means of PolSAR to reflect the fluctuation processes of an ancient lake in Lop Nur: Remote Sensing of Environment, v. 175, p. 148-157.

    Liu, C. L., M. L. Wang, P. C. Jiao, W. D. Fan, Y. Z. Chen, Z. C. Yang, and J. G. Wang, 2008, Sedimentary characteristics and origin of polyhalite in Lop Nur Salt Lake,Xinjiang: Mineral Deposits.

    Liu, C. L., J. F. Zhang, P. C. Jiao, and S. Mischke, 2016b, The Holocene history of Lop Nur and its palaeoclimate implications: Quaternary Science Reviews, v. 148, p. 163-175.

    Ma, C., F. Wang, Q. Cao, X. Xia, S. Li, and X. Li, 2008, Climate and environment reconstruction during the Medieval Warm Period in Lop Nur of Xinjiang, China: Chinese Science Bulletin, v. 53, p. 3016-3027.

    Ma, L., T. K. Lowenstein, B. Li, P. Jiang, C. Liu, J. Zhong, J. Sheng, H. Qiu, and H. Wu, 2010, Hydrochemical characteristics and brine evolution paths of Lop Nor Basin, Xinjiang Province, Western China: Applied Geochemistry, v. 25, p. 1770-1782.

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    Brine evolution and origins of potash ore salts: Primary or secondary? Part 1 of 3

    John Warren - Wednesday, October 31, 2018

    Introduction

    There is a dichotomy in mineralogical associations and precipitation series in both modern and ancient potash ore deposits. Interpretations of ancient potash ore mineralogies across time are generally tied to the evolution of the hydrochemical proportions in modern and ancient oceans. We have already discussed this in previous Salty Matters articles and will not repeat the details here (see August 10, 2015; July 31, 2018).

    At times in the past, such as in the Devonian and the Cretaceous, the world ocean was depleted in Mg and SO4 relative to the present-day ocean (Figure 1a). In the relevant literature, this has led to the application of the term MgSO4-depleted versus MgSO4-enriched oceans. In terms of brine evolution, this is related to the gypsum divide, with the term MgSO4-enriched used to describe the ocean chemistry of today and other times in the past, such as in the Permian, when MgSO4 bittern salts typify co-precipitates with sylvite/carnallite (Figure 1b).


    The validity of the ocean chemistry argument is primarily based on determinations of inclusion chemistries as measured in chevron halites (Figure 1a; Lowenstein et al., 2014). Inclusions in growth-aligned primary halite chevrons are assumed to preserve the chemical proportions in the ambient oceanic brine precipitating the halite. That is, the working assumption is that pristine aligned-halite chevrons have not been subject to significant diagenetic alteration once the salt was deposited and permeability was lost due to ongoing halite cementation in the shallow (eogenetic) subsurface realm.

    The same assumption as to the pristine nature of chevron halite is applied to outcomes of biological experiments where Permian archaeal/halobacterial life has been re-animated using ancient salt samples (Vreeland et al., 2000).

    Primary potash ore?

    But does the same assumption of pristine texturing across time also apply to the halite layers associated with the world’s potash ores? In my experience of subsurface potash ores and their textures, I have rarely seen primary-chevron halite interlayered with potash ore layers of either sylvite or carnallite. An obvious exception is the pristine interlayering of chevron halite and sylvite in the now-depleted Eocene potash ores of the Mulhouse Basin, France (Lowenstein and Spencer, 1990). There, the sylvite layers intercalate at the cm-scale with chevron halite, and the alternating layering is thought to be related to precipitation driven by temperature fluctuations in a series of shallow density-stratified meromictic brine lakes (documented in the third Salty Matters article).

    More typically, ancient potash ore textures are diagenetic and indicate responses to varying degrees of dissolution, brine infiltration and alteration. The simpler styles of brine infiltration consist of a background matrix dominated by cm-dm scale chevron halite layers that have been subject to dissolution and karstification during shallow burial. Resultant cm-dm scale voids typically retain a mm-thick selvedge of CaSO4 lathes and needles, followed by fill of the remaining void by varying amounts of sparry halite, carnallite and sylvite. This type of texture dominates Quaternary stratoid potash layers in the southern Qaidam Basin in China and Cretaceous carnallite-rich layers in the Maha Sarakham Fm in NE Thailand and southern China (Warren, 2016). Then there are the even more altered and recrystallised, but still bedded, textures in the potash ore zones of Devonian Prairie Evaporite of western Canada (Wardlaw, 1968) and potash layers in the Permian Basin in west Texas and New Mexico (Lowenstein, 1988; Holt and Powers, 2011). Beyond this level of diagenetic texturing are the flow-orientated and foliated structural textures of the Permian potash ores in potash mines in the diapiric Zechstein evaporites of Germany and Poland, the Kungurian diapirs of the Cis-Urals of Russia and the Devonian diapirs of the Pripyat Basin.

    And so, herein lies the main point of discussion for this and the next two Salty Matters articles, namely, what, where and when is(are) the mechanism(s) or association(s) of hydrochemical mechanisms that sufficiently concentrate or alter a brine’s chemistry to where it precipitates economic levels of a variety of potash salts, as either muriate of potash or sulphate of potash. Notably, there are no Quaternary-age solid-state ore systems that are mined for potash.

    In this article, we look at the main modern brine systems where muriate of potash (MOP) is produced economically by solar evaporation (Salar de Atacama, Chile; Qarhan sump, China; and the southern Basin of the Dead Sea). In the second article we will focus on sulphate of potash (SOP) production in Quaternary saline sumps (Great Salt Lake, USA and Lop Nur, China). In the third article we shall discuss depositional and diagenetic characteristics of solid-state potash ores some of the world’s more substantial deposits (e.g. Devonian of western Canada) and relate the observations of ancient potash texture to time-based evolution of potash precipitating brines, and subsequent alteration or the ore textures, which are typically driven by later cross-flushing by one or more pulses of diagenetically-evolved brines.


    Potash from brine in Salar de Atacama (MOP in a simple near-uniserial set of brine concentration pans)

    Potash production in Salar de Atacama is a byproduct of the output of lithium carbonate from shallow lake brines pumped into a series of solar concentration pans (Figure 2). The inflow feed to the concentrator pans comes from fields of brine wells extracting pore waters from the salt nucleus facies across the central and southern part of the Atacama saltflats (Figure 3a,b). However, Atacama pore brines are not chemically homogeneous across the salar (Alonso and Risacher, 1996; Risacher and Alonso, 1996; Carmona et al., 2000; Pueyo et al. 2017). The most common primary inflow brines to the Atacama sump are sulphate-rich (SO4/Ca > 1), but there are areas in the salt flat at the southern end of the playa, such as those near the Península de Chépica, where pore brines are richer in calcium (SO4/Ca < 1- Figure 4). These brines also contain elevated levels of lithium (Figure 3c; Risacher et al., 2003).


    Ion proportions in the natural salar inflows and pore waters are dominated by sodium and chloride, followed by potassium, then magnesium then sulphate in the more saline regions of the salar sump (Figure 4; Lowenstein and Risacher, 2009). In addition, owing to the progressive reduction of porosity with depth, driven mainly by diagenetic halite cementation, the pore brine in the upper 40 meters of the salar sediment column accumulates by advection in the area of greatest porosity, i.e., in this top 40 m of sediments of the salt flat at the southern end of Atacama (Pueyo et al., 2017). When pumped from the hosting salar sediments into the concentrator pans, the final brines contain elevated levels of lithium chloride (≈ 6000 ppm). These lithium-enriched acidic waters are then pumped to a nearby industrial plant and processed to obtain lithium carbonate as the main commercial product.

    In a benchmark paper, Pueyo et al. (2017) document the brine evolution and products recovered in the solar pans of Rockwood Lithium GmbH (Figure 2b; formerly Sociedad Chilena del Litio) in the Península de Chépica. There, a bittern paragenesis of salts precipitates that is mostly devoid of magnesium sulphate salt due to the low levels of sulphate attained in the various concentator pans via widespread precipitation of gypsum in the early concentrator pans (Figures 4, 5).


    The depletion of sulphate levels in the early concentrators is done via artificial manipulation of ionic proportions in the feeders. Without alteration of the ionic proportion in halite-stage brines, the evaporation of the saltflat brine feeds, which are rich in sulfate, would result in assemblages that, in addition to potassic chlorides, would contain contain problematic magnesium sulfates (such as schoenite, kainite, glaserite as in Great Salt Lake). The presence of such sulphate salts and ions in the liquor feeding the lithium carbonate plant would complicate the lithium carbonate extraction process. So the aim in the Atacama pans is to remove most of the sulphate via constructing a suitably balanced chemistry in the early concentrator brine stage (compare ionic proportion in sulphate between early and end-stage bitterns, as illustrated in Figures 4 and Figure 5a).


    Such a sylvite/carnallite brine paragenesis, sans sulphate (as seen in Figure 5), is similar to that envisaged as the feed chemistry for ancient Mg-sulphate-free marine potash deposits (Braitsch, 1971). That is, as the brines pass through the concentrators, with successive pans transitioning to higher salinities, the potash salts carnallite and sylvite precipitate, without the complication of the widespread magnesium sulphate salts, which complicate the processing of modern marine-derived bitterns. Such MgSO4 double-salts typify SOP production in the Ogden Salt flats, with their primary feed of sulphate-rich Great Salt Lake waters. Relative proportions of sulphate are much higher in the Great Salt Lake brine feed (see article 2 in this series).


    Once the balance is accomplished by mixing a Ca-rich brine from further up the concentration series, with the natural SO4-brine in an appropriate ratio, the modified brines are then pumped and discharged into the halite ponds of the saltwork circuit (ponds number 17 and 16, as seen in Figures 5 and 6). In these ponds, halite precipitates from the very beginning with small amounts of accessory gypsum as brines are saturated with both minerals. Subsequently, the brines are transferred to increasingly smaller ponds where halite (ponds 15 and 14), halite and sylvite (pond 13), sylvite (ponds 12, 11 and 10), sylvite and carnallite (pond 9), carnallite (pond 8), carnallite and bischofite (pond 7), bischofite (ponds 6, 5 and 4), bischofite with some lithium-carnallite [LiClMgCl26H2O] (pond 3), and lithium-carnallite (ponds 2 and 1) precipitate. The brines of the last ponds (R-1 to R-3), whose volumes undergo a reduction to 1/50th of the starting volume, are treated at the processing factory to obtain lithium carbonate as the main commercial product.

    As documented in Pueyo et al. (2017), the average daily temperature in the Salar de Atacama ranges between 22 °C in February and 8 °C in July, with a maximum oscillation of approximately 14 °C. Wind speed ranges daily from< 2 ms−1 in the morning to 15 ms−1 in the afternoon. Rainfall in the area of the salt flat corresponds to that of a hyperarid desert climate with an annual average, for the period 1988–2011, of 28 mm at San Pedro de Atacama, 15.1 mm at Peine and 11.6 mm at the lithium saltworks, in the last case ranging between 0 and 86 mm for individual years. The adjoining Altiplano to the east has an arid climate with an average annual rainfall of approximately 100 mm. The average relative humidity in the saltpan area, for the period 2006–2011, is 19.8% with a maximum around February (27%) and a minimum in October (15%) and with a peak in the morning when it may reach 50%. The low relative humidity and the high insolation (direct radiation of 3000 kWh m−2 yr−1) in the salt flat increase the efficiency of solar evaporation, giving rise to the precipitation and stability of very deliquescent minerals such as carnallite and bischofite. The average annual evaporation value measured in the period 1998–2011, using the salt flat interstitial brine, is approximately 2250 mm with a peak in December–January and a minimum in June–July. This cool high-altitude hyperarid climatic setting, where widespread sylvite and carnallite accumulates on the pan floor, is tectonically and climatically distinct from the hot-arid subsealevel basinwide desert seep settings envisaged for ancient marine-fed potash basins (as discussed in the upcoming third article in this series).

    MOP from brine Dabuxum/Qarhan region, Qaidam Basin, China

    The Qarhan saltflat/playa is now the largest hypersaline sump within the disaggregated lacustrine system that makes up the hydrology of Qaidam Basin, China (Figure 7a). The Qaidam basin sump has an area of some 6,000 km2, is mostly underlain by bedded Late Quaternary halite. Regionally, the depression is endorheic, fed by the Golmud, Qarhan and Urtom (Wutumeiren) rivers in the south and the Sugan River in the north, and today is mostly covered by a layered halite pan crust. Below, some 0 to 1.3m beneath the playa surface, is the watertable atop a permanent hypersaline groundwater brine lens (Figure 7b).


    The southern Qaidam sump entrains nine perennial salt lakes: Seni, Dabiele, Xiaobiele, Daxi, Dabuxum (Dabsan Hu), Tuanjie, Xiezuo and Fubuxum north and south lakeshore (Figure 7). Dabuxum Lake, which occupies the central part of the Qarhan sump region, is the largest of the perennial lakes (184 km2; Figures 7b, 8a). Lake water depths vary seasonally from 20cm to 1m and never deeper than a metre, even when flooded. Salt contents in the various lakes range from 165 to 360 g/l, with pH ranging between 5.4 and 7.85. Today the salt plain and pans of the Qarhan playa are fed mostly by runoff from the Kunlun Mountains (Kunlun Shan), along with input from a number of saline groundwater springs concentrated along a fault trend defining an area of salt karst along the northern edge of the Dabuxum sump, especially north of Xiezuo Lake (Figure 8a).


    The present climate across the Qaidam Basin is cool, arid to hyperarid (BWk), with an average yearly rainfall of 26 mm, mean annual evaporation is 3000–3200 mm, and a yearly mean temperature 2-4° C in the central basin (An et al., 2012). The various salt lakes and playas spread across the basin and contain alternating climate-dependent evaporitic sedimentary sequences. Across the basin the playa sumps are surrounded by aeolian deposits and wind-eroded landforms (yardangs). In terms of potash occurrence, the most significant region in the Qaidam Basin is the Qarhan sump or playa (aka Chaerhan Salt Lake), which occupies a landscape low in front of the outlets of the Golmud and Qarhan rivers (Figure 7a, b). Overall the Qaidam Basin displays a typical exposed lacustrine geomorphology and desert landscape, related to increasing aridification in a cool desert setting. In contrast, the surrounding elevated highlands are mostly typified by a high-alpine tundra (ET) Köppen climate.


    Bedded and displacive salts began to accumulate in the Qarhan depression some 50,000 years ago (Figure 9). Today, outcropping areas of surface salt crust consist of a chaotic mixture of fine-grained halite crystals and mud, with a rugged, pitted upper surface (Schubel and Lowenstein, 1997; Duan and Hu, 2001). Vadose diagenetic features, such as dissolution pits, cavities and pendant cements, form wherever the salt crust lies above the watertable. Interbedded salts and siliciclastic sediments underlying the crust reach thicknesses of upwards of 70m (Kezao and Bowler, 1986).

