Salty Matters

The Blog is written by me, John Warren. Once every three or four weeks or so I will post an article or two on an evaporite topic that has piqued my interest. On the Saltwork Publications webpage (under "the Works") there is a growing library of pdfs and epubs based on these blogs. These articles on the website have much higher resolution extractable graphics in than in the blog. There is also a link to this set of pdfs and epubs on the home page (www.saltworkconsultants.com).

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

John Warren - Sunday, September 30, 2018

 


Introduction

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

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

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

Along the brine pool edge in the Gulf of Mexico

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

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


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

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


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

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

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

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

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

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

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


Hardgrounds, seafloor stability & stable isotopes

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

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


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

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

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

Clathrates and methane seeps in the Gulf of Mexico

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


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


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

Brine lake biota in the Mediterranean Ridges

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

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

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

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

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


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


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

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


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

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

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

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

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

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

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

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


Life in the Red Sea brine deeps

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

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


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

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

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

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

Life in the Dead Sea

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

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

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

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


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

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

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


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

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

Salt dissolution, seafloor salinity and halophilic extremophile populations

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

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


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

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

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

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

References

Albuquerque, L., M. Taborda, V. La Cono, M. Yakimov, and M. S. da Costa, 2012, Natrinema salaciae sp. nov., a halophilic archaeon isolated from the deep, hypersaline anoxic Lake Medee in the Eastern Mediterranean Sea: Systematic and Applied Microbiology, v. 35, p. 368-373.

Antunes, A., D. K. Ngugi, and U. Stingl, 2011, Microbiology of the Red Sea (and other) deep-sea anoxic brine lakes: Environmental Microbiology Reports, v. 3, p. 416-433.

Augustin, N., F. M. van der Zwan, C. W. Devey, M. Ligi, T. Kwasnitschka, P. Feldens, R. A. Bantan, and A. S. Basaham, 2016, Geomorphology of the central Red Sea Rift: Determining spreading processes: Geomorphology, v. 274, p. 162-179.

Bergquist, D. C., T. Ward, E. E. Cordes, T. McNelis, S. Howlett, R. Kosoff, S. Hourdez, R. Carney, and C. R. Fisher, 2003, Community structure of vestimentiferan-generated habitat islands from Gulf of Mexico cold seeps: Journal of Experimental Marine Biology and Ecology, v. 289, p. 197-222.

Bergquist, D. C., F. M. Williams, and C. R. Fisher, 2000, Longevity record for deep-sea invertebrate: Nature, v. 403, p. 499-500.

Bernhard, J. M., C. R. Morrison, E. Pape, D. J. Beaudoin, M. A. Todaro, M. G. Pachiadaki, K. A. Kormas, and V. P. Edgcomb, 2015, Metazoans of redoxcline sediments in Mediterranean deep-sea hypersaline anoxic basins: BMC Biology, v. 13, p. 105.

Borin, S., L. Brusetti, F. Mapelli, G. D'Auria, T. Brusa, M. Marzorati, A. Rizzi, M. Yakimov, D. Marty, G. J. De Lange, P. Van der Wielen, H. Bolhuis, T. J. McGenity, P. N. Polymenakou, E. Malinverno, L. Giuliano, C. Corselli, and D. Daffonchio, 2009, Sulfur cycling and methanogenesis primarily drive microbial colonization of the highly sulfidic Urania deep hypersaline basin: Proceedings of the National Academy of Sciences, v. 106, p. 9151-9156.

Bougouffa, S., J. K. Yang, O. O. Lee, Y. Wang, Y. Batang, A. Al-Suwailem, and G. Qian, 2013, Distinctive Microbial Community Structure in Highly Stratified Deep-Sea Brine Water Columns: Appl. Environ. Microbiol., v. 79, p. 3425-3437.

Brusa, T., E. Del Puppo, A. Ferrari, G. Rodondi, C. Andreis, and S. Pellegrini, 1997, Microbes in deep-sea anoxic basins: Microbiol.Res., v. 152, p. 45-56.

Burkova, V. N., E. A. Kurakolova, N. S. Vorob'eva, M. L. Kondakova, and O. K. Bazhenova, 2000, Hydrocarbons of the hypersaline environment of the Tyro deep-sea depression (eastern Mediterranean): Geochemistry International, v. 38, p. 883-894.

Camerlenghi, A., 1990, Anoxic Basins of the eastern Mediterranean: geological framework: Marine Chemistry, v. 31, p. 1-19.

Canet, C., R. M. Prol-Ledesma, E. Escobar-Briones, C. Mortera-Gutierrez, R. Lozano-Santa Cruz, C. Linares, E. Cienfuegos, and P. Morales-Puente, 2006, Mineralogical and geochemical characterization of hydrocarbon seep sediments from the Gulf of Mexico: Marine and Petroleum Geology, v. 23, p. 605-619.

Cita, M. B., 2006, Exhumation of Messinian evaporites in the deep-sea and creation of deep anoxic brine-filled collapsed basins: Sedimentary Geology, v. 188-189, p. 357-378.

Cita, M. B., K. A. Kastens, F. W. McCoy, F. Aghib, A. Cambi, A. Camerlenghi, C. Corselli, E. Erba, M. Giambastiani, T. Herbert, C. Leoni, P. Malinverno, A. Nosetto, and E. Parisi, 1985, Gypsum precipitation from cold brines in an anoxic basin in the eastern Mediterranean: Nature (London), v. 314, p. 152-154.

Cordes, E. E., D. C. Bergquist, and C. R. Fisher, 2009, Macro-Ecology of Gulf of Mexico Cold Seeps, Annual Review of Marine Science: 1, p. 143-168.

Corselli, C., and F. S. Aghib, 1987, Brine formation and gypsum precipitation in the Bannock Basin (eastern Mediterranean): Marine Geology, v. 75, p. 185-199.

Dalthorp, M., and T. H. Naehr, 2011, Structural and Stratigraphic Relationships of Hydrocarbon Seeps in the Northern Gulf of Mexico and Geological Factors Contributing to Migration Variations: Gulf Coast Association of Geological Societies Transactions, v. 61, p. 105-122.

Danovaro, R., A. Dell'Anno, A. Pusceddu, C. Gambi, I. Heiner, and R. Møbjerg Kristensen, 2010, The first metazoa living in permanently anoxic conditions: BMC Biology, v. 8, p. 30.

Danovaro, R., C. Gambi, A. Dell’Anno, C. Corinaldesi, A. Pusceddu, R. C. Neves, and R. M. Kristensen, 2016, The challenge of proving the existence of metazoan life in permanently anoxic deep-sea sediments: BMC Biology, v. 14, p. 43.

Dattagupta, S., L. L. Miles, M. S. Barnabei, and C. R. Fisher, 2006, The hydrocarbon seep tubeworm Lamellibrachia luymesi; primarily eliminates sulfate and hydrogen ions across its roots to conserve energy and ensure sulfide supply: Journal of Experimental Biology, v. 209, p. 3795.

de Lange, G. J., J. J. Middleburg, C. H. van der Weijden, G. Catalano, G. W. Luther, III, D. J. Hydes, J. R. W. Woittiez, and G. P. Klinkhammer, 1990, Composition of anoxic hypersaline brines in the Tyro and Bannock Basins, eastern Mediterranean: Marine Chemistry, v. 31, p. 63-88.

Edgcomb, P. V., and M. J. Bernhard, 2013, Heterotrophic Protists in Hypersaline Microbial Mats and Deep Hypersaline Basin Water Columns: Life, v. 3.

Erba, E., 1991, Deep mid-water bacterial mats from anoxic basins of the eastern Mediterranean: Marine Geology, v. 100, p. 83-101.

Erba, E., G. Rodondi, E. Parisi, H. L. Ten Haven, M. Nip, and J. W. De Leeuw, 1987, Gelatinous pellicles in deep anoxic hypersaline basins from the Eastern Mediterranean: Marine Geology, v. 75, p. 165-183.

Feng, D., H. H. Roberts, P. Di, and D. Chen, 2009, Characteristics of hydrocarbon seep-related rocks from the deep Gulf of Mexico: Gulf Coast Association of Geological Societies Transactions, v. 59, p. 271-275.

Ferrer, M., J. Werner, T. N. Chernikova, R. Bargiela, L. Fernández, V. La Cono, J. Waldmann, H. Teeling, O. V. Golyshina, F. O. Glöckner, M. M. Yakimov, P. N. Golyshin, and M. S. C. The, 2012, Unveiling microbial life in the new deep-sea hypersaline Lake Thetis. Part II: a metagenomic study: Environmental Microbiology, v. 14, p. 268-281.

Fisher, C. R., I. R. MacDonald, R. Sassen, C. M. Young, S. A. Macko, S. Hourdez, R. S. Carney, S. Joye, and E. McMullin, 2000, Methane Ice Worms: Hesiocaeca methanicola Colonizing Fossil Fuel Reserves: Naturwissenschaften, v. 87, p. 184-187.

Freytag, J. K., P. R. Girguis, D. C. Bergquist, J. P. Andras, J. J. Childress, and C. R. Fisher, 2001, A paradox resolved: Sulfide acquisition by roots of seep tubeworms sustains net chemoautotrophy: Proceedings of the National Academy of Sciences of the United States of America, v. 98, p. 13408-13413.

