Salty Matters

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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



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.


    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.


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

    John Warren - Sunday, September 30, 2018



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

    John Warren - Saturday, December 31, 2016

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

    Extrasalt gas and brine intersections

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

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

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


    Caprock and other salt periphery-held gases

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

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

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

    Intrasalt gas

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

    Gas types and sources at the local and basin scale

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

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

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

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

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

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


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


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    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.

    Tikhomirov, V. V., 2014, Molecular nitrogen in salts and subsalt fluids in the Volga-Ural Basin: Geochemistry International, v. 52, p. 628-642.

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

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



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

    John Warren - Wednesday, November 30, 2016

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

    What’s the gas?

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

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


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

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


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


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

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


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

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


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

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


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

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

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

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


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

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

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

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

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

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

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

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

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

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

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

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

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

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

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


    Anderson, S. B., and W. P. Eastwood, 1968, Natural Gas in North Dakota, Natural Gases of North America, Volume Two, American Association of Petroleum Geologists Memoir 9, p. 1304-1326.

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    Gases in Evaporites; Part 1 - Rockbursts and gassy outbursts

    John Warren - Monday, October 31, 2016

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

    Gases in evaporites can create problems

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

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

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


    Gassy salt (knistersalz)

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

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

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

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

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


    Gassy outbursts and rockbursts in salt

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

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

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

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

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

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

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

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

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

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


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


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

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

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


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