    Bedded potash, as carnallite, precipitates naturally in transient volumetrically-minor lake strandzone (stratoid) beds about the northeastern margin of Lake Dabuxum (Figure 8a) and as cements in Late Pleistocene bedded deposits exposed in and below nearby Lake Tuanje in what is known as the sediments of the Dadong ancient lake (Figure 8b). Ongoing freshened sheetflow from the up-dip bajada fans means the proportion of carnallite versus halite in the evaporite unit increases with distance from the Golmud Fan across, both the layered (bedded) and stratoid (cement) modes of occurrence.

    At times in the past, when the watertable was lower, occasional meteoric inflow was also the driver for the brine cycling that created the karst cavities hosting the halite and carnallite cements that formed as prograde cements during cooling of the sinking brine (Figure 9). Solid bedded potash salts are not present in sufficient amounts to be quarried, and most of the exploited potash resource resides in interstitial brines that are pumped and processed using solar ponds.

    Modern halite crusts in Qarhan playa contain the most concentrated brine inclusions of the sampled Quaternary halites, suggesting that today may be the most desiccated period in the Qarhan-Tuanje sump recorded over the last 50,000 years (values in the inset in figure 9 were measured on clear halite-spar void-fill crystals between chevrons). Inclusion measurements from these very early diagenetic halite show they formed syndepositionally from shallow groundwater brines and confirm the climatic record derived from adjacent primary (chevron) halite. The occurrence of carnallite-saturated brines in fluid inclusions in the diagenetic halite in the top 13 m of Qarhan playa sediments also imply a prograde diagenetic, not depositional, origin of carnallite, which locally accumulated in the same voids as the more widespread microkarst halite-spar cements.

    Today, transient surficial primary carnallite rafts can accumulate along the northern strandline of Lake Dabuxum (Figure 9; Casas, 1992; Casas et al., 1992). Compositions of fluid inclusions in the older primary (chevron) halite beds hosting carnallite cements in the various Qarhan salt crusts represent preserved lake brines and indicate relatively wetter conditions throughout most of the Late Pleistocene (Yang et al., 1995). Oxygen isotope signatures of the inclusions record episodic freshening and concentration during the formation of the various salt units interlayered with lacustrine muds. Desiccation events, sufficient to allow halite beds to accumulate, occurred a number of times in the Late Quaternary: 1) in a short-lived event ≈ 50,000 ka, 2) from about 17 - 8,000 ka, and 3) from about 2,000 ka till now (Figure 9).

    The greatest volume of water entering Dabuxum Lake comes from the Golmud River (Figure 7b). Cold springs, emerging from a narrow karst zone some 10 km to the north of the Dabuxum strandline and extending hundreds of km across the basin, also supply solutes to the lake. The spring water discharging along this fault-defined karst zone is chemically similar to hydrothermal CaCl2 basin-sourced waters as defined by Hardie (1990), and are interpreted as subsurface brines that have risen to the surface along deep faults to the north of the Dabuxum sump (Figure 9, 10; Spencer et al., 1990; Lowenstein and Risacher, 2009). Depths from where the Ca–Cl spring waters rise is not known. Subsurface lithologies of the Qaidam Basin in this region contains Jurassic and younger sediments and sedimentary rock columns, up to 15 km thick, which overlie Proterozoic metamorphic rocks (Wang and Coward, 1990).


    Several lakes located near the northern karst zone (Donglin, North Huobusun, Xiezhuo, and Huobusun) receive sufficient Ca–Cl inflow, more than 1 part spring inflow to 40 parts river inflow, to form mixtures with chemistries of Ca equivalents > equivalents HCO3 + SO4 to create a simple potash evaporation series (this is indicated by the Ca-Cl trend line in Figure 10a). With evaporation such waters, after precipitation of calcite and gypsum, evolve into Ca–Cl-rich, HCO3–SO4-poor brines (brines numbered 5, 7-12 in figure 11a).


    Dabuxum is the largest lake in the Qarhan region, with brines that are Na–Mg–K–Cl dominant, with minor Ca and SO4 (Figure 10d, 11a). These brines are interpreted by Lowenstein and Risacher (2009) to have formed from a mix of ≈40 parts river water to 1 part spring inflow, so that the equivalents of Ca ≈ equivalents HCO3 + SO4 (Figure 10b). Brines with this ratio of river to spring inflow lose most of their Ca, SO4, and HCO3 after precipitation of CaCO3 and CaSO4, and so form Na–K–Mg–Cl brines capable of precipitating carnallite and sylvite (Figure 11a). This chemistry is similar to that of ancient MgSO4-depleted marine bitterns (Figure 1)

    The chemical composition of surface brines in the various lakes on the Qarhan Salt plain vary and appear to be controlled by the particular blend of river and spring inflows into the local lake/playa sump. In turn, this mix is controlled geographically by proximity to river mouths and the northern karst zone. Formation of marine-like ionic proportions in some lakes, such as Tuanje, Dabuxum and ancient Dadong Lake, engender bitterns suitable for the primary and secondary precipitation of sylvite/carnallite (Figures 10b-d; 11a). The variation in the relative proportion of sulphate to chloride in the feeder brines is a fundamental control on the suitability of the brine as a potash producer.

    Figure 11b clearly illustrates sulphate to chloride variation in pore waters in the region to the immediate north and east of Dabuxum Lake. Brine wells in the low-sulphate area are drawn upon to supply feeder brines to the carnallite precipitating ponds. The hydrochemistry of this region is a clear indication of the regional variation in the sump hydrochemistry (Duan and Hu, 2001), but also underlines why it is so important to understand pore chemistry, and variations in aquifer porosity and permeability, when designing a potash plant in a Quaternary saline setting.

    Compared to the MOP plant in Atacama, there as yet no lithium carbonate extraction stream to help ameliorate costs associated with carnallite processing. Lithium levels in the Qaidam brines, whilee levated, are much lower than in the Atacama brine feeds. Regionally, away from the Tuanje-Dadong area, most salt-lake and pore brines in the Qaidam flats are of the magnesium sulphate subtype and the ratio of Mg/Li can be as high as 500. With such brine compositions, the chemical precipitation approach, which is successfully applied to lithium extraction using low calcium and magnesium brines (such as those from Zabuye and Jezecaka Lake on the Tibetan Plateau and in the Andean Altiplano), would consume a large quantity of chemicals and generate a huge amount of solid waste. Accordingly, brine operations in the Qarhan region are focused on MOP production from a carnallitite slurry using extraction techniques similar to those utilised in the Southern Dead Sea. But owing to its cooler climate compared to the Dead Sea sump, the pond chemistry is subject to lower evaporation rates, higher moisture levels in the product, and a longer curing time.

    Potash in the Qarhan region is produced by the Qinghai Salt Lake Potash Company, which owns the 120-square-kilometer salt lake area near Golmud (Figure 7). The company was established and listed on the Shenzhen Stock Exchange in 1997. Currently, it specialises in the manufacture of MOP from pore brines pumped from appropriate low-sulphate regions in the lake sediments (Figure 12). The MOP factory processes a carnallite slurry pumped from pans using a slurry processing stream very similar to the dual process stream utilized in the pans of the Southern Basin in the Dead Sea and discussed in the next section.

    The final potash product in the Qaidam sump runs 60-62% K2O with >2% moisture and is distributed under the brand name of “Yanqiao.” With annual production ≈3.5 million tonnes and a projected reserve ≈ 540 million tonnes, the company currently generates 97% of Chinese domestic MOP production. However, China’s annual agricultural need for potash far outpaces this level of production. The company is jointly owned by Qinghai Salt Lake Industry Group and Sinochem Corporation and is the only domestic producer of a natural MOP product.

    Dead Sea Potash (MOP operation in the Southern Basin)

    The Dead Sea water surface defines what is the deepest continental position (-417 m asl) on the earth’s current terrestrial surface. In the Northen Basin is our only modern example of bedded evaporitic sediments (halite and gypsum) accumulating on the subaqueous floor of a deep brine body, where water depths are hundreds of metres (Warren, 2016). This salt-encrusted depression is 80 km long and 20 km wide, has an area of 810 km2, is covered by a brine volume of 147 km3 and occupies the lowest part of a drainage basin with a catchment area of 40,650 km2 (Figure 13a). However, falling water levels in the past few decades mean the permanent water mass now only occupies the northern part of the lake, while saline anthropogenic potash pans occupy the Southern Basin, so that the current perennial “Sea” resides in the Northern Basin is now only some 50 km long (Figure 13b).


    Rainfall in the region is 45 to 90 mm, evaporation around 1500 mm, and air temperatures between 11 and 21°C in winter and 18 to 40°C in summer, with a recorded maximum of 51°C. The subsiding basin is surrounded by mountain ranges to the east and west, producing an orographic rain shadow that further emphasises the aridity of the adjacent desert sump. The primary source of solutes in the perennial lake is ongoing dissolution of the halokinetic salts of the Miocene Sedom Fm (aka Usdum Fm) a marine evaporite unit that underlies the Dead Sea and approaches the surface in diapiric structures beneath the Lisan Straits and at Mt. Sedom (Garfunkel and Ben-Avraham, 1996).

    A series of linked fractionation ponds have been built in the Southern Basin of the Dead Sea to further concentrate pumped Dead Sea brine to the carnallite stage (Figure 13). On the Israeli side this is done by the Dead Sea Works Ltd. (owned by ICL Fertilisers), near Mt. Sedom, and by the Arab Potash Company (APC) at Ghor al Safi on the Jordanian side. ICL is 52.3% owned by Israel Corporation Ltd.(considered as under Government control), 13.6% shares held by Potash Corporation of Saskatchewan and 33.6% shares held by various institutional investors and the general public (33.64%). In contrast, PotashCorp owns 28% of APC shares, the Government of Jordan 27%, Arab Mining Company 20%, with the remainder held by several small Middle Eastern governments and a public float that trades on the Amman Stock Exchange. This gives PotashCorp control on how APC product is marketed, but it does not control how DSW product is sold.

    In both the DSW and APC brine fields, muriate of potash is extracted by processing carnallitite slurries, created by sequential evaporation in a series of linked, gravity-fed fractionation ponds. The inflow brine currently pumped from the Dead Sea has a density of ≈1.24 gm/cc, while after slurry extraction the residual brine, with a density of ≈1.34 gm/cc, is pumped back into the northern Dead Sea basin water mass. The total area of the concentration pans is more than 250 km2, within the total area of 1,000 km2, which is the southern Dead Sea floor. The first stage in the evaporation process is pumping of Dead Sea water into header ponds and into the gravity-fed series of artificial fractionation pans that now cover the Southern Basin floor. With the ongoing fall of the Dead Sea water level over the past 60 years, brines from the Northern Basin must be pumped higher and over further lateral distances. This results in an ongoing need for more powerful brine pumps and an increasing problem with karst dolines related to lowered Dead Sea water levels. Saturation stages of the evolving pan brines are monitored and waters are moved from pan to pan as they are subject to the ongoing and intense levels of natural solar evaporation (Figure 13b, c; Karcz and Zak, 1987).

    The artificial salt ponds of the Dead Sea are unusual in that they are designed to trap and discard most of the halite precipitate rather than harvest it. Most other artificial salt ponds around the world are shallow pans purpose-designed as ephemeral water-holding depressions that periodically dry out so that salts can be scrapped and harvested. In contrast, the Dead Sea halite ponds are purpose-designed to be permanently subaqueous and relatively deep (≈4m). Brine levels in the ponds vary by a few decimetres during the year, and lowstand levels generally increase each winter when waste brine is pumped back into the northern basin.

    As the Dead Sea brine thickens, minor gypsum, then voluminous halite precipitates on the pan floor in the upstream section of the concentration series, where the halite-precipitating-brines have densities > 1.2 gm/cc (Figure 13c). As the concentrating brines approach carnallite-precipitating densities (around 1.3 gm/cc), they are allowed to flow into the carnallite precipitating ponds (Figure 13c). Individual pans have areas around 6-8 km2 and brine depths up to 2 metres. During the early halite concentration stages, a series of problematic halite reefs or mushroom polygons can build to the brine surface and so compartmentalise and entrap brines within isolated pockets enclosed by the reefs. This hinders the orderly downstream progression of increasingly saline brines into the carnallite ponds, with the associated loss of potash product.


    When the plant was first designed, the expectation was that halite would accumulate on the floor of the early fractionation ponds as flat beds and crusts, beneath permanent holomictic brine layers. The expected volume of salt was deposited in the pans each year (Talbot et al., 1996), but instead of accumulating on a flat floor aggrading 15-20 cm each year, halite in some areas aggraded into a series of polygonally-linked at-surface salt reefs (aka salt mushrooms). Then, instead of each brine lake/pan being homogenized by wind shear across a single large subaqueous ponds, the salt reefs separated the larger early ponds into thousands of smaller polygonally-defined inaccessible compartments, where the isolated brines developed different compositions (Figure 14). Carnallitite slurries crystallised in inter-reef compartments from where it could not be easily harvested, so large volumes of potential potash product were locked up in the early fractionation ponds (Figure 14a, b). Attempts to drown the reefs by maintaining freshened waters in the ponds during the winters of 1984 and 1985 were only partly successful. The current approach to the salt reef problem in the early fractionation ponds is to periodically breakup and remove the halite reefs and mushrooms by a combination of dredging and occasional blasting (Figure 14c).


    Unlike seawater feeds to conventional marine coastal saltworks producing halite with marine inflow salinities ≈35‰, the inflow brine pumped into the header ponds from the Dead Sea already has a salinity of more than 300‰ (Figure 15). Massive halite precipitation occurs quickly, once the brine attains a density of 1.235 (≈340‰) and reaches a maximum at a density of 1.24 (Figure 13c). Evaporation is allowed to continue in the initial halite concentrator ponds until the original water volume pumped into the pond has been halved. Concentrated halite-depleted brine is then pumped through a conveyance canal into a series of smaller evaporation ponds where carnallite, along with minor halite and gypsum precipitates (Figure 13c). Around 300–400 mm of carnallite salt slurry is allowed to accumulate in the carnallite ponds, with 84% pure carnallite and 16% sodium chloride as the average chemical composition (Figure 6a; Abu-Hamatteh and Al-Amr, 2008). The carnallite bed is harvested (pumped) from beneath the brine in slurry form and is delivered through corrosion-resistant steel pipes to the process refineries via a series of powerful pumps.

    This carnallitite slurry is harvested using purpose-specific dredges floating across the crystalliser ponds. These dredges not only pump the slurry to the processing plant but also undertake the early part of the processing stream. On the dredge, the harvested slurry is crushed and size sorted, with the coarser purer crystals separated for cold crystallisation. The remainder is slurried with the residual pan brine and then further filtered aboard the floating dredges. At this stage in the processing stream the dredges pipe the treated slurries from the pans to the refining plant.