Fu Chen, D., Q. Liu, Z. Zhang, L. M. Cathles Iii, and H. H. Roberts, 2007, Biogenic fabrics in seep carbonates from an active gas vent site in Green Canyon Block 238, Gulf of Mexico: Marine and Petroleum Geology, v. 24, p. 313-320.

Glover, A. G., A. J. Gooday, D. M. Bailey, D. S. M. Billett, P. Chevaldonne, A. Colaco, J. Copley, D. Cuvelier, D. Desbruyeres, V. Kalogeropoulou, M. Klages, N. Lampadariou, C. Lejeusne, N. Mestre, G. L. J. Paterson, T. Perez, H. Ruhl, J. Sarrazin, T. Soltwedel, E. H. Soto, S. Thatje, A. Tselepides, S. Van Gaever, and A. Vanreusel, 2010, Temporal change in deep-sea benthic ecosystems: a review of the evidence from recent time-series studies: Advances In Marine Biology, v. 58, p. 1-95.

Hartmann, M., J. C. Scholten, P. Stoffers, and K. F. Wehner, 1998, Hydrographic structure of the brine-filled deeps in the Red Sea - New results from the Shaban, Kebrit, Atlantis II, and Discovery deeps: Marine Geology, v. 144, p. 311-330.

Häusler, S., M. Weber, C. Siebert, M. Holtappels, B. E. Noriega-Ortega, D. De Beer, and D. Ionescu, 2014, Sulfate reduction and sulfide oxidation in extremely steep salinity gradients formed by freshwater springs emerging into the Dead Sea: FEMS Microbiol Ecol., v. 90.

Henneke, E., G. W. Luther, G. J. Delange, and J. Hoefs, 1997, Sulphur speciation in anoxic hypersaline sediments from the Eastern Mediterranean Sea: Geochimica et Cosmochimica Acta, v. 61, p. 307-321.

Ionescu, D., C. Siebert, L. Polerecky, Y. Y. Munwes, C. Lott, S. Häusler, M. B. Ionescu, J. Peplies, F. O. Glöckner, A. Ramette, T. Rödiger, T. Dittmar, A. Oren, S. Geyer, H.-J. Stärk, M. Sauter, T. Licha, J. B. Laronne, and D. de Beer, 2012, Microbial and Chemical Characterization of Underwater Fresh Water Springs in the Dead Sea. PLoS ONE 7(6): e38319. doi:10.1371/journal.pone.0038319: PlosOne, v. 7, p. e38319.

Joye, S. B., V. A. Samarkin, B. N. Orcutt, I. R. MacDonald, K.-U. Hinrichs, M. Elvert, A. P. Teske, K. G. Lloyd, M. A. Lever, J. P. Montoya, and C. D. Meile, 2009, Metabolic variability in seafloor brines revealed by carbon and sulphur dynamics: Nature Geoscience, v. 2, p. 349-354.

Kis-Papo, T., A. R. Weig, R. Riley, D. Peršoh, A. Salamov, H. Sun, A. Lipzen, S. P. Wasser, G. Rambold, I. V. Grigoriev, and E. Nevo, 2014, Genomic adaptations of the halophilic Dead Sea filamentous fungus Eurotium rubrum: Nature Communications, v. 5, p. 3745.

Kvitek, R. G., K. E. Coonlan, and P. J. Iampietro, 1998, Black pools of death: hypoxic, brine-filled ice gouge depressions become lethal traps for benthic organisms in a shallow Arctic embayment: Marine Ecology Progress Series, v. 162, p. 1-10.

MacDonald, I. R., 1992, Sea-floor brine pools affect behavior, mortality, and preservation of fishes in the Gulf of Mexico: lagerstatten in the making?: Palaios, v. 7, p. 383-387.

MacDonald, I. R., W. W. Sager, and M. B. Peccini, 2003, Gas hydrate and chemosynthetic biota in mounded bathymetry at mid-slope hydrocarbon seeps: Northern Gulf of Mexico: Marine Geology, v. 198, p. 133-158.

McMullin, E. R., K. Nelson, C. R. Fisher, and S. W. Schaeffer, 2010, Population structure of two deep sea tubeworms, Lamellibrachia luymesi and Seepiophila jonesi, from the hydrocarbon seeps of the Gulf of Mexico: Deep Sea Research Part I: Oceanographic Research Papers, v. 57, p. 1499-1509.

Milkov, A. V., and R. Sassen, 2001, Economic Geology of the Gulf of Mexico and the Blake Ridge Gas Hydrate Provinces: Gulf Coast Association of Geological Societies Transactions, v. 51, p. 219-228.

Nix, E. R., C. R. Fisher, J. Vodenichar, and K. M. Scott, 1995, Physiological ecology of a mussel with methanotrophic endosymbionts at three hydrocarbon seep sites in the Gulf of Mexico: Marine Biology, v. 122, p. 605-617.

Oren, A., 1993, The Dead Sea - Alive again: Experientia, v. 49, p. 518-522.

Oren, A., 1999, Bioenergetic aspects of halophilism: Microbiology & Molecular Biology Reviews, v. 63, p. 334-348.

Oren, A., 2005, A century of Dunaliella research: 1905-2005, in N. Gunde-Cimerman, A. Oren, and A. Plemenitaš, eds., Adaptation to Life at High Salt Concentrations in Archaea, Bacteria, and Eukarya: Dordrecht, Netherlands, Springer, p. 491-502.

Oren, A., 2015, Halophilic microbial communities and their environments: Current Opinion in Biotechnology, v. 33, p. 119-124.

Oren, A., D. Ionescu, M. Y. Hindiyeh, and H. I. Malkawi, 2008, Microalgae and cyanobacteria of the Dead Sea and its surrounding springs: Israel Journal of Plant Sciences, v. 56, p. 1-13.

Oren, A., and N. Ben Yosef, 1997, Development and spatial distribution of an algal bloom in the Dead Sea: A remote sensing study: Aquatic Microbial Ecology, v. 13, p. 219-223.

Oren, A., G. Bratbak, and M. Heldal, 1997, Occurrence of virus-like particles in the Dead Sea: Extremophiles, v. 1, p. 143-149.

Oren, A., and P. Gurevich, 1995, Dynamics of a bloom of halophilic Archaea in the Dead Sea: Hydrobiologia, v. 315, p. 149-158.

Oren, A., P. Gurevich, D. A. Anati, E. Barkan, and B. Luz, 1995, A bloom of Dunaliella parva in the Dead Sea in 1992: biological and biogeochemical aspects: Hydrobiologia, v. 297, p. 173-185.

Raiswell, R., and D. E. Canfield, 1998, Sources of iron for pyrite formation in marine sediments: American Journal of Science, v. 298, p. 219-245.

Reilly, J. F., I. R. MacDonald, E. K. Biegert, and J. M. Brooks, 1996, Geologic controls on the distribution of chemosynthetic communities in the Gulf of Mexico, in D. Schumacher, and M. A. Abrams, eds., Hydrocarbon Migration and its Near-Surface Expression, American Association of Petroleum Geologists Memoir 66, p. 39-62.

Sass, A. M., H. Sass, M. J. L. Coolen, H. Cypionka, and J. Overmann, 2001, Microbial communities in the chemocline of a hypersaline deep- sea basin (Urania basin, Mediterranean Sea): Applied and Environmental Microbiology, v. 67, p. 5392-5402.

Sassen, R., I. R. MacDonald, A. G. Requejo, N. L. Guinasso, Jr., M. C. Kennicutt, II, S. T. Sweet, and J. M. Brooks, 1994, Organic geochemistry of sediments from chemosynthetic communities, Gulf of Mexico slope: Geo-Marine Letters, v. 14, p. 110-119.

Schroeder, W., S. Brooke, J. Olson, B. Phaneuf, J. McDonough, III, and P. Etnoyer, 2005, Occurrence of deep-water Lophelia pertusa and Madrepora oculata in the Gulf of Mexico, in A. Freiwald, and J. M. Roberts, eds., Cold-Water Corals and Ecosystems: Erlangen Earth Conference Series, Springer Berlin Heidelberg, p. 297-307.

Sheu, D. D., 1987, Sulfur and organic carbon contents in sediment cores from the Tyro and Orca basins: Marine Geology, v. 75, p. 157-164.

Siam, R., G. A. Mustafa, H. Sharaf, A. Moustafa, A. R. Ramadan, A. Antunes, V. B. Bajic, U. Stingl, N. G. R. Marsis, M. J. L. Coolen, S. Mitchell, A. J. S. Ferreira, and H. El Dorry, 2012, Unique Prokaryotic Consortia in Geochemically Distinct Sediments from Red Sea Atlantis II and Discovery Deep Brine Pools: PLoS ONE, v. 7, p. e42872.

Smith, E. B., K. M. Scott, E. R. Nix, C. Korte, and C. R. Fisher, 2000, Growth and Condition of Seep Mussels (Bathymodiolus childressi) at a Gulf of Mexico Brine Pool: Ecology, v. 81, p. 2392-2403.

Steinle, L., K. Knittel, N. Felber, C. Casalino, G. de Lange, C. Tessarolo, A. Stadnitskaia, J. S. Sinninghe Damsté, J. Zopfi, M. F. Lehmann, T. Treude, and H. Niemann, 2018, Life on the edge: active microbial communities in the Kryos MgCl2-brine basin at very low water activity: The ISME Journal.