    On arrival at the processing plant, raw product is then used to manufacture muriate of potash, salt, magnesium chloride, magnesium oxide, hydrochloric acid, bath salts, chlorine, caustic soda and magnesium metal (Figure 16a). Residual brine after carnallitite precipitation contain about 11-12 g/l bromide and is used for the production of bromine, before the waste brine (with a density around 1.34 gm/cc) is returned to the northern Dead Sea water mass. The entire cycle from the slurry harvesting to MOP production takes as little as five hours.

    In the initial years of both DSW and APC operations, MOP was refined from the carnallite slurry via hot leaching and flotation. In the coarser-crystalline carnallitite feed, significant volumes of sylvite are now produced more economically in a cold crystallisation plant (Figure 16b). The cold crystallisation process takes place at ambient temperature and is less energy-intensive than the hot crystallisation unit. The method also consumes less water but requires a higher and more consistent grade of carnallite feed (Mansour and Takrouri, 2007; Abu-Hamatteh and Al-Amr, 2008). Both hot (thermal) and cold production methods can be utilized in either plant, depending on the quality of the slurry feed.

    Sylvite is produced via cold crystallisation using the addition of water to incongruently dissolve the magnesium chloride from the crystal structure. If the carnallite slurry contains only a small amount of halite, the solid residue that remains after water flushing is mostly sylvite. As is shown in Figure 16b, if the MgCl2 concentration is at or near the triple-saturation point (the point at which the solution is saturated with carnallite, NaCl, and KCl), the KCl solubility is suppressed to the point where most of it will precipitate as sylvite. For maximum recovery, the crystallising mixture must be saturated with carnallite at its triple-saturation point. If the mixture is not saturated, for example, it contains higher levels of NaCl, then more KCl will dissolve during the water flushing of the slurry. Industrially, the cold crystallizers are usually fed with both coarse and fine carnallite streams, such that 10% carnallite remains in the slurry, this can be achieved by adjusting addition of process water (Mansour and Takrouri, 2007).

    Successful cold crystallisation depends largely on a consistent high-quality carnallite feed. If a large amount of halite is present in the feed slurry, the resulting solid residue from cold crystallisation is sylvinite, not sylvite. This needs to be further refined by hot crystallisation, a more expensive extraction method based on the fact that the solubility of sylvite varies significantly with increasing temperature, while that of salt remains relatively constant (Figure 16c). As potash brine is hot leached from the sylvinite, the remaining halite is filtered off, and the brine is cooled under controlled conditions to yield sylvite.

    Residual brine from the crystallisation processes then undergoes electrolysis to yield chlorine, caustic soda (sodium hydroxide) and hydrogen. Chlorine is then reacted with brine filtered from the pans to produce bromine. The caustic soda is sold, and the hydrogen is used to make bromine compounds, with the excess being burnt as fuel. Bromine distilled from the brine is sold partly as elemental bromine, and partly in the form of bromine compounds produced in the bromine plant at Ramat Hovav (near Beer Sheva). This is the largest bromine plant in the world, and Israel is the main exporter of bromine to Europe. About 200,000 tons of bromine are produced each year.

    Residual magnesium chloride-rich solutions created by cold crystallisation are concentrated and sold as flakes for use in the chemical industry and for de-icing (about 100,000 tons per year) and dirt road de-dusting. Part of the MgCl2 solution produced is sold to the nearby Dead Sea Periclase plant (a subsidiary of Israel Chemicals Ltd.). At this plant, the brine is decomposed thermally to give an extremely pure magnesium oxide (periclase) and hydrochloric acid. In Israel, Dead Sea Salt Work’s (DSW) production has risen to more than 2.9 Mt KCl since 2005, continuing a series of increments and reflecting an investment in expanded capacity, the streamlining of product throughput in the mill facilities, and the amelioration of the effects salt mushrooms, and increased salinity of the Dead Sea due to extended drought conditions (Figure 17).

    On the other side of the truce line in Jordan, the Arab Potash Co. Ltd. (APC) output rose to 1.94 Mt KCl in 2010 The APC plant now has the capacity to produce 2.35 Mt KCl and like the DSW produces bromine from bittern end brines. Early in the pond concentration stream, APC also has to remove salt mushrooms from its ponds, a process which when completed can increase carnallite output by over 50,000 t/yr. Currently, APC is continuing with an expansion program aimed at increasing potash capacity to 2.5 Mt/yr.


    MOP brines and Quaternary climate

    As mentioned in the introduction, exploited Quaternary potash deposits encompass both MOP and SOP mineral associations across a range of climatic and elevation settings. This article focuses on the three main MOP producing examples, the next deals with SOP Quaternary producers (Great Salt Lake, USA and Lop Nur, China). Interestingly, both sets of Quaternary examples are nonmarine brine-fed depositional hydrologies. All currently-active economic potash plants hosted in Quaternary systems do not mine a solid product but derive their potash solar evaporation of pumped hypersaline lake brines. For MOP processing to be economic the sulphate levels in the brines held in bittern-stage concentrator pans must be low and Mg levels are typically high, so favoring the precipitation of carnallite over sylvite in all three systems.

    In Salar de Atacama the low sulphate levels in the bittern stage is accomplished by artificially mixing a CaCl2 brine from further up the evaporation stream with a less saline more sulphate-enriched brine. The mixing proportions of the two brine streams aims to maximise the level of extraction/removal of CaSO4 in the halite pans prior to the precipitation of sylvite and carnallite. In the case of the pans in the Qarhan sump there is a similar but largely natural mixing of river waters with fault-fed salt-karst spring waters in a ratio of 40:1 that creates a hybrid pore brine with a low sulphate chemistry suitable for the precipitation of both natural and pan carnallite. In the case of the Dead Sea brine feed, the inflowing Dead Sea waters are naturally low in sulphate and high in magnesium. The large size of this natural brine feed systems and its homogeneous nature allows for a moderate cost of MOP manufacture estimated in Warren 2016, chapter 11 to be US$ 170/tonne. The Qarhan production cost is less ≈ US$ 110/tonne but the total reserve is less than in the brine system of the Dead Sea. In Salar de Atacama region the MOP cost is likely around US$ 250-270/tonne, but this is offset by the production of a bischofite stage brine suitable for lithium carbonate extraction.

    Outside of these three main Quaternary-feed MOP producers there are a number of potash mineral occurrences in intermontane depressions in the high Andes in what is a high altitude polar tundra setting (Koeppen ET), none of which are commercial (Figure 18a). Similarly, there a number of non-commercial potash (SOP) mineral and brine occurrences in various hot arid desert regions in Australia, northern Africa and the Middle East (Koeppen BWh) that we shall look at in the next article. In the Danakhil depression there is the possibility of a future combined MOP/SOP plant (see Salty Matters April 19, 2015; April 29, 2015; May 1, 2015; May12, 2015 and Bastow et al., 2018). In the Danakhil it is important to distinguish between the current non-potash climate (BWh - Koeppen climate) over the Dallol saltflat in Ethiopia, with its nonmarine brine feed and the former now-buried marine fed potash (SOP)/halite evaporite system. The latter is the target of current exploration efforts in the basin, focused on sediments now buried 60-120m below the Dallol saltflat surface. Nowhere in the Quaternary are such dry arid desert climates (BWh) associated with commercial accumulations of potash minerals.


    Climatically most commercial potash brine systems in Quaternary-age sediments are located in cooler endorheic intermontane depressions (BWk, BSk) or in the case of the Dead Sea an intermontane position in the sump of the Dead Sea, the deepest position of any continental landscape on the earth’s surface (-417 msl). The association with somewhat cooler and or less arid steppe climates implies a need for greater volumes of brine to reside in a landscape in order to facilitate the precipitation of significant volumes of potash bitterns (Figure 18a,b).

    In summary, all three currently economic Quaternary MOP operations are producing by pumping nonmarine pore or saline lake brines into a series of concentrator pans. The final bittern chemistry in all three is a low-sulphate liquour, but with inherently high levels of magnesium that favors the solar pan production of carnallite over sylvite that is then processed to produce the final KCl product. The brine chemistry in all three examples imitates the ionic proportions obtained when evaporating a ancient sulphate-depleted seawater (Figure 1). The next article will discuss the complexities (the double salt problem at the potash bittern stage when concentrating a more sulphate-enriched mother brine.

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    Life in modern Deepsea Hypersaline Lakes and Basins - DHALs and DHABs

    John Warren - Sunday, September 30, 2018

     


    Introduction

    Exuded salt karst brine on the deep ocean floor has a much higher density that the overlying seawater and so if there is an ongoing supply it tends to pond in seafloor lows (Figure 1a). The longterm character (hydrological stability over hundreds to thousands of years) of such density-stratified brine lakes, which form the centrepieces in deepsea hypersaline anoxic basins (DHAB), facilitate longterm ecologic niche sthe tability. The upper surface of a brine lake is marked by a halocline, which typically defines one or more nutrient, thermal and salinity interfaces (Figure 1b). There a light-independent chemosynthetic seep and lake biota can grow and flourish (Figure 1a). Escaping subsurface brines can entrain both hydrocarbons (mostly methane) and H2S, which are nutrients in the base of the chemosynthetic food chain. The salinity layering created by the halocline can be positioned as ; 1) a pelagic biotal interface, or 2) a brine lake edge (or shore) interface or 3) out in the lake the brine column base (i.e. a hypersaline-sediment interface) (Figure 1b).

    In other places on a deep seafloor, the escaping salt-karst brines, with entrained methane and H2S, can form diffuse outflow or seep areas, without ever developing into a free-standing brine lake (position 4 in Figure 1a). Highly specialised chemosynthetic communities tend to dolonise the resulting density and salinity-stratified interfaces. And so, some chemosynthetic communities occupy a halocline interface in a pelagic position atop an open brine lake, while others inhabit a benthic position where the halocline intersects the deep seafloor (Figure 1). Anoxic hypersaline brine can also pond on the shallow seafloor in high latitude regions where the formation of sea ice create cryogenic brines (Kvitek et al, 1998). But this style of cryogenic seaflooor brine lake is more ephemeral and is not tied to major evaporite deposits, so is not considered further.

    Two groups of megafauna with symbiotic methanotrophic or thiotrophic bacteria dominate chemoosynthetic communities in the salt-floored Gulf of Mexico: 1) bivalves, including bathymodiolin mussels and multiple families of clams and 2) vestimentiferan tubeworms in the polychaete family Siboglinidae. Both the vestimentiferan siboglinids and clams harbour microbial endosymbionts that utilise sulphide as an energy source, whereas different species of bathymodiolin mussels harbour either methanotrophic, thiotrophic, or both, types of symbionts (Figure 2).

    Along the brine pool edge in the Gulf of Mexico

    Hence, the mussel-tubeworm dominated brine-lake edge and seep biostromes in the Gulf of Mexico are dependent on chemosynthesising microbes as a food source. This community is the cold-water counterpart to warm-water chemosynthetic hydrothermal communities flourishing in high temperature waters the vicinity of black smoker vents (MacDonald, 1992; MacDonald et al., 2003). In both settings, it is methane and sulphide, not light, that provides the than DHALs energy source for the bacteria and archaea that make up the base of the chemosynthetic food chain.

    Methanotrophic bacteria live symbiotically on a seep mussel’s gills, taking in methane and converting it to nutrients that nourish the mussels. The seep mussels (Bathymodiolus childressi and Calyptogena ponderosa continually waft methane-rich water through their gills to help their chemo-autotrophic bacterial symbionts grow and periodically harvest some of the excess growth. Their lifestyle means that seep mussels need to live near a supply of dissolved gas, so they can inhabit isolated seep outflows on the deep seafloor where gas is bubbling out, including the edges of mud volcano pools, but do best about the more stable and relatively quiescent edges of methane-saturated brine pools and lakes.


    There they grow as a fringe to the brine pool, and exist about the pool rim, wherever they can keep their syphons above the halocline Figure 2a-d). They tend to construct a biogenic edge (biostrome) to the brine pool atop with sediment piles generally cemented by methanogenic calcite. Such rims typically extend some 5-10 metres behind the pool edge (Figure 2a; Smith et al., 2000). The inner edge of the mussel biostrome is elevated only a few centimetres from the surface of the pool and is distinguishable by an abundance of smaller individuals, present in high densities (Figure 2b). At the outer edge of the mussel biostrome, there is a high frequency of disarticulated shells and low densities of still living larger individuals.

    Also living atop seafloor seeps and about some brine pools are knots and clusters of chemosynthetic polychaete tubeworms (Figures 2c, 3; Lamellibrachia luymesi and Seepiophila jonesi). Individual tubeworms (aka seep beard-worms) in a colony can be up to 2.5 m long with a microbe-dependent metabolism evolved to exploit the abundant H2S and methane seeping through the seafloor. Tubeworm colonies grow as rims and clumps atop H2S seeps, as at Bush Hill on the floor of the Gulf of Mexico (Figure 3a; Reilly et al., 1996; Dattagupta et al., 2006; McMullin et al., 2010). Tubeworm “bushes” in cold seep regions of the Gulf of Mexico are typically rooted in the H2S-rich muds (Figure 3b). Growing individual tubes actively extend down into the H2S-rich mud as well as up into the O2-rich water column giving the cluster a morphology similar to a tree or shrub. Their “roots” extend into the earth, while “branches” extend above. Continuing the plant analogy, it seems that tubeworm shrubs absorb H2S through their “roots” and O2 through their “branches” (Freytag et al., 2001; Bergquist et al., 2003). As a group, seep tubeworms are related to the giant rift tubeworm (Riftia pachptila), which inhabits active hydrothermal seeps in active seafloor rifts.


    Via a specialised haemoglobin molecule, vestimentiferan tubeworms in the Gulf of Mexico provide H2S and O2 as nutrients to sulphur-oxidising bacteria living symbiotically in trophosome structures, which extend for up to 75% of the length of each tubeworm. Unlike hydrothermal tubeworms such as Riftia pachptila that grow to lengths of more than 2 metres in less than two years, Lamellibrachia luymesi grow very slowly for most of their lives. It takes from 170 to 250 years to grow to 2 meters in length, making them perhaps the longest living known invertebrate species (Bergquist et al., 2000). With five or six species currently known to flourish there, the brine-fed cold seeps of the Gulf of Mexico host the highest biodiversity of vestimentiferan siboglinid tubeworms worldwide.

    There is a time-based evolution in the biotal make-up of chemosynthetic communities in the Gulf of Mexico (Glover et al., 2010 and references therein). The earliest stage of a cold seep is characterised by a high seepage rate and the release of large amounts of biogenic and thermogenic methane, H2S and oil (Sassen et al., 1994). As authigenic carbonates with specific negative δ13C values precipitate as a metabolic byproduct of microbial methanogenesis, they provide a necessary stable substrate for the settlement of larval vestimentiferans and seep mussels. These seep communities begin with mussel (Bathymodiolus childressi) beds containing high biomass communities of low diversity and high endemicity. Individual mussels live for 100–150 years, whereas mussel beds may persist for even longer periods, with growth rates of mussels primarily controlled by methane concentrations (Nix et al., 1995).

    The next successional stage consists of vestimentiferan tubeworm aggregations dominated by Lamellibrachia luymesi and Seepiophila jonesi. Young tubeworm aggregations often overlap in time with, and usually persist past the stage of mussel beds. These tubeworm aggregations and their associated faunas go through a series of successional stages over a period of hundreds of years. Declines in seepage rates result from ongoing carbonate precipitation occluding pores and so forming aquitards, as well as the influence of L. luymesi on the local biogeochemistry as it extracts ever-larger volumes of H2S. In older tubeworm aggregations, biomass, density, and number of species per square metre decline in response to reduced sulphide concentrations.

    Once seep habitat space becomes available, more of the non-endemic background species, such as amphipods, chitons, and limpets, can colonise the mussel and tubeworm aggregations. Due to the lowering concentrations of sulphide and methane, the free-living microbial primary productivity is reduced. The number of associated taxa is positively correlated with the size of the tubeworm-generated habitat, so diversity in this stage remains relatively high although the proportion of endemic species is smaller in the older aggregations. This final stage may last for centuries, as individual vestimentiferan tubeworms can live for over 400 years (Cordes et al., 2009).

    Even as seepage of hydrocarbons declines in a particular seep site, the authigenic carbonate layers of relict seeps can still provide a stable seafloor substrate for marine filter feeders, such as cold-water corals. The scleractinians Lophelia pertusa and Madrepora oculata, several gorgonian, anthipatharian, and bamboo coral species form extensive reef structures atop now inactive seeps on the upper slope of the Gulf of Mexico (Schroeder et al., 2005). The corals obtain their food supply form the water column and are not dependent on chemosynthetic microbes. The coral communities also harbour distinct associated assemblages, consisting mainly of the general background marine fauna, but also contain a few species exclusively associated with the corals and a few species that are common to both coral and seep habitats

    Although individual tubeworms and molluscs in chemosynthetic brine pool communities may live for more than 300-400 years, vagaries in the rate of brine and nutrient supply to the seafloor mean many mussel and tubeworm colonies are overwhelmed by a rising halocline and so die in a shorter space of time. Their partially decomposed remains can spread out as part of the organic-rich debris atop the halocline, along with bacterial, algal and faecal residues, where it is acted upon by a rich community of aerobic and anaerobic decomposers. If the organic matter is mineralised or attaches to other interface precipitates such as pyrite, it sinks to the anoxic brine pool bottom, where it is largely preserved and protected from further biodegradation.

    The inherently unstable nature of the seafloor in the vicinity of active salt allochthons and brine lakes means it is subject to slumping, especially in the vicinity of brine fed mud volcanoes. In such settings, parts of the carbonate-rich biostrome rim are periodically killed “en masse” as sediment about a brine pool edge collapses, slumps and slides into anoxic pool waters, carrying with it the chemosynthetic community. As well as further elevating levels of preserved organics in the brine pool bottom sediments, this process also creates potential fossil lagerstaette. Death of seep communities, even if survives such catastrophic events, ultimately comes when the supply of seep gases and liquid hydrocarbons is cut off to any single seep.


    Hardgrounds, seafloor stability & stable isotopes

    Associated with the brine-pool communities, and helping form an initial stable seafloor substrate for the colonising seep invertebrates, are calcite-cemented biogenic crusts. These cemented hardgrounds precipitate as a microbial byproduct wherever methane and H2S are bubbling up in and around brine pool edges, and gases are being metabolised by chemosynthetic archaea and bacteria (Canet et al., 2006; Fu Chen et al., 2007; Feng et al., 2009). The resulting biogenic calcite crusts have δ13CPDB values ranging to as low as -53‰, which is characteristic of methanogenic carbon (Figure 4a). Seep sediments retain a group of unsaturated 2,6,10,15,19-pentamethylicosane (PMID) compounds, also produced by methane-oxidising archaea, with δ13CPDB values ranging from -107.2 to -115.5‰. In combination, the isotope values, textures and biomarkers indicate a combination of bacterially catalysed methane oxidation and sulphate reduction plexi in the crusts.

    Fabrics of the two flat sides of methanogenic calcite crusts crust are texturally distinct. The “top” side is composed entirely of microcrystalline calcite, while the bottom is composed entirely of “wormy” carbonate cement that is interpreted as a random, low fidelity replacement of bacteria. (Figure 4b) “Wormy” carbonate cement coats microcrystalline calcite in the interior of the thick crust and dispersed pyrite framboids appear to be indicators of collaborating colonies of methane-oxidising archaea and sulphate-reducing bacteria. Fu Chen et al., (2007) propose that the “wormy” carbonate texture, particularly with microcrystalline calcite and pyrite framboids present, is a likely indicator of biologically controlled fabrics produced during methane oxidation and sulphate reduction.


    Hypersaline brines and entrained gases escaping and pooling on the Gulf of Mexico seafloor do so either into quiescent brine lakes and pools or as mud chimneys and volcanoes (Figure 5; Joye et al., 2009). Both environments are anoxic and hypersaline, brine pools are typified by low fluid-flow rates and waters free of suspended sediment, while flow rates in mud volcano chimneys are more vigorous and the waters tend to be more turbulent and carry more suspended load. The sharp salinity transition between hypersaline brine and seawater typifies the water column in both settings, and a higher suspended particle load underscores the more rapid fluid-flow regime of the mud volcano (Figure 5a, f). Brines in both are mildly sulphidic; concentrations of dissolved inorganic carbon are elevated relative to seawater. Microbial abundance is 100 times higher in brines than in the overlying seawater (Figure 5a, f), showing that brine-derived substrates produce high microbial biomass. The brines are gas charged; the dominant dissolved alkane is methane (94-99.9%) with a stable carbon isotopic composition, 13C, of -62‰.

    The feeder brines to the chemosynthetic communities in much of the Gulf of Mexico form via halite dissolution and so contain little to no sulphate. Seawater sulphate diffuses into the brine, and concentrations decrease with depth, reflecting a combination of microbial consumption through sulphate reduction (both sites) and upward advection of sulphate-free brine in a mud volcano (Figure5b, g). The hydrogen profile in the mud volcano brine is relatively uniform (hundreds of nanomolar), reflecting the potential importance of autotrophic acetogenesis and/or hydrogenotrophic methanogenesis. In the brine pool, however, hydrogen concentration increases to micromolar levels between depths ≈25 and 100 cm and remains high (≈µ6 M) to 180 cm, promoting acetogenesis. Such high hydrogen concentrations indicate active fermentation and substantial inputs of labile organic matter. Concentrations of dissolved organic carbon (DOC) increases with depth (Figure 5b, g), suggesting a deep-subsurface DOC source (thermogenic?). In the brine pool, extra labile DOC, probably coming from the surrounding chemosynthetic community can further stimulate fermentation (Joye et al., 2009)

    Rates of acetate production and levels of sulphate reduction are much higher in brine pools, whereas the mud volcano supports much higher rates of methane production (Figure 5d, i). Joye et al. (2009) found no evidence of anaerobic oxidation of methane (AOM), despite high methane fluxes in both settings. It suggests both these systems are leaking methane into the overlying water column. Joye et al. conclude that the different halo-adapted microbial community compositions and metabolisms are linked to differences in dissolved-organic-matter input from the deep subsurface and different fluid advection rates between the two settings.

    Clathrates and methane seeps in the Gulf of Mexico

    Across the slope and rise in the Gulf of Mexico, where sea bottom temperatures are suitably low, methane hydrates (clathrates) form atop focused outflow zones and oil seeps are common at the sea surface above vent clathrates (Dalthorp and Naehr, 2011). Gas hydrate or clathrate is an ice-like crystalline mineral in which hydrocarbon and non-hydrocarbon gases are frozen within rigid molecular cages of water. They can be thought of gaseous permafrost. Their occurrence is not just tied to the cold temperature portion of the deep seafloor; clathrates are the dominant seals to large gas reservoirs in the permafrost regions of Siberia. Methane hydrates are common associations where methane, which can be thermogenically or biogenically sourced, occurs just below the deep cold seafloor. In much of world, it accumulates in seafloor regions independent of any underlying evaporite occurrence (Thakur and Rajput, 2011). Evaporite edges just tend to focus the outflow zones (Figure 6).


    Clathrate formation on the seafloor requires bottom temperatures not encountered until the seafloor bottom lies beneath a water column 450-500 m deep. Beneath the clathrate-covered seafloor, temperature increases with depth and this limits the depth at which gas hydrates will occur, so below most clathrate layer is an accumulation of free gas is likely. Clathrates seeps in the vicinity off brine pools are not unique to, but are often very obvious about, salt allochthon edges where salt flow induces extensional faulting and funnels a focused rise of methane, degraded oil and H2S to the cold seafloor (Chapter 6). Hence, breaks in the lateral extent of the various salt sheets act as a focusing mechanism for escaping thermogenic and biogenic methane and other gases and fluids (Figures 3, 6; Fisher et al., 2000; MacDonald et al., 2003). Rapid burial of organic-entraining sediments in supra-allochthon minibasins encourages the creation of biogenic methane that sources much of the gas escaping to the seafloor away from salt-edge focused seeps. Hence, in the salt allochthon province of the northern Gulf of Mexico, there is a definite association between brine pool chemosynthetic communities, thicker gas hydrates and the edges of minibasins (Figure 6; Reilly et al., 1996; Milkov and Sassen, 2001).


    In all these setting clathrates are a food source for various methanogenic microbes, and so there are different multi-cellular lifeforms dependent on these microbes. One obvious dependency is seen in the eco-niche occupied by a small 2-4 cm-long highly specialised polychaete called Hesiocaeca methanicola (Figure 7). It was discovered in 1997 flourishing in regions of methane hydrate atop the deep seafloor in the Gulf of Mexico (Fisher et al., 2000). These “ice worms” inhabit indentations (“burrows”) in blocks and layers of methane clathrate and glean or harvest biofilms of the methanotrophic bacteria that are metabolising methane on the block surface. In turn, the ice worm supplies oxygen to the methanotrophs and via its movement appears to contribute to the dissolution of hydrates. Mature ice worms can survive in an anoxic environment for up to 96 hours. The experiments oof Fisher et al., (2000) also showed that the larvae were dispersed by currents, and died after 20 days if they did not find a place to feed.

    Brine lake biota in the Mediterranean Ridges

    Eight brine lakes, L’Atalante, Bannock, Discovery, Kryos, Medee, Thetis, Tyro and Urania, have been discovered and studied in the Mediterranean Ridge region of the deep eastern Mediterranean over the last 20 years (Figure 8a; see part 1). The surfaces of these brine lakes lie between 3.0 and 3.5 km below sea level, and the salinity of their brines ranges from five to 15 times higher than that of seawater. In the Bannock Basin, the various brine-filled depressions or sub-basins create a closed outer moat around a central seafloor mound that is 10 km across (Figure 8b). The chemical composition of the Tyro Basin bottom brine is related to the dissolution of the underlying halite-dominated evaporites, while the chemical composition of the Bannock Basin (Libeccio Basin in the Bannock area) implies derivation from dissolving bittern salts (de Lange et al., 1990). In the “anoxic lakes region”, sodium chloride is predominantly sourced in the L’Atalante and Urania lakes, but L’Atalante is much richer in potassium chloride than the other lakes. The Discovery basin brine is almost exclusively the product of dissolution of magnesium chloride (bischofite) salts. It has a density of 1330 kg/m3, which makes it the densest naturally occurring brine yet discovered in the marine environment (Wallmann et al., 2002). Its concentration profile in sediment beneath the brine lake shows the age of this lake is between 700 and 2000 yr. The high concentration of magnesium chloride drives the dissolution of biogenic calcium carbonate but facilitates excellent preservation of siliceous microfossils and organic matter. In basin bottom muds there are large euhedral crystals of gypsum, up to 10cm across, precipitating from these brines (Cita 2006).

    Of the Mediterranean brine lakes, Lake Medee is the largest, and fills a narrow depression at the Eastern edge of the abrupt cliffs of the small evaporite ridge located 70 nautical miles SW of Crete (Figure 8a). The lake depression is approximately 50 km in length with a surface area of about 110 km2 and a volume of nearly 9 km3, which places Lake Medee among the largest of the known DHALs in the deep-sea environment. Although all the Mediterranean DHALs lie geographically close to each other, their hydrochemical diversity suggests that dissolving salt mineralogies were different. Salinity levels are much higher in some dues to the presence off nearby bittern layers. For example, Discovery Lake and Lake Kryos have salinities and MgCl2 proportions indicative of bischofite dissolution. Even so, it seems like, mostly sulphate-reducers can still metabolise in the extremely saline MgCl2 waters of Lake Kryos (Steinle et al., 2018).

    In contrast to the brine lakes and seeps in salt-allochthon terrane of the Gulf of Mexico, seep megafauna is so far absent in the various documented modern brine lakes along the Mediterranean Ridges (Figure 8d). The brine lakeshore edge communities are mostly microbial, as are the lifeforms that make up the pelagic biota off the halocline. Biological studies on the anoxic basins of the Eastern Mediterranean started after the discovery of gelatinous matter of organic origin in the brine lake sediments (Figure 8c; Brusa et al., 1997). The laminar gelatinous matter was observed within the cores containing anoxic sediments obtained during oceanographic expeditions for geological study of the Mediterranean Ridge. Microbiological and ultrastructural investigations were carried out on core sediment samples and on the overlying water. Various authors demonstrated the organic nature of the mucilaginous pellicles found in the cores and their relation with numerous microbic forms present in all the samples. Viable microorganisms, prevalently Gram-negative and aerobic as well as facultative anaerobes, were found in the halocline water samples. Different microbic forms were isolated in pure culture: a vibrio (Nitrosovibrio spp.), a coccus (Staphylococcus sp.) and some rods of the family Pseudomonadaceae. In addition, laminar formations were observed in a growth medium of mixed cultures that could be interpreted as the first stages of the mucilaginous pellicles seen in the cores. Earlier studies described the geological and physiochemical characteristics of such habitats (Erba et al. 1987; Cita et al. 1985). Subsequent work using metagenomic techniques have documented a prosperous microbial community inhabiting the halocline of most of the Mediterranean brine lakes.

    DHAL interfaces in the Mediterranean Sea deeps act as hot spots of deep-sea microbial activity that significantly contribute to de novo organic matter production. Metabolically active prokaryotes are sharply stratified across the halocline interfaces in the various brine lakes and likely provide organic carbon and energy that sustain the microbial communities of the underlying salt-saturated brines. Since metagenomic analysis of DHALs is still in its infancy, the metabolic patterns prevailing in the organisms residing in the interior of DHALs remains mostly unknown. What is known is that the redox boundary at the brine/seawater interface provides energy to various types of chemolithic and heterotrophic communities. Aerobic oxidations of reduced manganese and iron, sulphide and intermediate sulphur species, diffusing from anaerobic brine lake interior to the oxygenated upper layers of the haloclines are highly exergonic processes capable of supporting an elevated biomass at DHAL interfaces (Yakimov et al., 2013). Depending on availability of oxygen and other electron acceptors bacterial autotrophic communities belonging to Alpha-, Gamma- and Epsilon-proteobacteria fix CO2 mainly via the Calvin-Benson-Bassham and the reductive tricarboxylic acid (rTCA) cycles, respectively.