Thakur, N. K., and S. Rajput, 2011, Exploration of Gas Hydrates - Geophysical Techniques: Berlin, Springer, 281 p.

Wallmann, K., F. S. Aghi, D. Castradori, M. B. Cita, E. Suess, J. Greinert, and D. Rickert, 2002, Sedimentation and formation of secondary minerals in the hypersaline Discovery Basin, eastern Mediterranean: Marine Geology, v. 186, p. 9-28.

Wang, Y., J. T. Li, L. S. He, B. Yang, Z. M. Gao, H. L. Cao, Z. B. Batang, A. Al-Suwailem, and P.-Y. Qian, 2015, Zonation of Microbial Communities by a Hydrothermal Mound in the Atlantis II Deep (the Red Sea): PLoS ONE, v. 10, p. e0140766.

Warren, J. K., 2011, Evaporitic source rocks: mesohaline responses to cycles of “famine or feast” in layered brines, Doug Shearman Memorial Volume, (Wiley-Blackwell) IAS Special Publication Number 43, p. 315-392.

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

Yakimov, M. M., V. La Cono, R. Denaro, G. D'Auria, F. Decembrini, K. N. Timmis, P. N. Golyshin, and L. Giuliano, 2007, Primary producing prokaryotic communities of brine, interface and seawater above the halocline of deep anoxic lake L'Atalante, Eastern Mediterranean Sea: ISME J, v. 1, p. 743-755.

Yakimov, M. M., V. La Cono, V. Z. Slepak, G. La Spada, E. Arcadi, E. Messina, B. Mireno, L. S. Monticelli, D. Rojo, C. Barbas, O. V. Golyshina, M. Ferrer, P. N. Golyshin, and L. Giuliano, 2013, Microbial life in the Lake Medee, the largest deep-sea salt-saturated formation. : Scientific reports, 3; http://www.nature.com/srep/2013/131219/srep03554/full/srep03554.html.

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

 

Deepsea Hypersaline Anoxic Lakes & Basins (DHALs & DHABS)

John Warren - Friday, August 31, 2018

 

Introduction

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


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

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


Hydrology

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

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

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

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

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

Occurrences

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


Orca Basin, Gulf of Mexico

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

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

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


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

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


Mediterranean Ridge Accretionary Wedge

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

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

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


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

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


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

Red Sea Deeps are DHALs

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


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


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

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


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

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

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


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

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


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

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

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

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

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


Dead Sea (partial continental DHAL counterpart)

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


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

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

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

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

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

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



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


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

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

Implications

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

References

Addy, K. S., and E. W. Behrens, 1980, Time of accumulation of hypersaline anoxic brine in Orca basin (Gulf of Mexico): Marine Geology, v. 37, p. 241-252.

Alsop, G. I., S. Marco, R. Weinberger, and T. Levi, 2016, Sedimentary and structural controls on seismogenic slumping within mass transport deposits from the Dead Sea Basin: Sedimentary Geology, v. 344, p. 71-90.

Alsop, G. I., R. Weinberger, T. Levi, and S. Marco, 2015, Deformation within an exposed salt wall: Recumbent folding and extrusion of evaporites in the Dead Sea Basin: Journal of Structural Geology, v. 70, p. 95-118.

Augustin, N., C. W. Devey, F. M. van der Zwan, P. Feldens, M. Tominaga, R. A. Bantan, and T. Kwasnitschka, 2014, The rifting to spreading transition in the Red Sea: Earth and Planetary Science Letters, v. 395, p. 217-230.

Batang, Z. B., E. Papathanassiou, A. Al-Suwailem, C. Smith, M. Salomidi, G. Petihakis, N. M. Alikunhi, L. Smith, F. Mallon, T. Yapici, and N. Fayad, 2012, First discovery of a cold seep on the continental margin of the central Red Sea: Journal of Marine Systems, v. 94, p. 247-253.

Blanc, G., and P. Anschutz, 1995, New stratification in the hydrothermal brine system of the Atlantis II Deep, Red Sea: Geology, v. 23, p. 543-546.

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

Borin, S., L. Brusetti, F. Mapelli, G. D'Auria, T. Brusa, M. Marzorati, A. Rizzi, M. Yakimov, D. Marty, G. J. De Lange, P. Van der Wielen, H. Bolhuis, T. J. McGenity, P. N. Polymenakou, E. Malinverno, L. Giuliano, C. Corselli, and D. Daffonchio, 2009, Sulfur cycling and methanogenesis primarily drive microbial colonization of the highly sulfidic Urania deep hypersaline basin: Proceedings of the National Academy of Sciences, v. 106, p. 9151-9156.

Bregant, D., G. Catalano, G. Civitarese, and A. Luchetta, 1990, Some chemical characteristics of the brines in Bannock and Tyro Basins: salinity, sulphur compounds, Ca , F, pH, At, PO4, SiO2, NH3: Marine Chemistry, v. 31, p. 35-62.

Burkova, V. N., E. A. Kurakolova, N. S. Vorob'eva, M. L. Kondakova, and O. K. Bazhenova, 2000, Hydrocarbons of the hypersaline environment of the Tyro deep-sea depression (eastern Mediterranean): Geochemistry International, v. 38, p. 883-894.

Camerlenghi, A., 1990, Anoxic Basins of the eastern Mediterranean: geological framework: Marine Chemistry, v. 31, p. 1-19.

Charrach, J., 2018, Investigations into the Holocene geology of the Dead Sea basin: Carbonates and Evaporites.

Cita, M. B., 2006, Exhumation of Messinian evaporites in the deep-sea and creation of deep anoxic brine-filled collapsed basins: Sedimentary Geology, v. 188-189, p. 357-378.

de Lange, G. J., J. J. Middleburg, C. H. van der Weijden, G. Catalano, G. W. Luther, III, D. J. Hydes, J. R. W. Woittiez, and G. P. Klinkhammer, 1990, Composition of anoxic hypersaline brines in the Tyro and Bannock Basins, eastern Mediterranean: Marine Chemistry, v. 31, p. 63-88.

Degens, E. T., and D. A. Ross, 1969, Hot Brines and recent heavy metal deposits in the Red Sea: New York, N.Y., Springer Verlag, 600 p.

Ehrhardt, A., C. Hübscher, and D. Gajewski, 2005, Conrad Deep, Northern Red Sea: Development of an early stage ocean deep within the axial depression: Tectonophysics, v. 411, p. 19-40.

Feldens, P., and N. Mitchell, 2015, Salt Flows in the Central Red Sea, in N. M. A. Rasul, and I. C. F. Stewart, eds., The Red Sea: Springer Earth System Sciences, Springer Berlin Heidelberg, p. 205-218.

Frumkin, A., 1994, Hydrology and denudation rates of halite karst: Journal of Hydrology, v. 162.

Garfunkel, Z., and Z. Ben-Avraham, 1996, The structure of the Dead Sea: Tectonophysics, v. 155-176.

Garfunkel, Z., I. Zak, and R. Freund, 1981, Active faulting in the Dead Sea Rift: Tectonophysics, v. 62, p. 37-52.

Girdler, R. W., and R. B. Whitmarsh, 1974, 28. Miocene Evaporates in Red Sea Cores, Their Relevance to the Problem of the Width and Age of Oceanic Crust beneath the Red Sea: Woods Hole Oceanogr. Inst., Collect. Repr., v. 23, p. 913-922.

Henneke, E., G. W. Luther, G. J. Delange, and J. Hoefs, 1997, Sulphur speciation in anoxic hypersaline sediments from the Eastern Mediterranean Sea: Geochimica et Cosmochimica Acta, v. 61, p. 307-321.

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

Kiro, Y., S. L. Goldstein, B. Lazar, and M. Stein, 2015, Environmental implications of salt facies in the Dead Sea: Geological Society of America Bulletin.

Larsen, B. D., Z. Ben-Avraham, and H. Shulman, 2002, Fault and salt tectonics in the southern Dead Sea basin: Tectonophysics, v. 346, p. 71-90.

Mitchell, N. C., M. Ligi, V. Ferrante, E. Bonatti, and E. Rutter, 2010, Submarine salt flows in the central Red Sea: Geological Society of America Bulletin, v. 122, p. 701-713.

Monnin, C., and C. Ramboz, 1996, The anhydrite saturation index of the ponded brines and sediment pore waters of the Red Sea deeps: Chemical Geology, v. 127, p. 141-159.

Neev, D., and J. K. Hall, 1979, Geophysical investigations in the Dead Sea: Sedimentary Geology, v. 25, p. 209-238.

Neugebauer, I., A. Brauer, M. J. Schwab, N. D. Waldmann, Y. Enzel, H. Kitagawa, A. Torfstein, U. Frank, P. Dulski, A. Agnon, D. Ariztegui, Z. Ben-Avraham, S. L. Goldstein, and M. Stein, 2014, Lithology of the long sediment record recovered by the ICDP Dead Sea Deep Drilling Project (DSDDP): Quaternary Science Reviews, v. 102, p. 149-165.