    Biomarker associations of the organics accumulating in the brine lakes define two depositional styles: typical marine and hypersaline (Burkova et al., 2000). For example, algal and bacterial biomarkers typical of saline environments were found in layers 0.60 to 0.75 m below the sediment surface in the Tyro Lake Basin, as well as standard marine indicators derived from pelagic fallout (“rain from heaven”). Saline indicators include; regular C-25 isoprenoids, squalane, lycopane, isolycopane, tetraterpenoid and tetrapyrrolic pigments, monoalkylcyclohexanes, tricyclic diterpanes, steranes, hopanes, bio- and geohopanes. According to Burkova et al. (2000), the saline organic signatures come from microbial mat layers, redeposited from a Messinian source into the sapropels of the modern depression. Alternatively, they may indicate the activities of a chemoautotrophic community, which flourishes at the halocline or around active brine vents. As in the Orca Basin, the organic content of the bottom sediments of the Mediterranean brine pools is much higher than is found in typical deep seafloor sediment (Figure 9a).


    Anoxic hypersaline brines in Mediterranean brine lakes are highly sulphidic and among the most sulphidic bodies of water in the marine realm, with H2S concentrations consistently greater than 2-3 mmol (Table 1; Henneke et al., 1997). The brine body below the Urania chemocline is more than 100 m thick and contains up to 11 mM hydrogen sulphide, making it the most sulphidic water body in the known marine realm. In combination with the sulphide are very high levels of methane both in and below the halocline (≈5.56 mM; Borin et al., 2009). In contrast, there is little to no H2S in the anoxic bottom brine of the Orca Basin (Table 1). There the iron concentration is 2 ppm, a value more than 1000 times higher than in the overlying Gulf of Mexico seawater. Such high levels of reducible iron in the Orca Basin are thought to explain the lack of H2S in the bottom brine and a preponderance of framboidal pyrite and extractable iron in the bottom sediments (Sheu, 1987). Both the Orca Basin and the brine pools on the floor of the Mediterranean, show sulphate levels that can be more than twice that of the overlying seawater.


    So, organic debris first formed at the halocline can then accumulated as pellicle layers within the pyritic bottom muds (laminites). Pellicular debris is also carried to the bottom during the emplacement of turbidites when the halocline is disturbed by turbid overflow (Figure 10; Erba, 1991). Hence, pellicular layers are typically aligned parallel to lamination, or are folded parallel to the sandy bases of the turbidite flows, or line up parallel to deformed layers within slumped sediment layers. Individual pellicle layers are 0.5 to 3 mm thick and dark greenish-grey in colour. Similar pellicular layers cover the surface of, or are locked within, recent gypsum crystals recovered from bottom sediments of the Bannock area. This gypsum is growing today on the bottom of the Bannock Basin, atop regions about the brine pool margin that are directly underlain by dissolving Miocene evaporites (Corselli and Aghib, 1987; Cita 2006). Other than the Dead Sea, it is one of the few modern examples of a deepwater evaporite, but its seepage-fed genesis means it is a poor analogue for deepwater basinwide salt units.

    The community of bacteria and archaea flourishing at the halocline in sulphidic marine brine pools on the deep Mediterranean floor is quite diverse, mostly independent of primary production in the euphotic zone, with the number of identified unique halobacteria and haloarchea species expanding every year (Albuquerque et al., 2012). Bottom brine in the Urania brine lake has a salinity of 162‰, and the chemocline of the brine lake is some 3490m below the ocean surface, so only a minimal amount of phytoplanktonic organic carbon ever reaches the 20m thick chemocline. Yet the oxic waters of the upper part of the chemocline support a rich bacterial and archaeal assemblage in and below the interface between the hypersaline brine and the overlying seawater, much like the chemosynthetic bacterial community associated with the halocline in Lake Mahoney (Sass et al., 2001; Borin et al., 2009).


    Sulphide concentration in the Urania Basin increases from 0 to 10 mM within a vertical interval of 5 m across the interface (Figure 11a). Within the halocline, the total bacterial cell counts and the exoenzyme activities are elevated and biogenic activity continues below the halocline. Bacterial sulphate reduction rates measured in this layer are ≈ 14 nmol SO4 cm-3 d-1 and are among the highest in the marine realm. They correspond to the zone of maximum bacterial activity in the chemocline (Figure 11b). Particulate organic content is 15 times greater than that in the overlying normal marine waters. A similar focus of microbial occurrence (bacterial and archaeal) is seen at the halocline in l’Atalante Basin and is probably typical of all chemocline layers in the various Bannock brine lakes (Yakimov et al., 2007)

    Employing 11 cultivation methods, Sass et al. 2001 isolated a total of 70 bacterial strains from the chemocline in the Urania Basin (Figure 11a). These strains were identified as the flavobacteria, Alteromonas macleodii, and Halomonas aquamarina. All 70 strains could grow chemo-organoheterotrophically under oxic conditions. Twenty-one of the isolates could grow both chemo-organotrophically and chemo-lithotrophically (decomposers and fermenters). While the most probable numbers in most cases ranged between 0.006 and 4.3% of the total cell counts, an unusually high value of 54% was determined above the chemocline with media containing amino acids as the carbon and energy source.

    Subsequent detailed work focused on the various layers that make up the Urania halocline showed the high sulphide levels in and below the halocline, make it a mecca for bacterial sulphate reducers, as do high levels of methane for the methanogens (Figure 11b; Borin et al., 2009). Microbial abundance showed a rapid increase by two orders of magnitude from 3.9 x 104 cells mL-1 in the deep oxic seawater immediately above the basin, up to 4.3 x 106 cells mL-1 in the first half of interface 1. Although less pronounced than in the first chemocline, a second increase in microbial counts occurred in interface 2. Deceleration of falling particulate organic matter from the highly productive interface 1, is probably responsible for stimulating microbial growth and hence cell numbers in interface 2. That is, compared to the overlying seawater column, bacterial cell numbers increased up to a hundred-fold in interface 1 and up to ten-fold in interface 2. This is a consequence of elevated nutrient availability, with higher numbers in the upper interface where the redox gradient was steeper. Bacterial and archaeal communities, analysed by DNA fingerprinting, 16S rRNA gene libraries, activity measurements, and cultivation, were highly stratified within the various layers of the chemocline and metabolically more active along the various chemocline layers, compared with normal seawater above, or the uniformly hypersaline brines below.

    Detailed metagenome analysis of 16S rRNA gene sequences revealed that in both chemocline interfaces the e- and d-Proteobacteria were abundant, predominantly as sulphate reducers and sulphur oxidisers, respectively (Figure 11b). The only archaea in the first 50 cm of interface 1 were Crenarchaeota, which consist of organisms having sulphur-based metabolism, and hence could play a role in sulphur cycling in the upper interface. In the deepest layers of the basin below the halocline, MSBL1, putatively responsible for methanogenesis, dominated among archaea (Figure 11b). The work of Borin et al. (2009) illustrate that a well adapted and complex microbial community is thriving in the Urania basin’s extreme chemistry, The elevated biomass centred on the halocline is driven mainly by sulphur cycling and methanogenesis.

    Similarly detailed studies of interface-controlled chemosynthetic communities in other Mediterranean DHALs have been documented in Lake Thetis (Ferrer et al., 2012; Oliveri et al., 2013) and Lake Medee (Yakimov et al., 2013). Medee Lake is the largest known DHAL on the Mediterranean seafloor and has two unique features: a complex geobiochemical stratification and an absence of chemolithoautotrophic Epsilonproteobacteria, which usually play the primary role in dark bicarbonate assimilation in DHALs interfaces worldwide. Presumably, because of these features, Medee is less productive and exhibits a reduced diversity of autochthonous prokaryotes in its interior brine layers. Indeed, the brine community almost exclusively consists of the members of euryarchaeal and bacterial KB1 candidate divisions which a ubiquitous in the DHAL biota worldwide. In Medee, as elsewhere, they are thriving on small organic molecules produced by a combination of degraded marine plankton and moderate halophiles living in the overlying stratified brine column.

    Outside off the microbial makeup of DHAL communities, one of the more exciting discoveries in the brine lakes of the Mediterranean ridges is the likely discovery of multicellular life of the Phylum Loricifera (“Beard shells) capable of living and reproducing in the absence of oxygen. Loricifera (from Latin, lorica, corselet (armour) + ferre, to bear) is a phylum made up of very small to microscopic marine cycloneuralian sediment-dwelling animals with 37 described species. Their size ranges from 100 µm to ca. 1 mm and individuals are characterised by a protective outer case called a lorica and by their habitat, which is in the spaces between marine sediment particles. The phylum was first discovered in tidal sediments in 1983 and is among the most recently discovered groups of Metazoans. Individuals attach themselves quite firmly to the sediment substrate, and hence the phylum remained undiscovered for so long. In 2010, viable specimens of Spinoloricus cinziae, along with two other newly discovered species, Rugiloricus nov. sp. and Pliciloricus nov. sp., were found in the sediment core from below the anoxic L'Atalante basin of the Mediterranean Sea (Danovaro et al., 2010, 2016). The species cellular innards appear to be adapted for a zero-oxygen life as their mitochondria appear to act as hydrogenosomes, organelles which already provide energy in some anaerobic single-celled creatures known. Before their discovery, living and reproducing exclusively in an oxygen-free setting was thought to be a lifestyle open only to viruses and single-celled microorganisms. The ability of these anoxic brine-dwelling creatures to live solely in an oxygen-free environment is questioned still by other workers (Bernhard et al., 2015).

    Neither Tyro nor Bannock Basin bottom sediments show a significant correlation between pyritic sulphur and the organic carbon in the bottom sediments, suggesting predominantly syngenetic pyrite evolution in bottom sediments of these brine lakes (Henneke et al., 1997). That is, both pyritic and humic sulphur preserved in the bottom sediments formed either in the lower water column or at the sediment-brine interface, not in the sediment itself. Ongoing diagenetic processes within the bottom sediments only form an additional 5% of the total pyrite. Van der Sloot et al. (1990) clearly showed that metal sulphides, as well as organics and other minerals, precipitate at the brine-seawater interface in the Tyro Basin, as they do in the Orca Basin. They found extremely high concentrations of Co (0.015%), Cu (1.35%) and Zn (0.28%) in suspended matter at the halocline. These high particulate Co, Cu and Zn concentrations correspond to sharp increases in dissolved sulphide across the interface (a redox front), and indicate precipitation of metal sulphides at the interface. Humic sulphur in the bottom sediments correlates with the pyritic sulphur distribution and is related to the amount of gelatinous pellicle derived from bacterial mats growing at the halocline between oxic seawater and bottom brine (Erba, 1991, Henneke et al., 1997).

    Additionally, the degree of pyritisation in the sediments (DOP ≈ 0.62) indicates that present-day pyrite formation is limited by the reactivity of Fe in the Bannock and Tyro basins and not by the availability of organic matter, the latter being the process that limits pyrite formation in most normal marine settings (Figure 9b). The degree of pyritisation (DOP) is defined as [(pyritic iron)/(pyritic iron + reactive iron)]. Raiswell et al. (1988) showed that DOP in ancient sediments can distinguish anoxic from normal marine sediments. Anoxic sediments show DOP values between 0.55 and 0.93, while normal marine sediments have DOP values less than 0.42. The DOP levels in the Bannock and Tyro basins confirm observations made in ancient anoxic sediments. Thus, although the Tyro and Bannock basin brines differ in their major element chemistry, reflecting a different salt source, their reduced sulphur species chemistry appears to be similar, but is significantly different from standard marine systems and capable of precipitating metal sulphides above the sediment surface.


    Life in the Red Sea brine deeps

    The Atlantis II Deep marks the northern-most end of the Atlantis II Shagara- Erba Trough section, hosting numerous sub-deeps like the Discovery and Aswad Deep (Figure 12). In general, the Atlantis II Deep area has a smoother bathymetric character than the Thetis-Hadarba-Hatiba and Shagara-Aswad-Erba Troughs, due to massive inflow of salt and sediments from nearly all sides into the deep. In the Atlantis II deep, Siam et al. (2012) identified metagenomic archaeal groups in high relative abundance at the bottom of a sediment core from the Atlantis II Deep, which, as in the Kebrit Deep, are another case of the dominance of Archaea. Their results showed that the dominant archaeal inhabitants in the bottom layer (3.5 m depth to the seafloor) included Marine Benthic Group E, and the archaeal ANME-1 ( anaerobic methane consumers metagenome. The presence of the latter was also confirmed in a study of a barite mound in the Atlantis II Deep (Wang et al., 2015), but the former was not detected in this later study.

    In metagenomic studies of the Atlantis II sediments, Cupriavidus (Betaproteobacteria) and Acinetobacter (Gammaproteobacteria) are the most abundant species in the surface layer (12 cm) and the bottom layer (222 cm) of a sediment core obtained in 2008. Both bacterial species were not the dominant inhabitants in the ABS core analysed in the present study. Due to tremendous differences between brine water and sediment chemistry in the Deep, their microbial communities differ remarkably. The lower convective layers of the Atlantis II and Discovery brine pools are dominated by Gammaproteobacteria, while Alphaproteobacteria and Betaproteobacteria are the major bacterial groups in the upper layers of Atlantis II sediment (Bougouffa et al., 2013). All the above discrepancies in composition of microbial communities in the two Deeps were probably caused by 1) primer selection for amplification of rRNA genes; 2) different microenvironments in the sampling sites; 3) taxonomic assignment criteria employed by different studies; 4) different experimental procedures, and 5) sampling bias due to low biomass in sampling sites. Except for these potential problems, this study demonstrates the profound changes in microbial communities in deep-sea hydrothermal sediment under the influence of extensive mineralisation process. Many of the groups detected in the S-rich Atlantis II section are likely to play a dominant role in the cycling of methane and sulphur due to their phylogenetic affiliations with bacteria and archaea involved in anaerobic methane oxidation and sulphate reduction.


    In the Kebrit Deep on the deep floor of the Red Sea, an assemblage of halophilic archaea and bacteria similar to that of the DHALs of the Mediterranean Deeps flourish in hypersaline waters below the chemocline (Figure 13). Kebrit Deep (24°44’N, 36°17’E) measures 1 by 2.5 km, with a maximum depth of 1549 m and is one of the smallest salt allochthon-associated brine-pools of the Red Sea. It is located around 300 km nothwest the well-known metalliferous Atlantis II deep (see previous article). The Kebrit Deep is filled by an 84 m thick, anaerobic, slightly acidic brine lake (pH approximately 5.5) with a salinity of 260‰ and a temperature of 23.3°C (Antunes et al., 2011). The brine has a high gas content that is made up mainly of CO2, H2S, small amounts of N2, methane and ethane, with remarkably high quantities of H2S (12–14 mg S l-1; Hartmann et al., 1998). The presence of sulphur is self-evident by the strong, characteristic odour present in brine samples, and hence the name of the basin (Kebrit is the Arabic word for sulphur). Like the Atlantis II deep there are impregnated massive sulphides accumulations on the floor of Kebrit Deep. Kebrit samples are porous and fragile, and consist mainly of pyrite and sphalerite. Prior to gene sequencing studies, sulphur isotope values provided substantial evidence for biogenic sulphate reduction being involved in sulphide-forming processes in Kebrit Deep. They are linked to bacterial methane oxidation and sulphate reduction centred on the brine-seawater interface (see Chapter 15 in Warren 2016 for metallogenic details).

    Most of the archaeal metagenomic sequences in Kebrit Deep cluster within the Thermoplasmatales (Marine group II, Marine Benthic group D, and the KTK-4A cluster) among the Euryarchaeota, while the remaining sequences do not show high similarity to any of the known phylogenetic groups (Figure 13). One of these sequences was shown to cluster with the later-described SA2 group, while another (accession number AJ133624) clusters together with two gene sequences from L’Atalante Basin waters, defining a novel deeply-branching phylogenetic lineage within the Crenarchaeota.

    Gene sequencing studies on water samples from the brine-seawater interface in the Kebrit deep retrieved sequences from the KB1 group, as well as Clostridiales (mostly Halanaerobium), Spirochetes (ST12-K34/MSBL2 cluster), Epsilonproteobacteria and Actinobacteria, but no archaeal sequences were detected in these interface samples (Antunes et al.,2011). Under strictly anaerobic culture conditions, novel halophiles were isolated from samples of these waters and belong to the halophilic genus Halanaerobium. They are the first representatives of the genus obtained from deep-sea, anaerobic brine pools (Eder et al., 2001). Within the genus Halanaerobium, they represent new species that grow chemo-organotrophically at NaCl concentrations ranging from 5 to 34%. They contribute significantly to the anaerobic degradation of organic matter, which formed at the brine-seawater interface and is slowly settling into the bottom brine.

    Similarities in the makeup of the Archaeal population, tied to similar metabolic process sets at the brine interface across various deep seafloor brine lakes in the Gulf of Mexico, the Mediterranean and the Red Sea. Compared with other hydrothermal sediments around the world, the Atlantis II hydrothermal field is unique in that sulphur and nitrogen oxides are low in the pore water of the sediments. This probably leads to lack of ANME . It seems, different geochemical conditions of hydrothermal marine and cool seep sediments across the deepsea sub-seafloor resulted in various niche-specific microbial communities.

    Life in the Dead Sea

    As defined in the salty matters article previous to this, the Dead Sea can be considered a continental counterpart of a marine DHAL where there is no overlying body of marine water. Instead, the Dead Sea brine mass is in direct contact with the atmosphere.

    The Dead Sea provides one of nature’s supreme tests of survival of life. The negative-water balance in the Dead Sea hydrology over recent decades resulted in ever-rising salinity and divalent-cation ratios, cumulating in the current highly drawdown situation (See Warren 2016, Chapter 4 for a summary of the relevant hydrological evolution. Today the brines have reached a salinity level more than 348 /l total dissolved salts, with a high ratio of (Ca + Mg) to Na. Water activity (Aw, a measure based on the partial pressure of water vapour in a substance, and correlated with the ability to support microorganisms) of the Dead Sea is extremely low (Aw ≈ 0.669), even lower than that of saturated-NaCl solution (Aw ≈ 0.753±0.004), and is thus unbearable for most life forms (Kis-Papo et al., 2014).

    Nevertheless, a number of halobacteria (Archaea), one green algal species (Dunaliella parva), and several fungal taxa withstand these extreme conditions(Kis-Papo et al., 2014). Most organisms in the Dead Sea survive in fresher-water spring refugia or in their dormant stages or and only revive when salinity is temporarily reduced during rare massive flooding events (Ionescu et al., 2012.

    Effects of occasional freshening on biomass in stratified brine columns that are supersaline, not mesohaline, is clearly seen in the present “feast or famine” productivity cycle of the Dead Sea (Warren, 2011; Oren and Gurevich, 1995; Oren et al., 1995; Oren 2005). Dunaliella sp, a unicellular green alga variously described in the past as Dunaliella parva or Dunaliella viridis, is the sole primary producer in the Dead Sea waters. Then there are several types of halophilic archaea of the family Halobacteriaceae (prokaryotes) which consume organic compounds produced by the algae.


    Two distinct periods of organic productivity (feast) have been documented in the upper lake water mass since the Dead Sea became holomictic in 1979 (Oren, 1993, 1999). The first mass developments of Dunaliella sp. (up to 8,800 cells/ml) began in the summer of 1980 following dilution of the saline upper water layers by the heavy winter rains of 1979-1980 Figure 14a, b). The rains drove a rapid rise of 1.5 metres in lake level and an increase in the level of phosphates in the lake’s surface waters (Figure 14c). This bloom was quickly followed by a blossoming in the numbers of red halophilic archaea (2 x 107 cells/ml), Dunaliella numbers then declined rapidly following the complete remixing of the water column and the associated increase in salinity of the upper water mass. By the end of 1982, Dunaliella had disappeared from the main surface water mass. Archaeal numbers underwent a slower decline.

    During the period 1983-1991 the lake was holomictic, halite-saturated and no Dunaliella blooms were observed. Viable halophilic and halotolerant archaea were probably present in refugia about the lake edge during this period but in meagre numbers. Then heavy rains and floods of the winter of 1991-1992 raised the lake level by 2 metres and drove a new episode of meromictic stratification as the upper five metres of the water column was diluted to 70% of its normal surface salinity (Figure 14d). High densities of Dunaliella reappeared in this upper less saline water layer (up to 3 x 104 cells/ml) at the beginning of May 1992, rapidly declining to less than 40 cells/ml at the end of July 1992 (Figure 15). An associated bloom of heterotrophic haloarchaea (3 x 107 cells/ml) continued past July and continued to impart a reddish colour to the surface and nearsurface waters.

    Much of the archaeal community was still present at the end of 1993, but the amount of carotenoid pigment per cell had decreased two- to three-fold between June 1992 and August 1993 (Oren and Gurevich, 1995). A remnant of the 1992 Dunaliella bloom maintained itself at the lower end of the pycnocline at depths between 7 and 13 m (September 1992- August 1993), perhaps chasing nutrients rather than light. Its photosynthetic activity was low, and very little stimulation of archaeal growth and activity was associated with this algal community (Figure 15). It seems that once stratification ends and the new holomictic period begins, the remaining Archaeal community, which was primarily restricted to the upper water layers above the halocline, spreads out more evenly over the entire upper water column until it too dies out. No substantial algal and archaeal blooms have developed in the Dead Sea since the winter floods of 1992-1993 until today


    Underwater freshwater to brackish springs are likely refugia to much of the life in the Dead Sea and are inhabited by interesting microbial communities including chemolithotrophs, phototrophs, sulphate reducers, nitrifiers, iron oxidisers, iron reducers, and others. The springs also host numerous cyanobacterial and diatomatous mats with sulfate-reducers near the base of the foood chain (Oren et al., 2008; Ionescu et al., 2012). Sequences matching the 16S rRNA gene of known sulphate-reducing bacteria (SRB) and sulphur oxidising bacteria (SOB) were detexcted in all microbial mats centered on freshwater springs as well as in the Dead Sea water column (Häusler et al., 2014). Generally, sequence abundance of SRB and SOB was higher in the microbial mats than in the Dead Sea, indicating that the conditions for both groups are more favorable in the spring environments.

    The springs also supply nitrogen, phosphorus and organic matter to the Dead Sea microbial communities. Due to frequent fluctuations in the freshwater flow volumes in the springs and local salinity, microorganisms that inhabit these springs must be capable of withstanding large and rapid salinity fluctuations and the population proportions vary according to the Spring chemistry (Ionescu et al., 2012).

    Salt dissolution, seafloor salinity and halophilic extremophile populations

    In most DHALs, the rate of vertical mixing across the extreme density gradients between brine and overlying seawater is extremely slow (Steinle et al., 2018). Hydrochemically, depending on the nature of the dissolving salt supply, seawater and DHAL brines can differ sharply in their solute composition, in particular, in the concentrations of the critical electron donors and acceptors so crucial to the functioning of life. In that a narrow (1– 3 m) chemocline (halocline) forms a transition zone between the two quite-different hydrologies that define a DHAL water column, microbial ecologies have evolved to inhabit particular portions of the halocline as well as the brine lake and the normal marine deepwater columns (Figure 16).

    In contrast to the overlying seawater, the bottom brines are anoxic but contain electron acceptors other than oxygen most importantly sulphide and methane. Hence, hotspots of chemosynthetic (not photosynthetic) activity have evolved that flourish at these brine-seawater interfaces, where the principal reactions at the base of the food chain are anoxic and encompass sulphate reduction, methanogenesis, and microbial heterotrophy. Highly-adapted microbial life continues to function even in the most extreme hypersaline conditions found in some DHALs, such as in Lake Kryos where MgCl2-rich chemistries dominate, or in the Atlantis II Deep where there is a combination of extreme temperatures and salinities.


    In the Gulf of Mexico, an endosymbiotic megafauna constructs methanogenically-cemented carbonate biostromes as lake fringe mussel-dominated communities or polychaete forests atop cool water H2S seeps. Both the microbial population and the megafauna that exploits this chemosynthetic base to the food chain flourish best in seafloor regions defined by the long-term focused escape of methane or H2S (Figure 16). Cool-seep brine lakes were first discovered in the Gulf of Mexico in the early 1980s, but similar hydrocarbon-dependent cool-seep communities with their own megafauna accumulations are now documented in other parts of the world characterised by the naturally-focused escape of hydrocarbons to the seafloor (for example, atop cool-water brine seeps along the slope and rise of the east and west coasts of North America and in the Black Sea.

    The relative long-term stability of cool-seep ecology, tied to the chemical stability of the niche, is seen when lifespans of hydrothermal endosymbiotic communities living chemosynthetically about thermal vents along mid-oceanic ridges are compared to Gulf of Mexico communities. Endosymbiotic polychaete and clam species in the brine lakes and seeps of the Gulf of Mexico can live for a hundred or more years, while lifespans in similar endosymbiotic polychaete and clam species in hydrothermal ridges communities are less than 30-50 years.

    Moving onshore, into the partial analogue offered by the salt-karst fed Dead Sea depression, we see Dead Sea biomass is subject to much shorter-term changes in the salinity and nutrient content of its uppermost water mass (Feast and Famine cycles as documented in Warren, 2011, 2016 Chapter 9). The freshening water mass above a lake halocline his ephemeral in the current longterm holomictic hydrology of the Dead Sea (see Warren 2016 chapter 4 for details). The changes in surface water salinity are tied to the periodic influx of a freshened upper water mass. These climatically-driven fluctuation to the the extent and activity of the halotolerant and halophilic community in the upper water mass, and the Feast or Famine responses of the Dead Sea biota, are different to the longterm niche stability created by the presence of a perennial oceanic water mass over a salt-karst induced halocline and brine lake in a DHAL sump on the deep seafloor. The latter is continually resupplied brine and chemosynthetic nutrients via the dissolution and focusing effect of the underlying salt sheet. The hydrology of a DHAL system only shuts down when all the mother salt is dissolved or cut off.

    Accordingly, rather than the hundreds of years of longterm growth (albeit at relatively slow metabolic rates) that we see in a DHAL, in the Dead Sea we see that freshening facilitates a rapid spread of a halotolerant alga (Dunaliella sp.) and associated halophilic microbes and viruses. The propagation and persistence of a large biomass pulse in the Dead Sea is measured in timeframes of months. The halotolerant photo-synthesisers can only spread out from long-term refugia communities once the surface salinities fall to levels that allow the photosynthesising base too the Lake food chain inhabit fresher water springs regions about the lake margins. Comparison to the DHAL and Dead Sea communities underlines how life will evolve into any neighbourhood, even if conditions are extremely challenging

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    Ziegenbalg, S. B., D. Birgel, L. Hoffmann-Sell, C. Pierre, J. M. Rouchy, and J. Peckmann, 2012, Anaerobic oxidation of methane in hypersaline Messinian environments revealed by 13C-depleted molecular fossils: Chemical Geology, v. 292-293, p. 140-148.

     

    Deepsea Hypersaline Anoxic Lakes & Basins (DHALs & DHABS)

    John Warren - Friday, August 31, 2018

     

    Introduction

    Since the 1980s, a new salt-accumulating subaqueous brine-lake style, tied to the dissolution of shallow sub-seafloor salt has been documented on the deep seafloor below a normal marine salinity water column. These are known as DHAL (deeps hypersaline anoxic lake) or DHAB (deeps hypersaline anoxic basin) deposits. They are described in salt allochthon regions on the deep seafloors of the Gulf of Mexico, the Mediterranean Sea and the Red Sea. All possess hydrologies and sediment columns characterised by prolonged separation of the bottom brine mass from the upper marine water column; a stratification that is due to a lack of mixing controlled by extreme conditions of elevated salinity, anoxia, and relatively high hydrostatic pressure and temperatures in the bottom waters.


    DHABs form in depressions where dense anoxic brines pond in stratified hypersaline lakes or basins on the seafloor, as vented hypersaline brines seep into closed seafloor depressions (Figure 1). The ponded bottom brines create distinctive brine interfaces with the overlying seawater, while the laminites deposited in the brine ponds are subject to occasional slump events. Both the interface and the bottom brine host well-adapted chemosynthetic communities and are described in detail in the next article in this series. DHABs typically form via local subsidence atop dissolving shallow allochthonous salt sheets or atop areas of salt withdrawal. Accordingly, DHABs tend to form adjacent to characteristic growth-faults or salt welds and to occur within rim syncline depressions; both features that are seismically resolvable in halokinetic terrains.

    This first article on DHALs focuses on the hydrology and physical geology/sedimentology of these interesting systems. The next will focus on the chemosynthetic communities that inhabit these brine lakes.


    Hydrology

    A DHAL or DHAB is a depression holding hypersaline water more saline than the overlying seawater (Table 1). Their deep-sea position, usually a few kilometres below the sea surface means DHALS are regions of a high-pressure bottom (> 35 MPa), total darkness, anoxicity and extreme salt-conditions (>250-350‰ salinity), some 5-10 times higher than normal seawater (≈35-40‰). Bottom brine chemistries typically have high concentrations of sulfides, manganese and ammonium, but at levels that vary independently across different basins (Table 1). The high density of the brine prevents it from mixing with overlying oxic seawater, so the water column is always density-stratified with permanently structured depth profiles typified by a chemocline or halocline interface (suboxic) separating the brine layer below (anoxic) and the normal marine (oxic) water column above.