Pilcher, R. S., and R. D. Blumstein, 2007, Brine volume and salt dissolution rates in Orca Basin, northeast Gulf of Mexico: Bulletin American Association Petroleum Geologists, v. 91, p. 823-833.

Pottorf, R. J., and H. L. Barnes, 1983, Mineralogy, geochemistry, and ore genesis of hydrothermal sediments from the Atlantis II Deep, Red Sea: Economic Geology Monographs, v. 5, p. 198-223.

Ramboz, C., and M. Danis, 1990, Superheating in the Red Sea? The heat-mass balance of the Atlantis II Deep revisited: Earth & Planetary Science Letters, v. 97, p. 190-210.

Shanks III, W. C., and J. L. Bischoff, 1980, Geochemistry, sulfure isotope composition, and accumulation rates of Red Sea geothermal deposits: Economic Geology, v. 75, p. 445-459.

Sheu, D. D., 1987, Sulfur and organic carbon contents in sediment cores from the Tyro and Orca basins: Marine Geology, v. 75, p. 157-164.

Sirota, I., Y. Enzel, and N. G. Lensky, 2017, Temperature seasonality control on modern halite layers in the Dead Sea: In situ observations: Geological Society America Bulletin, v. 129, p. 1181-1194.

Smit, J., J. P. Brun, X. Fort, S. Cloetingh, and Z. Ben-Avraham, 2008, Salt tectonics in pull-apart basins with application to the Dead Sea Basin: Tectonophysics, v. 449, p. 1-16.

Stein, M., A. Starinsky, A. Agnon, A. Katz, M. Raab, B. Spiro, and I. Zak, 2000, The impact of brine-rock interaction during marine evaporite formation on the isotopic Sr record in the oceans: evidence from Mt. Sedom, Israel: Geochimica et Cosmochimica Acta, v. 64, p. 2039-2053.

Torfstein, A., S. L. Goldstein, Y. Kushnir, Y. Enzel, G. Haug, and M. Stein, 2015, Dead Sea drawdown and monsoonal impacts in the Levant during the last interglacial: Earth and Planetary Science Letters, v. 412, p. 235-244.

Trefry, J. H., B. J. Presley, W. L. Keeney-Kennicutt, and R. P. Trocine, 1984, Distribution and chemistry of manganese, iron, and suspended particulates in Orca Basin: Geomarine Letters, v. 4, p. 125-130.

Tribovillard, N., V. Bout-Roumazeilles, T. Sionneau, J. C. M. Serrano, A. Riboulleau, and F. Baudin, 2009, Does a strong pycnocline impact organic-matter preservation and accumulation in an anoxic setting? The case of the Orca Basin, Gulf of Mexico: Comptes Rendus Geoscience, v. 341, p. 1-9.

Wallmann, K., F. S. Aghi, D. Castradori, M. B. Cita, E. Suess, J. Greinert, and D. Rickert, 2002, Sedimentation and formation of secondary minerals in the hypersaline Discovery Basin, eastern Mediterranean: Marine Geology, v. 186, p. 9-28.

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

Williams, D. F., and I. Lerche, 1987, Salt domes, organic-rich source beds and reservoirs in intraslope basins of the Gulf Coast region, in I. Lerche, and J. J. O'Brien, eds., Dynamical geology of salt and related structures: New York, Academic Press, p. 751-830.

Zak, I., and R. Freund, 1981, Asymmetry and basin migration in the Dead Sea rift: Tectonophysics, v. 80, p. 27-38.

Zierenberg, R. A., and W. C. Shanks III, 1983, Mineralogy and geochemistry of epigenetic features in metalliferous sediment, Atlantis II Deep, Red Sea: Economic Geology, v. 78, p. 57-72.

 

Polyhalite, Geological origins of an alternate low-chloride potash fertiliser

John Warren - Tuesday, July 31, 2018

Introduction

Polyhalite is the hydrated sulfate of potassium, calcium and magnesium, with the formula (K2Ca2Mg(SO4)42H2O). Polyhalite crystallises in the triclinic system, but individual euhedral crystals are very rare in nature where the usual habit is fibrous to massive. It is typically colourless, grey to white, although the natural colour in some evaporite deposits tends to brick red due to iron oxide inclusions. Primary (syndepositional) precipitates can be layered at a mm to cm scale. Its Mohs hardness is 3.5 with a specific gravity of 2.8. It was first described from a Salzburg mine in 1818 and the name comes from a Latin root that refers to the “many salts”evident in its chemical formula.

Polyhalite is relatively easy to distinguish from associated evaporites by simple field tests. Its hardness separates it from most evaporite salts, other than anhydrite. It has a bitter taste and is water soluble (incongruent dissolution), with remnants of gypsum and syngenite (K2Ca(SO4)22H2O), which is also soluble, leaving behind a final residue of gypsum (unlike sylvite and halite which dissolve wholly and congruently in fresh water). Polyhalite is not deliquescent, unlike carnallite, and gives a purple flame result when held in a gas flame due to its potassium content, unlike the non-potassium salts. Table 1 lists the constituents of the various evaporite salts mentioned in this article.


The complete equilibrium evaporation of modern seawater at 25ºC produces the mineral sequence: CaCO3 (calcite and aragonite), CaSO4 (gypsum and anhydrite), halite (NaCl), other sulphates (glauberite, polyhalite, epsomite, hexahydrite, and kieserite) and chlorides (carnallite and bischofite). Halite is the dominant mineral because of the high concentration of Na and Cl in seawater (Figure 1; after Harvie et al., 1980).

This mineral sequence with abundant sulphate bitterns reflects the relatively high concentration of sulphate in modern seawater (molar SO4 > Ca), which, following evaporation and precipitation of CaCO3 and CaSO4, produces a SO4-rich, Ca-depleted brine at halite saturation and beyond, from which Mg-sulphate bitterns precipitate. Polyhalite can form syndepositionally at temperatures below 30 °C, via back-reaction of the evaporating K-Mg-SO4 brine with early-formed gypsum or anhydrite (Hardie 1984). Mg and Cl reach high concentrations during the latest bittern stages of evaporation resulting in the precipitation of carnallite and bischofite (Figure 1).

Utilisation

Today polyhalite is mined as a prime ore target in only one place in the world, the Israel Chemical Limited (ICL)-owned Boulby Mine, located below the North Sea, off the North Yorkshire coast of the UK. Currently, ICL-UK distributes around 500 kt/y from its Boulby mine as a direct application fertiliser or bulk-blend/compound NPK additive and is in the process of expanding its polysulphate output as it downgrades its MOP operations. A second major polyhalite mine, located near the Boulby Mine, is proposed by Sirius Minerals PLC. This renewed interest in polyhalite as an economical source of potash fertiliser is why I am writing this article.

Sulphate of Potash (SOP) fertiliser was initially derived from polyhalite/langbeinite in the US during the first half of the twentieth century. But then during the 1940s, after its discovery in vast quantities by prospectors searching for oil in Saskatchewan, the focus for the world’s potash fertiliser supply moved to Canada and the mining of sylvite (muriate of potash- KCl), (Warren, 2016). The geology of SOP was discussed in an earlier Salty Matters article (May 15, 2015). All aspects of SOP and MOP geology and mining are discussed in detail in Chapter 11, Warren (2016).

ICL currently markets crushed and processed polyhalite from the Boulby mine as polysulphate (Table 2). The purity of the Polysulphate product from the Boulby mine is very high (95% polyhalite) with <5% sodium chloride (NaCl) and traces of boron (B) and iron (Fe) at 300 and 100 ppm, respectively (Yermiyahu et al., 2017). The declared minimum analysis of polyhalite for S, K, Mg and Ca is 48% sulphur trioxide (SO3), 14% potassium oxide (K2O), 6% magnesium oxide (MgO) and 17% calcium oxide (CaO), respectively. This compares to a K2O content of contains 60–63% in MOP and around 50% in SOP.

Polyhalite’s make up in terms of K, Mg, S and Cl proportions is similar to the other major potassium-magnesium-sulphate (SOPM) fertilisers: Langbeinite, Schöenite and Patentkali® (Table 2). All are marketed as low-chloride potash fertilisers with additional magnesium and sulphur components. ICL-UK markets polyhalite as a multi-nutrient, low-chloride fertiliser under the brand name Polysulphate®. Sirius plans to market its polyhalite as POLY4®.

CRU estimates the global consumption of potassium-magnesium-sulphate (SOPM) fertilisers in 2017 at 1.7 Mt total product; a comparatively small total compared to the widely traded 65.5 Mt potassium chloride (MOP) market (https://www.crugroup.com "Will polyhalite disrupt the fertiliser industry?” published online April 2018; last accessed 12 July 2018). Polyhalite accounts for around 450-460 kt of current SOPM fertilisers. ICL-UK is currently ramping-up production to 1 Mt/y (140 kt/y K2O) by 2020, as it simultaneously phases out MOP production at the Boulby mine. At current production levels, this will be equivalent to almost 40% of the current SOPM market in K2O terms.

Sirius Minerals’ planned Phase I mine capacity (10 Mt/y product) is on a different scale altogether; around four times larger than the current SOPM market in K2O terms. This volume is almost the same size as the current global potassium sulphate (SOP) market, which is the most popular low-chloride potash fertiliser, outside China. The success of any future expanded SOPM application in agriculture is contentious; the majority of SOPM consumption has traditionally been concentrated close to production sites, nurtured by the local marketing efforts of producers. The proposed worldwide expansion will be tied to increasing acceptance by the agricultural community of polyhalite as an acceptable cheaper substitute for SOP and perhaps MOP.

Polyhalite as a fertiliser.

In the last few years, the use of polyhalite as an agricultural fertiliser has been tested successfully in a number of studies supported by Sirius minerals (Mello et al., 2018; Pavuluri et al., 2017). Polyhalite supplies four nutrients, is less water soluble than the more conventional potassium sources and may conceivably provide a slower release of nutrients. Studies comparing polyhalite to other K and Mg fertilisers have shown that polyhalite is at least as effective as potassium sulphate (K2SO4) as a slow release source of K, and at least as effective as potassium chloride (KCl) plus magnesium sulphate (MgSO4) as a source of K and Mg (Barbarick, 1991).


The possibility of successful use of polyhalite as a fertiliser is illustrated by the positive effects of its application on the growth of a potato crop at Tapira Brazil (Figure 2; Mello et al., 2017). Its impact on tuber starch and tuber dry matter exceeded that of either MOP or SOP applications. Cultivar Asterix at Tapira is mainly used for frying and chip-making (fries) in the food processing industry. High dry matter and starch content improve texture, and lower sugar contents result in less darkening of fries which is desirable. High dry matter percentage enables lower oil absorption while frying, resulting in lower oil usage per unit product. Tuber firmness is essential to handle mechanical stresses that may occur during tuber harvesting, transport, and storage. Crunchiness and hardness are positively related to starch and dry matter contents and specific gravity.

It seems polyhalite products are probably suitable as a low chloride fertiliser replacement of sylvite in some agricultural applications especially in arid acid, infertile soils as found in parts of Israel and other dry growing areas in the Middle East where salinisation due to fertiliser residues is a known problem. Yermiyahu et al., 2017, found the transport and leaching of Ca, Mg, K and S in soil following polyhalite application is lower than following the application of the equivalent sulphate salt fertilisers. The residual effect of polyhalite fertiliser on the subsequently grown crop is higher than the impact from the equivalent sulphate salts, especially regarding Ca, Mg and S. Irrigation management, as determined by the leaching fraction, has a substantial effect on the efficiency of polyhalite as a source of K, Ca, Mg and S for plant nutrition.

Geology of polyhalite

Polyhalite is a common constituent of many ancient evaporite sequences, especially in Permian and Neogene deposits, due to evaporation of Na-K-Mg-Cl-SO4 marine brines. These sulphate bittern assemblages correspond to periods of MgSO4-enriched ocean chemistries (Lowenstein et al., 2003; Demicco et al., 2005).


Modern polyhalite occurrences

The presence of syndepositional polyhalite in the supratidal evaporite flats around the Ojo de Liebre lagoon was first discovered by Holser (1966), who attributed its origin to the diagenesis of gypsum by interstitial marine brines. A large part of this area is now occu­pied by artificial salt ponds. However, some remnants of the an­cient evaporite flats are still accessible, for example, on the southeast coast of the evaporitic complex, where sedimentolog­ical, chemical and isotopic investigations were performed on evaporitic sediments and interstitial solutions (Figure 3a, Pierre, 1983). In May 1979, the evaporitic succession was mainly composed of gypsum; a few centimetres below the surface, polyhalite was present in the form of small nodules that were partially replacing former gypsum crystals (Figure 3b,c). In May 1980, this evaporitic succession was drastically modified, since polyhalite replaced gypsum sediments lying below the water table. This gives an exact timing for the mineral transformation from gypsum to halite of one year to replace 10cm thick interval of gypsum with a 10cm interval of polyhalite, which points to a chemical evolution of the solutions permeating the sediments.

During this one-year period, ionic concentrations of interstitial brine increased from thirteen to 18 times with respect to seawater concentration (Pierre, 1983). SO4 levels of interstitial solutions in the sabkha were higher than in normal marine brine and progressively increased in a landward direction, suggesting gypsum dissolution by groundwater crossflows.

Isotope study suggests both water and aqueous sulphate in the mudflat porewater have a mixed marine and continental origin (Figure 4). Thus, it appears that sulphate ions are provided in part by marine brines, in part by continental waters which have dissolved Pleistocene interstitial gypsum present at depth. The replacement of gypsum by polyhalite requires not only high Mg2+ and K+, but also high SO4 concentrations in the crossflowing solu­tions (Braitsch, 1971).


The polyhalite in Ojo de Liebre mudflats is diagenetic but also penecontemporaneous with the crystallisation of gypsum. However, in brines with temperatures >30°C, polyhalite may also be a primary co-precipitate with halite, as is occurring in recent saltworks near Santa Pola, SE Spain (as observed by B.C. Schreiber), and in cool-zone (cryogenic) salt lakes associated with widespread mirabilite-glauberite such as in Karabogazgol (Andriyasova, 1972).

Polyhalite is also a minor but widespread phase associated with glauberite in the Late Pleistocene-early Holocene sediments of Lop Nur China (Ma et al., 2016). There, the natural lake evaporites are nonmarine assemblages of mirabilite-glauberite-polyhalite-bloedite-gypsum-halite. The evaporitic stages of the lake fill contain massive amounts of glauberite and polyhalite compared to the other salts present. Polyhalite in the upper 40 m of the lake column and its predominance, is indicative of pervasive back-reactions, as is the presence of very minor amounts of carnallite and sylvite in the same section (Ma et al., 2010; Dong et al., 2012).

Ancient occurrences

Most ancient occurrences are interpreted as early diagenetic, formed in shallow brine crossflows as replacement of anhydrite or gypsum. Even so, there is a direct association between higher volumes of polyhalite in marine evaporite basins and times of MgSO4 enrichment of ocean waters.

Neogene polyhalite

Polyhalite is not found as a widespread primary precipitate in rocks of this age even though ocean chemistries are MgSO4-enriched. Instead, polyhalite is typically a minor but extensive early burial replacement of anhydrite or gypsum. The better-documented examples of this type of replacement polyhalite are found in association with gypsum and thenardite/glauberite in various Tertiary lacustrine basins of Spain. For example, below the exploited thenardite beds in the Madrid (Tajo) basin, Spain, the succession in the upper part of the lower Miocene unit is characterised by glauberite layers made up of a mixture of glauberite (45.8 %) and halite (41.7 %), with a minor polyhalite (7.8 %), dolomite (2.1 %), and clay minerals (1.8 %) (Herrero et al 2015).

The Madrid Basin is a large Tertiary intra-cratonic depression that contains some of the largest fossil sodium sulphate and sepiolite deposits in the world. Bedded sodium sulphates (glauberite and thenardite) are restricted to the Lower Saline Unit, where they are associated with anhydrite, halite, magnesite, polyhalite and minor clays. Glauberite and thenardite are thought to have been deposited in the most central part of a permanent saline lake. The accumulation of thenardite might have taken place during a stage of contraction of the lake system at the beginning of the middle Aragonian (middle Miocene).


Polyhalite occurs as a diagenetic saline phase related both to calcium and sodium sulphates occurrences. Both sepiolite and bentonite deposits are widely distributed within peripherally in distal fan and marginal lacustrine sequences in the so-called Intermediate Unit of the Miocene (middle to upper Aragonian). Thick beds of nearly pure sepiolite were deposited in ponds extended at the toes of arkosic alluviums. Sepiolite is also found within calcrete profiles in these environments. Minor amounts of sepiolite are commonly recognised along with palygorskite in open lacustrine areas. On the other hand, Mg-bentonites characteristically occur associated with dolostones and fine micaceous sands in sequences that provide evidence of fluctuations in the lake level. Polyhalite typically occurs as felty and spherulitic aggregates that alternate with centimetre-thick halite layers or millimetre-thick glauberite laminae in the Lower Saline Unit(Figure 5). The polyhalite crystals are always associated with micritic magnesite). In its turn, the felty polyhalite may be related to skeletal glauberite crystals. The hal­ite crystals commonly exhibit chevron-type mor­phologies. The thickness of the individual layers of halite ranges from 1 to 6 cm.

Similar polyhalite proportions are entrained in a number of glauberitic mineral assemblage in gypsiferous Neogene continental basins across the Iberian Peninsula, such as those of the Zaragoza (Salvany et al., 2007) or Lerín gypsum units (Salvany and Ortí, 1994), both occurrences are in the Ebro basin. In all cases, the polyhalite tends to be either massive or more typically a fibrous rim on large glauberite crystals.


Polyhalite also occurs as a minor phase in some potash regions the Messinian evaporites of the Mediterranean. In the mined succession exposed in the Realmonte mine,(southern Sicily) the halite unit is approximately 400 m-thick. From the bottom to the top, it consists of irregular anhydrite and marly mudstone breccia layer up to 2 m thick followed by units A to D (Figure 6; Lugli et al., 1999). Unit A, up to 50 m thick, contains evenly laminated halite with anhydrite nodules and laminae passing upward to massive halite beds with irregular mudstone bed some decimeters thick. Unit B (approximately 100m thick) consists of massive even layers of halite inter-bedded with thin kainite laminae, along with millimeter to centimetre-thick layers dominated by polyhalite spherulites and anhydrite laminae Figure 7; Garcia-Veigas et al., 1995). It may well be that along with kainite, the layers of polyhalite spherulites are primary co-precipitates at the potash bittern stage. The upper part of the succession contains several kainite layers up to 12 m thick. The 70–80 m thick unit C, consists of halite 10 to 20 cm thick layers separated by irregular mud laminae and it too contains minor polyhalite and anhydrite. Unit D, up to 60 m thick, begins with a grey anhydrite-rich mudstone passing to an anhydrite laminate sequence, followed by halite millimetre- to centimetre-thick layers intercalated with anhydrite laminae and decimetre-thick halite beds.