    One of the interesting features of a DHABs is the perennial halocline; this is the zone where hypersaline waters meet the normal seawater above them. Because of an inherently high salt content, the bottom brine in a DHAB is so dense that it mixes very little with the overlying seawater. As you move down through the halocline, the salt concentration goes from normal seawater salinity to hypersaline. Along that gradient, the density of the water goes from that of normal seawater (≈1.04) to very high (1.1-1.2), and the oxygen concentration drops from normal seawater concentrations to zero. In some basins the halocline is only a meter thick, in others, it is more than a few metres thick.

    The temperature profile in a DHAL water column is distinct; it is always characterised by warmer bottom DHAL brine and cooler upper marine brine. Across some haloclines the temperature contrast is only a degree or two, in others, like some Red Sea deeps, the temperature contrast is tens of degrees.

    While a salt-karst-fed brine continues to supply the depression, a DHAL brine mass and its halocline show long-term stability. This long-term stability of the chemical interface facilitates laminite deposition, periodic bottom slumps and long-term chemical reactions at the brine interface, so facilitating the evolution of lifeforms well suited to a chemosynthetic habitat.

    By definition, a DHAB is a basin (closed seafloor depression), with walls that come up like the sides of a bowl. The halocline sits on top of the very salty water in the basin and touches the sides of the basin. Researchers sometimes call that area of intersection of the halocline with the basin floor area the “bathtub ring” because it is like the ring of soap scum and dirt that forms on a bathtub when the water is drained out. The sediment in this narrow "scum" zone has a little bit of oxygen and less salt than sediments inside the DHAB.

    Occurrences

    DHABS need a long-term brine source and so are found in halokinetic seafloor provinces where salt has flowed into a sufficiently shallow sub-seafloor position to be dissolving (salt karst). Often there is a faulted margin acting as a preferential brine conduit and seep zone supplying the nearby salt-withdrawal depression (Figure 1).


    Orca Basin, Gulf of Mexico

    The Orca Basin is a brine-filled minibasin atop a shallow salt allochthon at a depth of 2,400 metres, and some 600m below the surrounding seafloor. It is one of more than 70 such brine-soaked minibasins atop the allochthonous salt canopy in the northeast Gulf of Mexico (Figures 2, 3a).

    3D seismic images published by Pilcher and Blumstein (2007) show the Orca brine lake is surrounded by clay-rich slope sediments, which in the NE flank have slumped to “expose” shallow Louann salt to dissolution and seafloor karstification. They argue that dense anoxic brines in the Orca brine lake come mostly from this shallow salt (bright orange area in Figure 2a). The brine seeps downslope to pond in the sump of the basin as a 123 km2 lake of hypersaline brine, which is up to 220 m deep. Time-averaged addition of salt to the brine lake is calculated to be ≈0.5 million t/yr, and the resulting 13.3 km3 volume of the brine lake represents the dissolution of some 3.62 billion tons of Louann salt. The seismic shows that the depression hosting the closed brine lake area is a salt-withdrawal mini-basin.

    The Orca Lake sump encloses a 200m column of highly saline (259‰) anoxic brine, which is more than a degree warmer than the overlying seawater column (Figure 3b). The pool is stable and has undergone no discernable change since it was first discovered in the 1970s. It is a closed dissolution depression fed by brines seeping from a nearby subsurface salt allochthon (Addy and Behrens, 1980). A significant portion of the particulate matter settling into the basin is trapped at the salinity interface between the two water bodies. Trefry et al. (1984) noted that the particulate content was 20-60 µg/l above 2,100m and 200-400µg/l in the brine column below 2,250m. In the transition zone, the particulate content was up to 880 µg/l and contained up to 60% organic matter.


    A core from the bottom of the Orca brine pool captured laminated black pyritic mud from the seafloor to 485 cm depth and entrained three intralaminite turbidite beds of grey mud with a total thickness of 70cm (Figure 3c; Addy and Behrens, 1980). Grey mud underlies this from the 485 cm depth to the bottom of the core at 1079 cm. The laminated black mud was deposited in a highly anoxic saline environment, while grey mud deposition took place in a more oxic setting. The major black-grey boundary at 485 cm depth has been radiocarbon dated at 7900 ± 170 years and represents the time when escaping brine began to pond in the Orca Basin depression. Within the dark anoxic laminates of the Orca Basin, there are occasional mm- to cm-thick red layers where hematite and other iron hydroxides dominate the iron minerals and not pyrite. These reddish layers represent episodes of enhanced mixing across the normally stable oxic-anoxic halocline and indicate the short-term destruction of bottom brine stratification. When the plot of leachable iron is plotted, it is obvious that the pore brines in the black mud intervals can store iron in its soluble ferric (3+) form, a reflection of the anoxia typifying these black-mud pore-brines.

    Although the bottom brines are perennially anoxic, the levels of organic matter in the laminites are less than 1.2% (Tribovillard et al., 2009). Marine-derived amorphous organic matter dominates the organic content. However, the organic assemblage is unexpectedly degraded in terms of hydrogen content, which may be accounted for by a relatively long residence time of organic particles at the halocline-pycnocline. It seems the organic particles are temporarily trapped at the halocline and sokept in contact with the dissolved oxygen-rich overlying water mass.


    Mediterranean Ridge Accretionary Wedge

    Deep Hypersaline Anoxic Basins (DHABs) in the Mediterranean Sea are mostly located south of Crete between Greece and the North African coast of Libya (ranging from 34°17’N; 20°0’E to 33°52’N; 26°2’E from west to east) at a depth of 3000-4000 m. In the last few decades a number of salty basin areas have been discovered, namely; L’Atalante, Urania, Discovery, Bannock, Tyro, Thetis, Medee and Kryos basins (Figure 4).

    The brines that create these hypersaline anoxic seafloor depressions first formed as thick salt beds accumulated during the deep drawdown of the Mediterranean Sea some 5.45 million years ago, in an event known as the Messinian Salinity Crisis. A few million years later, ongoing basin closure along the Mediterranean suture and uplift of the Mediterranean Ridge drove inversion of some  portions of the buried salt. This brought thick salt masses back into the marine phreatic, where the evaporites began to dissolve, more rapidly from the upper edges of the Messinian salt mass. And so, hypersaline brine haloes ultimately vented onto the seafloor.

    The various brine lakes on the deep-sea floor of the Mediterranean, today occur thousands of metres below the photic zone, within depressions entraining bottom lake brine chemistries up to ten times as saline as Mediterranean seawater (Figure 4). In the Bannock region, the various brine-filled depressions or sub-basins create a closed outer moat around a central seafloor mound that is 10 km across (Figure 5a). The chemical composition of the Tyro Basin bottom brine is related to the dissolution of the underlying halite-dominated evaporites, while the MgCl2 dominant chemical composition of the Libeccio Basin in the Bannock area, with its elevated salinities approaching 400‰, imply derivation from dissolving bittern salts (de Lange et al., 1990). In the L' Atalante region, sodium chloride is predominantly sourced in the L’Atalante and Urania lakes, but L’Atalante is much richer in potassium chloride than the other nearby lakes.


    The Libeccio Basin (aka Bannock Basin)is almost exclusively the product of dissolution of magnesium chloride (bischofite) salts (Figure 5b). The bottom brine has a density of 1330 kg/m3, which makes it the densest naturally-occurring brine yet discovered in the marine realm (Wallmann et al., 2002). Its concentration profile in sediment beneath the brine lake shows the age of this lake is between 700 and 2000 yr. The high concentration of magnesium chloride drives the dissolution of biogenic calcium carbonate, but simultaneously facilitates excellent preservation of siliceous microfossils and organic matter. In the basin bottom muds there are large euhedral crystals of gypsum, up to 10cm across, precipitating from these magnesium chloride brines (Cita 2006).

    Biomarker associations in organics accumulating in the Mediterranean brine lakes define two depositional styles: typical marine and hypersaline (Burkova et al., 2000). Algal and bacterial biomarkers typical of saline environments are found in layers some 0.60 to 0.75 m below the sediment surface in the Tyro Lake Basin, as well as normal marine biomarkers derived from pelagic fallout (“rain from heaven”) in the same bottom sediments. Saline indicators include; regular C-25 isoprenoids, squalane, lycopane, isolycopane, tetraterpenoid and tetrapyrrolic pigments, monoalkylcyclohexanes, tricyclic diterpanes, steranes, hopanes, bio- and geohopanes. According to Burkova et al. (2000), the saline organic signatures come from microbial mat layers, redeposited from a Messinian source into the sapropels of the modern depression. Alternatively, they may indicate the activities of a chemoautotrophic community, which flourishes at the halocline or in rims around active brine vents. As in the Orca Basin, the organic content of the bottom sediments of the Mediterranean brine pools is much higher than is typical for deep seafloor sediment (Figure 6b).


    Anoxic hypersaline brines in Mediterranean brine lakes are highly sulphidic and among the most sulphidic bodies of marine water world-wide; in many lakes across the region H2S concentrations are consistently greater than 2-3 mmol (Table 1;  Henneke et al., 1997). The brine body below the Urania chemocline is more than 100 m thick and contains up to 11 mM hydrogen sulphide, making it the most sulphidic water body in the known marine realm. In combination with the sulphide are very high levels of methane both in and below the halocline (≈5.56 mM; Borin et al., 2009). In contrast, there is little to no H2S in the anoxic bottom brine of the Orca Basin (Table 1). There the iron concentration is 2 ppm, a value more than 1000 times higher than in the overlying Gulf of Mexico seawater. Such high levels of reducible iron in the Orca Basin are thought to explain the lack of H2S in the bottom brine and a preponderance of framboidal pyrite along with extractable iron in the bottom sediments (Sheu, 1987). Both the Orca Basin and the brine pools on the floor of the Mediterranean, show sulphate levels that can be more than twice that in the overlying seawater.

    Red Sea Deeps are DHALs

    Today the deep axial part of the Red Sea rift is characterised by a series of brine filled basins or deeps (Figure 7). Surrounding these deeps, the rift basement is covered by a thick sequence of middle Miocene evaporites precipitated in an earlier hydrographically isolated stage of rifting (Badenian – Middle Miocene). In the Morgan basin in the southern Red Sea the maximum thickness of rift-fill sediments, including halokinetic salt, is around 8,000m (Figures 7, 8, 9; Ehrhardt et al., 2005). Girdler and Whitmarsh (1974) conclude that Miocene evaporites first accumulated on Red Sea transitional crust but must have later flowed down-dip into now cover parts of the axial zone (basaltic) of the Plio-Pleistocene oceanic crust. At latitudes of 20° to 23° N, transform fracture zones provide focused passage-ways for such into-the-basin salt flow.


    Thick flowing halite enables the involvement of dissolving salt in axial hydrothermal circulation, so producing pools of dense hot brines and the topographic isolation of spreading segments into a series of evaporite-enclosed deeps (Figure 7; Feldens and Mitchell, 2015). Today, flow-like features, cored by Miocene evaporites, are situated along the axis of the Red Sea atop younger magnetic seafloor spreading anomalies. However, not all brine seeps occur in or near the deep axis of the Red Sea on the downdip edge of flowing Miocene salt, some occur in much shallower suprasalt positions sediment-floored nearer the coastal margins of the Red Sea, in waters just down-dip of actively-growing well-lit coral reefs (Batang et al., 2012).


    Six salt flows, most showing rounded fronts in plan-view, with heights of several hundred meters and widths between 3 and 10 km, are seen in high-resolution bathymetry and DSDP core material collected around the Thetis and Atlantis II deeps and between the Atlantis II Deep and the Port Sudan Deep (Figure 9; Feldens and Mitchell, 2015; Augustin et al., 2014; Mitchell et al., 2010). Relief on the underlying volcanic basement surface likely controls the positions of individual salt flow lobes. On the flow surfaces, along-slope and downslope ridge and trough morphologies have developed parallel to the local seafloor gradient, presumably due to the extension of the hemiplegic sediment cover or strike-slip movement within the evaporites.

    The local topographies of these salt flows, and the orientation of longitudinal ridges and troughs, indicate their downslope senses of flow. Where two allochthon tongues meet in the central rift, they form a suture along which the salt may turn to then flow parallel to the suture axis (Figure 9). Many volcanic ridges and fault scarps terminate where smooth rounded-lobes front salt, which then flows around obstructions in the basement (like volcanoes) to onlap them. The entire region between 23°N and 19°N shows signs of salt flow with no fault traces seen in areas covered by salt, which is up to 800 m thick (Augustin et al., 2014). Most normal faults, folds, and thrust fronts are parallel or perpendicular to the direction of maximum seabed gradient, while strike-slip shears tend to trend downslope.


    Dissolution of shallow, halokinetic, near-seafloor halite means that today, beneath more than a kilometre of seawater, there are 26 brine pools and deeps, some of which are underlain by metalliferous sediments (Figure 7; Blanc and Anschutz 1995, Blum and Puchelt, 1991). Because of varying size, age, and formation history across the various deeps, Ehrhardt and Hübscher (2015) discriminate between central and northern Red Sea deeps. The larger central Red Sea deeps are located in the axial trough and are separated by inter-trough zones. Young basaltic crust floors them and exhibits magnetic anomalies not older than 1.7 Ma. The northern Red Sea deeps are smaller and form only isolated deeps within the axial depression. Volcanic activity accompanies some of them. Many of the central Red Sea deeps contain bottom-water brines and metalliferous sediments, pointing to the hydrothermal circulation of seawater below a focusing salt layer (Schmidt et al., 2015). The largest and most prominent deep is the Atlantis II Deep, located in the central part of the Red Sea, in the vicinity of other large deeps such as the Chain Deep and Discovery Deep. Other prominent deeps further north are the Tethys and Nereus Deeps, but these deeps are still in the central part of the Red Sea (Figure 7).

    There are two types of brine-filled ocean deeps in the deeper parts of the salt-floored parts off the Red Sea: (a) volcanic and tectonically impacted deeps that opened by a lateral tear in the Miocene evaporites and Plio-Quaternary overburden; (b) non-volcanic deeps built by subsidence of Plio-Quaternary sediments due to evaporite subrosion (dissolution) processes. Type b) deeps develop as evaporite collapse structures (Figure 7: Ehrhardt and Hübscher, 2015). In contrast, the type (a) volcanic deeps can be correlated with their positions in NW–SE-oriented segments of the Red Sea, which are regions off "daylighted" volcanic segments. The N–S segments, between these volcanically active NW–SE segments, are called  “non-volcanic segment” as no volcanic activity is known (Ehrhardt and Hübscher, 2015). The interpreted lack of volcanism is in agreement with associated magnetic data that shows no major anomalies. Accordingly, the deeps in the “nonvolcanic segments” are evaporite collapse-related structures creating discontinuities and brine breakout zones in and atop the salt sheets, without the need for a seafloor spreading cell.

    However, evaporite collapse-type ocean deeps are not limited to the non-volcanic segments, subrosion processes driven by upwells in hydrothermal circulation are possible in any part of the axial depression, especially along fault damage zones. The combined interpretation of bathymetry and seismic reflection profiles gives a further insight into the nature of lateral salt gliding in the Red Sea. Salt rises are typically present where the salt flows above basement faults. The internal reflection character of the salt changes laterally from reflection-free to stratified, which suggests significant salt deformation during the salt deposition, as in the Santos Basin in the Aptian Atlantic salt province Warren, 2016). Acoustically-transparent halokinetic halite accumulated locally as evolving rim synclines were filled by stratified evaporite-related facies (Figure 10). Both types of deeps, as defined by Ehrhardt and Hübscher (2005), are surrounded by thick halokinetic masses of Miocene salt, with brine chemistries in the bottom brine layer signposting ongoing halite subrosion and dissolution.


    Red Sea deeps were discovered in the 1960s at a time when lateral translation of salt (gliding and spreading) and the formation of density stratification in deep-seafloor hypersaline anoxic lakes (DHALS) were not known (Warren, 2016). Today, with our knowledge of seeps and hypersaline seafloor depressions in halokinetic terranes on the slope and rise in the Gulf of Mexico and accretionary ridges in the parts of the Mediterranean Sea, we now know that the brine-filled deeps on the floor of the Red Sea are just another example of DHALs. What is most interesting in the Red Sea Dhals is the chemical make-up of a few deeps, with inherent elevated levels of iron, copper and lead, especially in the Atlantis II deep, which lies in one of the deeper and most hypersaline sets of linked brine lake depressions known  (Figure 9b). The association of copper-zinc hydrothermal mineralisation in the Atlantis II Deep was discussed in an earlier Salty Matters article (see April 29, 2016).

    In the last 28,000 years some 10 to 30 metres of the oxidic-silicatic-sulphidic laminites, along with hydrothermal anhydrites, have accumulated beneath the Atlantis II brine lake, atop a basement composed of a mixture of basaltic ridges and halokinetic salt (Figure 10b; Shanks III and Bischoff, 1980; Pottorf and Barnes, 1983; Anschutz and Blanc, 1995; Mitchell et al., 2010; Feldens et al., 2012). Metalliferous sediments beneath the floor of the deep are composed of stacked delicately banded (laminated)  mudstones with bright colours of red, yellow, green, purple, black or white. The colours indicate varying levels of oxidised or reduced iron and manganese, related to varying oxidation levels and salinities in the overlying brine column. Sediments in the laminites are typically anhydritic and very fine-grained, with 50-80% of the sediment less than 2µm in size. Intercrystalline pore brines constitute up to 95 wt% of the muds, with measured pore salinities as much as 26 wt% and directly comparable to the salinity/density of the overlying brine layer (Figure 11; Pottorf and Barnes, 1983).


    The sulphide-rich layers are a metre to several metres thick and form laterally continuous beds several kilometres across. Sulphides are dominated by very fine-grained pyrrhotite, cubic cubanite, chalcopyrite, sphalerite, and pyrite, and are interlayered with iron-rich phyllosilicates (Zierenberg and Shanks, 1983). Sulphur isotope compositions and carbon-sulphur relations indicate that some of these sulphide layers have a hydrothermal seawater component, whereas others were formed by bacterial sulphate reduction centred in the halocline interface. Ongoing brine activity began in the western part of the Deep some 23,000 years ago with deposition of a lower and upper sulphide zone, and an intervening amorphous silicate zone (Figure 11). The metalliferous and nonmetalliferous sediments in the W basin accumulated at similar rates, averaging 150 kg/k.y./m2, while metalliferous sediments in the SW basin accumulated at a higher rate of 700 kg/k.y./m2 (Figure 11; Anschutz and Blanc, 1995). The lowermost unit in the sediment pile in the W basin consists mainly of detrital biogenic carbonates, with occasional thin beds of red iron oxides (mostly fine-grained hematite) or dark interbeds entraining sulphide minerals.

    Hydrothermal anhydrite in the Atlantis II sediments occurs both as at-surface nodular hydrothermal beds around areas where hot fluid discharges onto the sea floor and as vein fills beneath the sea floor (Degens and Ross 1969, Pottorf and Barnes 1983, Ramboz and Danis 1990, Monnin and Ramboz 1996). White nodular to massive anhydrite beds in the W basin are up to 20 cm thick and composed of 20-50 µm plates and laths of anhydrite, typically interlayered with sulphide and Fe-montmorillonite beds. The central portion of individual anhydrite crystals in these beds can be composed of marcasite. The lowermost bedded unit in the SW basin contains much more nodular anhydrite, along with fragments of basalt toward its base. Its 4-metre+ anhydritic stratigraphy is not unlike that of nodular sekko-oko ore in a Kuroko deposit, except that any underlying volcanics are basaltic rather than felsic (see Chapter 16; Warren, 2016).

    The anhydrite-filled veins that crosscut the cored laminites acted as conduits by which hot, saline hydrothermal brines vent onto the floor of the Deep. Authigenic talc and smectite dominate in deeper, hotter vein fills, while shallower veins are rich in anhydrite cement (Zierenberg and Shanks III, 1983). The vertical zoning of vein-mineral fill is related to heating haloes, tied the same ascending hydrothermal fluids, with stable isotope ratios in the various vein minerals indicating precipitation temperatures ranging up to 300°C.

    Because of anhydrite’s retrograde solubility, it can form by a process as simple as heating hydrothermally-circulating seawater to temperatures over 150°C. Pottorf and Barnes (1983) concluded that the bedded anhydrite of the Atlantis II Deep, like the vein fill, is a hydrothermal precipitate. Based on marcasite inclusions in the anhydrite units, it precipitated at temperatures down to 160°C or less. At some temperature between 60 and 160°C, probably close to 100-120°C, hydrothermal anhydrite precipitation ceased. Thus, anhydrite distribution in the Atlantis II deep is related to the solution mixing and thermal anomalies associated with hydrothermal seawater circulation.

    The fact that Holocene sediments in the Atlantis II Deep contain sulphate minerals and that particulate anhydrite is still suspended in the lower brine body strongly suggests that anhydrite is stable in the temperatures found at the bottom of the water column or is at least only dissolving slowly. These conclusions were clarified by Monnin and Ramboz (1996), who found that the Upper Convective Layer (UCL; or Transition Zone) of the Atlantis II hydrothermal system was undersaturated with respect to hydrothermal anhydrite throughout their study period, 1965-1985. The system reached anhydrite saturation in the lower brine only for short periods in 1966 and 1976.


    Dead Sea (partial continental DHAL counterpart)

    The Dead Sea depression is a large strike-slip basin located within the Dead Sea transform; it lies in a plate boundary separating the Arabian plate from the African plate and connects the divergent plate boundary of the Red Sea to the convergent plate boundary of the Taurus Mountains in southern Turkey (Figure 12). Since the fault first formed, 105 km of left-lateral horizontal movement has occurred along the transform. In places along the transform where the crust is stretched or attenuated, plate stress is accommodated via several rapidly subsiding en-echelon rhomb-shaped grabens separated across west-stepping fault segments. The Dead Sea basin and the Gulf of Elat to its south are the largest of these graben depressions and are separated by the Yotvata Playa basin. The Dead Sea basin fill is 110 km long, 16 km wide and 6–12 km deep and located in the offset between two longitudinal faults, the Arava Fault and the Western Boundary (Jericho) Fault (Figure 12a, b; Garfunkel et al., 1981; Garfunkel and Ben-Avraham, 1996).


    Movement began 15 Ma in the Miocene with the opening of the Red Sea and is continuing today at a rate of 5 to 10 mm/yr. The Dead Sea basin floor is more strongly coupled to the western margin (Levantine plate), which is being left behind by the northward-moving Arabian plate (Figure 12b). Since the Miocene, depocentres in the Dead Sea region have moved 50 km northward along the shear zone (Zak and Freund, 1981) to create the offlapping style of sedimentation in the Dead Sea–Arava Valley, with a basin geometry reminiscent of the Ridge Basin in California. Continued extensional movement has triggered halokinesis in the underlying Miocene evaporites so that diapirs subcrop along the Western Boundary Fault and its offshoots (Figures 12b, 13; Neev and Hall, 1979; Smit et al., 2008). Salt in these structures is equivalent to the salt in the outcropping Mount Sedom diapir (Alsop et al., 2015).

    In the late Miocene (8-10 Ma), differential uplift along the transform edges and rapid subsidence of the basin led to a deep topographic trough. During this second stage (4-6 Ma) the trough was invaded by Mediterranean seawater, perhaps through the Yizre’el Valley, to create a highly restricted seepage arm that was periodically cut off from the ocean and so deposited a 2-3 km thick sequence of halite-rich evaporites that constitute the Sedom Formation (also known as the Usdum Fm.). This 2 to 3 km-thick section is now halokinetic in the Dead Sea region.

    Unlike the marine isotopic signatures of the salts in the Sedom Formation, isotopes in the evaporites of the various Pleistocene sequences in the Dead Sea depression indicate their precipitation from lacustral CaCl-rich connate brines. Groundwater inflow chemistries are created by rock-water interactions with original connate seawater brines, first trapped in sediments of the rift walls in “Sedom time” (Stein et al., 2000). After the final Pliocene disconnection from the sea and a lowering of the lake levels, these residual brines gradually seeped and leached back into the Sedom basin. At the same time, rapid accumulation of Amora and Samra sediments within a subsiding and extending valley, atop thick-bedded evaporites of the Sedom Fm. initiated several salt diapirs along the valley floor, the best known being Mt. Sedom (Figure 13b; Alsop et al., 2015; Smit et al., 2008; Larsen et al., 2002). Today the Mount Sedom diapir has pierced the surface atop a 200 m-high salt wall. Throughout the Holocene, salt has been rising in Mt. Sedom at a rate of 6-7 mm a-1 (Frumkin, 1994). The nearby Lisan ridge is also a topographic high underlain by halokinetic Sedom salt.

    Study of the halokinetic stratigraphy of Mt Sedom salt wall shows the structure has a moderate-steep west dipping western margin and an overturned (west-dipping) eastern flank (Figure 13b; Alsop et al., 2015). The sedimentary record of passive wall growth includes sedimentary breccia horizons that locally truncate underlying beds and are interpreted to reflect sediments having been shed off the crest of the growing salt wall. Structurally, the overturned eastern flank is marked by upturn within the overburden, extending for some 300 m from the salt wall. Deformation within the evaporites is characterised by ductile folding and boudinage, while a 200 m thick clastic unit within the salt wall formed a tight recumbent fold traceable for 5 km along strike and associated with a 500 m wide inverted limb. This overturned gently-dipping limb is marked by NE-directed folding and thrusting, sedimentary injections, and a remarkable attenuation of the underlying salt from ≈380 m to >20 m over just 200 m of strike length. The inverted limb is overlain by an undeformed anhydrite, gypsum and clastic caprock, thought to be the residue from a now-dissolved salt sheet that extruded over the top of the fold.

    Expulsion of salt down the regional slope towards the NE, combined with subsequent dissolution of evaporites, likely resulted in the local ‘pinching shut’ of the salt wall aperture, leading to its distinctive hour-glass map pattern. The pinched area also coincides with deposition of a thicker overlying clastic sequence, indicating continued subsidence of this part of the salt wall. The dissolution of the salt tongue, as well as other shallow salt, has contributed significant volumes of dissolved salt to the Dead Sea brine system so creating and maintaining the large halite-precipitating perennial saline lake in the basin sump

    Unlike the longterm stability of the deep seawater-covered top to a salt-karst induced density-stratified brine lake defining a classic oceanic DHAL hydrology, the continental setting of the Dead Sea salt-karst brine-sump means sediments accumulating below the perennial brine mass in the Dead Sea are deposited with a range of brine-pool bottom textures indicative of the presence for absence of a less saline uppermost brine mass (Figures 14, 15;Charrach, 2018; Sirota et al., 2017; Alsop et al., 2016; Kiro et al., 2015; Neugebauer et al., 2014).



    Since the beginning of the 20th century the water budget of the Dead Sea has been negative, leading to a continuous decrease in the water level. The extensive evaporation in the absence of major fresher water input led to an increase in the density of the upper water layer, which caused the lake to overturn in 1979 (Warren, 2016 for summary of the hydrochemical evolution). Since then, except after two rainy seasons in 1980 and 1992, the Dead Sea remained holomictic and has been characterized by a NaCl supersaturation and halite deposition on the lake bottom, with total dissolved salt concentrations reaching 347 g/l. Due to the continuous evaporation of the Dead Sea, Na+ precipitates out as halite, while Mg2+, whose salts are more soluble, is further concentrated and has become the dominant cation in the present holomictic water mass (Table 1).


    In situ observations in the Dead Sea by Sirota et al., 2017, within the current holomictic hydrology of the Dead Sea, link seasonal thermohaline stratification, halite saturation, and the the textural characterist of the actively forming halite-rich bottom sediments . The spatiotemporal evolution of halite precipitation in the current holomictic stage of the Dead Sea is influenced by (1) lake thermohaline stratification (temperature, salinity, and density), (2) degree of halite saturation, and (3) textural evolution of the active halite deposits. Observed relationships by Sirota et al., tie the textural characteristics of layered subaqueous halite deposits (i.e., grain size, consolidation, and roughness) to the degree of saturation, which in turn reflects the limnology and hydroclimatology of the lake sump. The current halite-accumulating lake floor is divided into two principal environments: 1) a deep, hypolimnetic (below thermocline) lake floor and, 2) a shallow, epilimnetic lake floor(above thermocline) (Figure 15).

    In the deeper hypolimnetic lake floor, halite, which is a prograde salt,  continuously precipitates with seasonal variations so that : (a) During summer, consolidated coarse halite crystals under slight supersaturation form rough crystal surfaces on the deep lake floor. (2) During the cooler conditions of winter, unconsolidated, fine halite crystals form smooth lake-floor deposits under high supersaturation. These observations support interpretations of the seasonal alternation of halite crystallisation mechanisms. The shallow epilimnetic lake floor is highly influenced by the seasonal temperature variations, and by intensive summer dissolution of part of the previous year’s halite deposit, which results in thin sequences with annual unconformities. This emphasises the control of temperature seasonality on the characteristics of the precipitated halite layers. In addition, precipitation of halite on the hypolimnetic floor, at the expense of the dissolution of the epilimnetic floor, results in lateral focusing and thickening of halite deposits in the deeper part of the basin and thinning of the deposits in shallow marginal basins.

    Implications

    All DHALs, either in a classic marine deep anoxic seafloor setting or a continental setting, require karstification of a shallowly buried halokinetic salt mass and a topographic depression capable of longterm retention of brine in the landscape. DHALs on the deep seafloor can create their topographic sumps via salt withdrawal (the Gulf of Mexico and the Red Sea) or regional tectonism as in The Mediterranean Ridges and the Dead Sea.

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