Lugli et al. (1999) proposed that these lithologies, including the early diagenetic polyhalite, reflect the shallowing and the desiccation of the evaporitic basin resulting from a possible combination of factors: (1) uplift of the basin floor by thrust activity, (2) simple evaporitic drawdown and (3) a basin-wide drop of the Mediterranean sea level.

Polyhalite is also common as a potash contributor along with, in the highly deformed bittern series in the Badenian (Middle Miocene) ores of the Carpathian Foredeep Figure 8a). These beds are highly distorted and host former potash mines extracting a kainite-langbeinite ore target (Figure 8b). These potash-entraining salt deposits occur in western Ukraine within two structural terranes: 1) Carpathian Foredeep (rock and potash salt) and (II) Transcarpathian trough (rock salt) (Figure 8a). Deposits differ in the thickness and lithology, depending on the regional tectonic location (Czapowski et al., 2009). In the Ukrainian part of Carpathian Foredeep, three main tectonic zones are distinguished (Figure 8a): (I) outer zone (Bilche-Volytsya Unit), in which the Miocene molasse deposits overlie the Mesozoic platform basement discordantly at a depth of 10-200 m, and in the foredeep they subsided under the overthrust of the Sambir zone and are at depths of 1.2-2.2 km (Bukowski and Czapowski, 2009); Hryniv et al., 2007); (II) central zone (Sambir Unit), in which the Miocene deposits were overthrust some 8-12 km onto the external part of the Foredeep deposits of the external zone occur at depths of 1.0-2.2 km; (III) internal zone (Boryslav-Pokuttya Unit), where Miocene deposits were overthrust atop the Sambir Nappe zone across a distance of some 25 km (Hryniv et al., 2007).


Potash evaporites of the Carpathian Foredeep host an interesting sulphates group that includes about 20 sulphate evaporite minerals. Exploited potash deposits of the foredeep are composed of kainite, langbeinite, kainite–langbeinite, sylvinite, polyhalite and carnallite rocks with layers of rock salt or interbedded clays and rock salt. In the areas of salt-bearing breccia, a polyhalite–anhydrite layer occurs along the contact with the potash salts bed. Halite, langbeinite and kainite dominated targeted ore levels in these potash deposits. Kieserite, polyhalite, anhydrite, sylvite and carnallite were present in smaller but significant quantities. These deposits, once a source of sulphate of potash, are no longer mined.

A study of sulphur isotopic composition of 10 of the sulphate minerals from the Kalush-Holyh and Stebnyk potash deposits shows that only the basal Ca-sulphates (anhydrite) from the Kalush-Holyn potash deposits has d34SCD values typical of Neogene marine evaporites (+21.0‰; Hryniv et al., 2007). Potash minerals related to the ore-associations in the deposits (polyhalite, anhydrite, kainite, langbeinite and kieserite) show d34SCD values from +15.28 ‰ to +17.54‰, while weathering zone minerals (picromerite, leonite, bloedite, syngenite and gypsum) in the Dombrovo Quarry show values ranging from +14.73‰ to +18.22‰ (Table 3).

According to Hyrniv et al. (2007) the recorded depletion of sulphur isotopic composition of the salt minerals in the Ukranian potash deposits (and their weathering zone) was probably caused by one or more factors as follow: 1) bacterial reduction of sulphate, 2) effect of crystallisation and 3) inflow of surface waters containing sulphates enriched in light sulphur isotopes due to pyrite oxidation. Accordingly, the observed sulphur isotopic composition of minerals from these potash deposits demonstrates the depletion of the original marine brines and continual inflow of new (concentrated) seawater and later meteoric access. The preponderance of lighter sulphur isotopic values recorded in the Stebnyk deposit can be explained by a more intensive inflow of surface waters from the Carpathian nappes or by the oxidation of a part of the pyrite hosted in the sediments. Whatever the case, it seems that once again polyhalite is an early diagenetic mineral.


Permian polyhalite

Permian polyhalite deposits are much more impressive in terms of volume and extent, compared to the Neogene, and are exemplified by massive occurrences in the USA and Europe

Permian polyhalite in West Texas and New Mexico

Polyhalite deposits are by far the most abundant, most numerous, and widespread of all potash mineral occurrences in the Delaware Basin of Texas and New Mexico (Jones 1972; Lowenstein, 1988; Harville and Fritz, 1986). However, langbeinite and sylvite are the economically important potash minerals and have been the focus of many studies, rather than polyhalite documentation (Figure 9a). Permian polyhalite in the Delaware Basin occurs both as massive and disseminated deposits in anhydrite and salt beds and less often in clay beds. Typically, massive deposits and all veins and lenses are composed predominantly of polyhalite, in distinctly compact units that have sharp, clear-cut outlines. Disseminated deposits generally are less defined, shapeless bodiesof spherules as cleavage-parallel growths in a host rock, chiefly in halite. Disseminated occurrences are many times more numerous than the massive deposits, but the amount of polyhalite present is minor in comparison with that present in most massive deposits in anhydrite beds.

Massive polyhalite occurrences outline a crude oval-shaped area in the basin, extending over a region about 325 km long and 220 km wide, covering practically the whole southern half of the area between the Pecos River and the eastern limit of salt in the Ochoa Series (Figure 9a). Occurrences range stratigraphically from low in the Tansill Formation (upper part of Guadalupe Series) in the North-western shelf to near the middle of the Rustler Formation in the north-east corner of the Delaware basin (Figure 9b). Polyhalite beds reach their highest number and size in the Salado Formation (Ochoan), where they have a wide distribution over much of the Delaware and Midland basins and adjacent platform and shelf areas (Figure 9). In the Salado Fm., thick clay seams occur as basal strata that underlie massive polyhalite/anhydrite beds (Harville and Fritz, 1986; Lowenstein, 1988). By virtue of the wide extent and number of massive deposits, polyhalite ranks next to halite and anhydrite among the major constituents of the Salado Formation. Sections with layered halite and polyhalite cover areas of 95,000 km2 and 70,000 km2, respectively (Jones 1972).

Massive polyhalite units are typically compact and fine-grained, exhibiting a variety of colours (grey to red) and textures (irregular to layered to laminated and fibrous to equicrystalline prismatic). Significant volumes are replacements of anhydrite beds, and although they may have almost any shape, most tend to be lenticular to sheet-like masses that spread out along the bedding and replace practically the entire section of anhydrite. Polyhalite units in the McNutt Potash zone, east of Carlsbad, have lateral continuities sufficient to act as marker beds, which separate and define layering in the sylvite-langbeinite ore zones (Figure 10).


As a general rule, sheet-like to crudely tabular polyhalite bodies occur in anhydrite layers where stacked polyhalite units are a few centimetres to a metre thick. Deposits that are more irregular in shape occur mostly in thicker beds of anhydrite (>1m). IN most cases the polyhalite is pseudomorphous after growth-aligned subaqueous and nodular gypsum or nodular anhydrite beds (Figure 11).

Practically all the deposits enclose residual strips and irregular remnants of magnesitic anhydrite, which are mottled and streaked with halitic and anhydritic pseudomorphs after gypsum. Commonly polyhalite crystals and multigrain aggregates project into the magnesitic anhydrite remnants either as elongate crystals and veinlike tongues or as aggregates having scalloped margins convex toward anhydrite.


In many places in the Carlsbad district and nearby parts of the north-western shelf, many of the massive polyhalite deposits grade laterally to an anhydritic hartsalz unit with ore grade levels of sylvite. This is the area known as the McNutt Member or the McNutt potash zone (Figures 9a, 10). The change from polyhalite to hartsalz coincides with a shift from unmineralized to sylvinitic rock peppered with sparse grains and veinlets of carnallite and other magnesium-bearing bittern minerals, such as langbeinite and polyhalite.

In 1988, Lowenstein recognised two types of metre-scale depositional cycles (Type I and Type II) within the McNutt Potash Zone (Figure 11). Both cycles record progressive drawdown and concentration of brine in a shallow, marginal marine drawdown basin. "Type I" cycles have a base of carbonate-siliciclastic mudstone, overlain by anhydrite-polyhalite that is pseudomorphous after primary bedded gypsum. This, in turn, is overlain by bedded halite and capped by muddy halite. Lowenstein (1988) concluded the McNutt Zone of the Salado Formation consists entirely of these two types of metre-scale sequences, variably stacked one upon another (Figure 11).

All units are interpreted as mostly marine-brine dominated units precipitated by evaporation of massive volumes of brines fed by marine seepage or periodic overflows of the Permian ocean water. The upper cap to Type II cycles influenced by inflows of continental groundwater (Figure 11).


A basal mudstone grades upsection into anhydrite-polyhalite that is commonly laminated. Laminae are defined by couplets of anhydrite or polyhalite separated by magnesite-rich mud (Figure 12a-c). The most significant feature of the anhydrite/polyhalite interval is the large number of crystal outlines that occur in the anhydrite-polyhalite laminae. These crystals are now composed of anhydrite, polyhalite, halite, or sylvite but are all interpreted as replacement pseudomorphs after primary gypsum because of their close similarity to typical bottom-nucleated subaqueous gypsum crystal habit,s such as "swallow-tail twins" (Figure 11). In some occurrences, the polyhalite is forming early diagenetic spherules in magnesite layers (Figure 13a). Elsewhere polyhalite directly replaces bottom-nucleated subaqueous gypsum or halite (Figures 12a, c, ), while yet elsewhere it grows as spherular clusters in halite that already has pseudomorphed aligned gypsum crystals Figure 12c). In other places, rippled gypsum beds are replaced by polyhalite and anhydrite. Syndepositional brine reflux likely drove replacement of subaqueous gypsum by anhydrite-polyhalite, in a fashion similar to that described by Hovorka (1992) for halite replacing growth-aligned gypsum.


At the microscopic scale, it is evident that polyhalite forms as a replacement (Figure 13). One of the most common modes of occurrence across the Salado Formation is as coalescing spherules growing in relatively undisturbed magnesite layers (Figure 13a). Elsewhere, coarser mm-scale polyhalite prisms have poikilotopically enclosed anhydrite crystals (Figure 13b). Felted fibrous polyhalite also surrounds euhedral halite (Figure 13c) or forms a replacement rim to halite in the langbeinite-sylvite ore layers (Figure 13d).

"Type II" cycles, lacking the basal mudstone and polyhalite/anhydrite beds, occur between Type I cycles and contain additional halite units (with thinly layered polyhalite) overlain transitionally by muddy halite (also with dispersed polyhalite). Complete brining-upward Type I and Type II cycles record a temporal evolution of depositional environment from a shallow saline lake to an ephemeral salt-pan-saline mudflat complex. The uppermost muddy halite unit interpreted as a continental-dominated sequence, sourced by meteoric inflow from surrounding land areas that mixed with variable amounts of seawater, either from residual pore waters or introduced into the Salado Basin by seepage.

Periodic invasions of seawater best explain the vertical stacking of Type I cycles in the Salado basin, perhaps coincident with eustatic sea-level rises (Lowenstein, 1988). The continental-dominated upper parts of Type I and II cycles formed during intervening periods of eustatic sea-level fall and low stand when nonmarine waters exerted more influence on the brine chemistry. According to Lowenstein (1988), the maximum time interval between major marine incursions averages 100,000 years. The layered nature of the polyhalite replacement implies that this occurred in each eustatic cycle, that is, the replacement was an integral part of the eogenetic hydrology and was not a burial diagenetic (mesogenetic) process.


Permian Polyhalite in Poland and Russia

According to Peryt et al., 2005 (and references therein) there are four polyhalite deposits in the Zechstein of northern Poland, and more than ten polyhalite-bearing areas in the adjacent part of Russia (Figure 14). In addition, K-Mg chlorides are found locally both in Poland and Russia. The K-Mg salts originated during the last stages of chloride accumulation within small, actively subsiding isolated salt basins of the salina type, which were probably tectonically controlled.

The paragenetic sequence in one polyhalite (Zdrada) deposit in the Zechstein of Poland was the result of a very early - penecontemporaneous polyhalitisation of anhydrite that had already pseudomorphed gypsum, much as is seen in the Delaware basin (Peryt et al. 1998). There polyhalite formed by altering anhydrite during crossflows of concentrated brines that were also responsible for potash deposition in local salt basins, while the sulphate-rich brines supplied by the dissolution of emergent parts of the sulfate platform (Peryt et al. 1998).


The timing of the polyhalitisation can be inferred from a S-O isotope crossplot (Figure 15; Peryt et al., 1998). The isotopic compositions of sulphate evaporites indicate that marine solutions were the only source of sulphate ions supplied to the Zechstein basin. The more negative oxygen values associated with the polyhalite compared to its anhydrite precursor indicates somewhat warmer solutions that drove the conversion to polyhalite. These solutions were more saline than those driving the initial shallow anhydritisation that replaced platform gypsum by a reaction with refluxing brines.

 

Polyhalite in the Zechstein of the UK

Polyhalite in the Boulby Mine and the proposed York mine both occur within the Permian Fordon Formation, of the 2nd Zechstein cycle (Z2) in northeast England (Figure 16; Table 4; Stewart 1963; Smith et al., 1986; Kemp et al. 2016). Although initially discovered in 1939, the deeper, polyhalite-bearing Fordon (Evaporite) Formation was largely overlooked until recently. ICL-UK operations at the Boulby Mine have largely depleted the sylvinite target in the Boulby Potash Member, so the mine is now transitioning into polyhalite extraction from the Fordon (Evaporite) Formation (Table 4). The historical output from the Boulby Mine was around 1 Mt/yr of refined KCl product and 0.6 Mt of road salt (Kemp et al., 2016). Polyhalite beds in the proposed York (Whitehall) mine are considered to be so high grade that they can be mined and marketed as SOPM fertiliser with no processing except crushing and sizing (Kemp et al., 2016).

Five evaporite cycles (EZ1-EZ5) are developed in the northwestern corner of the main Permian Zechstein basin where it comes onshore in the UK between Teesside and Lincolnshire (Table 2, Figures 16, 17).


The relationship between the evaporite sequence in the main Zechstein basin and its onshore, lateral gradation into shelf and then semi-continental clastic strata was described by Smith aet al., (1986). Potash salts are known from cycles EZ2, EZ3, and EZ4, and Britain’s only potash producer, the Boulby mine, exploits sylvite from the EZ3 Boulby Potash Member. Sylvite-bearing horizons are also known in the EZ2 cycle, but the key potash resource therein is polyhalite, first discovered in 1939 in the E2 oil exploration hole at Eskdale, Whitby (Stewart, 1949). The only known occurrence of potentially economic volumes of polyhalite in the UK is in the EZ2 Fordon (Evaporite) Formation in this area.

Mineral zonation in the Fordon (Evaporite) Formation was first described in detail by Stewart (1949, 1963) from the Eskdale and Fordon boreholes. Polyhalite was described as partly primary, but mostly a replacement of syndepositional anhydrite. Three subcycles were recognised at Fordon. The Lower subcycle was deposited in a basin that still displayed considerable topographic variation from a shallow-water shelf to a deepwater basin (Figure 17). It contains no known potash occurrences. The Middle subcycle, in which the polyhalite occurs, includes a large volume of basin-fill evaporites, chiefly halite, that filled accommodation space and smoothed out the shelf-basin geometry. Consequently, it shows considerable lateral variation in thickness. The Upper subcycle formed in uniformly shallow-water conditions with no clear distinction between shelf and basin. It hosts a persistent sylvite-bearing horizon known as the Gough Seam. Colter and Reed (1980) showed that Stewart’s mineral zones could be projected far beyond the Fordon borehole and were recognisable throughout much of the British section of the North Sea basin (Doornenbal and Stevenson, 2010).

The description of mineral zones at Eskdale and Fordon by Stewart (1949, 1963) relate to boreholes drilled through the shelf and basin, respectively. The precise correlation of the polyhalite- bearing sulfate deposits between these two environments, or zones, remains ambiguous (Kemp et al., 2016). At present, the polyhalite deposit is referred to as the Shelf seam in the Shelf zone, and the Basin seam in the Basin zone, with a Transition zone across the ramp and in its vicinity where great thicknesses of polyhalite and anhydrite occur with varying amounts of early diagenetic, displacive halite. In borehole SM2 there was solid evidence for overlapping Shelf and Basin seams, separated by 82 m of “sulphatic halite”. Both the shelf and basin polyhalite seams are considered to be of mineable thickness and grade in their relevant sectors, averaging over 12 m in thickness for high-grade sections of >85% polyhalite.

Kemp et al. (2016) argue that the polyhalite is almost entirely secondary, resulting from replacement reactions between freshly deposited anhydrite muds on the seabed, with dense, bottom flowing, K-Mg-rich brines. A sylvite-bearing bittern salt horizon is locally present near the top of the Middle subcycle in both the Basin and the Shelf (though less commonly) and is referred to here as the Pasture Beck seam, after the borehole (also known as SM1) where it was first cored and characterised (Figure 16).

Another sylvite-bearing bittern salt horizon is more commonly present near the top of the Upper subcycle in both Basin and Shelf and is referred to here as the Gough seam; described in the SM4 borehole, where it was first cored and characterised as containing relatively high-grade sylvite. It is not clear why this and the Pasture Beck potash seam are so localised and patchy in distribution, but they may result from bittern brine pools of limited area, cut off from each other as the aggrading basin filled up at the end of each subcycle.


At an even more local scale in the polyhalite ore intervals in the Boulby Mine, there are metre-scale domal= structures interpreted as a form of tepee structure (Figures 18; Abbott, 2017). The height of the domal-shaped structures exposed in the mine workings varies between ~0.4 m and 1.5 m (average = 0.9±0.1 m) and widths ranging from ~2.3 m to 10.5 m (average = 5.3±0.5 m).

Unlike the highly deformed halokinetic flow textures in the overlying sylvite of the Boulby Potash Member, it seems much of the polyhalite ore preserves mostly syndepositional diagenetic alteration structures. Most of the domal features do not show the overthrust brittle ridge crests that define most tepees (Kendall and Warren, 1987). Instead, the domal peak tends to be a fractured and folded local anticlinal culmination. Whether one calls these anticlinal deformations domes in the polyhalite a true tepee, depends on which definition of a tepee one chooses to use. The domal features are thought to be a soft sediment deformation features, formed via polyhalite dewatering, coupled with penecontemporaneous precipitation of halite in opening fractures and below anticline crests in shallow burial. Deformation was driven by fluid crossflows and escapes, as anhydrite converted to polyhalite.

Forming polyhalite?

Nowhere is the present or the past is there evidence of direct primary precipitation of polyhalite. By primary, I mean that to be considered a primary polyhalite, the crystals should drop out of a concentrating at-surface brine either as bottom-nucleated or foundered brine-surface crystals. Such primary textures are widespread in gypsum and halite units but not in polyhalite. Instead, polyhalite textures and isotopic signals indicate polyhalite forms via replacement of gypsum or anhydritised gypsum.

In the modern salt flats of Ojo de Liebre, we see polyhalite replacing gypsum. Likewise, in various Tertiary lacustrine basins in Spain, polyhalite is found in association with gypsum and thenardite/, and it is replacing a CaSO4 phase. In the Badenian marine evaporites of the Carpathian foredeep, the polyhalite is part of the kainite-langbeinite ore sequence. It is in the Permian of the UK and West Texas and New Mexico that the volumes of polyhalite become sufficient for it to become a potential ore target in its own right. Once again all the textural and isotopic evidence indicates polyhalitisation of anhydrite rather than primary precipitation. But this replacement is more likely to be eogenetic (driven by nearsurface hydrologies that were active in the depositional setting) rather than mesogenetic (burial).

The most likely driven mechanism was brine reflux moving highly saline seawater through shallowly buried units of platform or basinal gypsum and anhydrite. This shallow subsurface emplacement occurred while the gypsum anhydrite was still permeable, and so allowed the preservation of pristine texture (pseudomorphs) of the CaSO4 precursors.

Polyhalitisation of basinal and platform gypsum units in the mega-sulphate stages of a saline giant are driven by time separate hydrologies, tied to the changing brine levels in the drawndown basin (Figure 18; Warren, 2016). Marine-derived brine reflux through basinal anhydrites occurs during maximum drawdown in the mega-sulphate basin (blue arrow positions stage b in Figure 18), while reflux through the upper (marginal saltern) parts of a sulphate platform is a response to a relative highstand (blue arrow positions in stage c Figure 18; Warren, 2016 - Chapter 5). The likely loss of permeability as one goes deeper in a sulphate platform and the associated lessening in the volume brine crossflow probably explains why there is an interval of sub-economic polyhalitic sulphate separating the basinal from the shelfal ore zones. The sequence stratigraphic fill model also explains why the patchy potash intervals are located higher in the stratigraphy at the "fill and spill stage" of a hyperarid climate (stage e in Figure 16; Warren 2016).


A drawdown model encompassing two stages of polyhalitisation explains why much of the textures seen in the platform and basin polyhalite units contains evidence of both lamination and subaqueous shallow water deposition (Figure 19). In any megasulphate saline giant, the basin brine level can oscillate between shallow and deep and, depending on the nature of the overlying brine column, and we can deposit primary textures that are mm-cm laminates or upward-aligned gypsum growths, or displacive nodules. Bottom nucleated, upward aligned gypsum crystals indicate relatively stable and saturated bottom chemistries beneath a holomictic brine column (Chapters 1 and 2, Warren 2016). Without holomixis, brine reflux cannot occur. Layered and laminated gypsum sediments interlayered by carbonates indicate subaqueous deposition with fluctuating chemistries in the overlying column Laminites can form via changes in water chemistry in a meromictic deep water mass (as in the modern Dead Sea prior to 1979), or it can indicate a shallow overlying water mass subject to periodic freshening as in the salinas of southern Australia. If the layer and laminate gypsum.anhdrite is interlayered with units of bottom-aligned gypsum or its anhydritised "ghosts," as in west Texas and Poland, then the depositing waters in both units were shallow.


We can now take this reflux model for polyhalitisation and explain why the two polyhalite ore seams in the Forden Evaporite Formation are separated by a low quality "sulphatic halite-anhydrite" unit (Figure 20). At time 1 the basin is at its maximum lowstand and dense reflux brines are sinking into the basinal gypsum units. Water depths below the holomictic brine mass in the basin lows are relatively shallow. At time two the brine levels in the basin are much higher, and a gypsum platform is prograding into the basin. Water depths above the platform are shallow, while they are deep in the basin centre. When the water column is holomictic, brine reflux is occurring across the platform and out into the basin. However descending brines cannot penetrate into all parts of the platform due to compaction and earlier reflux of halite- saturated cements. Brines must pass beyond halite saturation to reach polyhalite (Figure 1). This early loss of permeability created a core of less altered anhydrite below the polyhalite replacement interval.

But we must now ask, why did polyhalitisation of large parts of sulphate platforms reach its zenith in the Permian. A pseudomorphing process with a halite-gypsum focus is seen throughout the rock record (Chapter 5, 7; Warren, 2016). But the volumes of polyhalite we see in the Permian saline giants are different to the much smaller volumes of the Neogen, which is also a time of MgSO4-enrieched waters. Polyhalite is never present in the gypsiferous units of the Messinian or Badenian saline giants in the same volumes we see in the Permian. Then again, the extreme hyperarid hydrologies we see in arid climate belts across the Pangean supercontinent are also unusual. But the seawater chemistry was not too different to that of today (Lowenstein et al., 2005).

References

Abbott, S., 2016, Depositional architecture and facies variability in anhydrite and polyhalite sequences: a multi-scale study of the Jurassic (Weald Basin, Brightling Mine) and Permian (Zechstein Basin, Boulby Mine) of the UK: Doctoral thesis, Imperial College London.

Andriyasova, G. M., 1972, Polyhalite in Kara Bogaz: Khim. Geol. Nauk, v. 3, p. 45-49.

Barbarick, K. A., 1994, Polyhalite application to sorghum-sudangrass and leaching in soil columns: Soil Science, v. 151, p. 159-166.

Bates, R. L., 1969, Potash Minerals: Geology of the industrial rocks and minerals: New York, Dover Publ., 370-385 and 439-440 p.

Braitsch, O., 1971, Salt Deposits: Their Origin and Compositions: New York, Springer-Verlag, 297 p.

Bukowski, K., G. Czapowski, S. Karoli, and M. Babel, 2007, Sedimentology and geochemistry of the Middle Miocene (Badenian) salt-bearing succession from East Slovakian Basin (Zbudza Formation): Geological Society, London, Special Publications, v. 285, p. 247-264.

Colter, V. S., and G. E. Reed, 1980, Zechstein 2 Fordon Evaporites of the Atwick No. 1 borehole, surrounding areas of N.E. England and the adjacent southern North Sea, in H. Fuchtbauer, and T. M. Peryt, eds., The Zechstein Basin with Emphasis on Carbonates: Stuttgart, E Schweizerbarts scheverlagsbuchhandlung, p. 115-129.

Czapowski, G., K. Bukowski, and K. Poborska-Młynarska, 2009, Miocene rock and potash salts of West Ukraine: Field geological-mining seminar of the Polish Salt Mining Society. Geologia (Przegląd Solny 2009), Wyd. AGH, Kraków, 35, 3: 479-490. (In Polish, English summary).

Decima, A., and F. Wezel, 1973, Late Miocene evaporites of the central Sicilian Basin; Italy: Initial reports of the Deep Sea Drilling Project, v. 13, p. 1234-1240.

Decima, A., and F. C. Wezel, 1971, Osservazioni sulle evaporiti messiniane della Sicilia centromeridionale: Rivista Mineraria Siciliana, v. 130–132, p. 172–187.

Demicco, R. V., T. K. Lowenstein, L. A. Hardie, and R. J. Spencer, 2005, Model of seawater composition for the Phanerozoic: Geology, v. 33, p. 877-880.

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

Doornenbal, H., and A. Stevenson, 2010, Petroleum geological atlas of the southern Permian Basin Area: Houten, Netherlands, EAGE Publications, 342 p.

Garcia-Veigas, J., F. Orti, L. Rosell, C. Ayora, R. J. M., and S. Lugli, 1995, The Messinian salt of the Mediterranean: geochemical study of the salt from the central Sicily Basin and comparison with the Lorca Basin (Spain): Bulletin de la Societe Geologique de France, v. 166, p. 699-710.

Griswold, G. B., 1982, Geologic overview of the Carlsbad potash-mining district: Circular New Mexico Bureau of Mines and Mineral Resources, v. 182, p. 17-22.

Harvie, C. E., J. H. Weare, L. A. Hardie, and H. P. Eugster, 1980, Evaporation of sea water; calculated mineral sequences: Science, v. 208, p. 498-500.

Harville, D. G., and S. J. Fritz, 1986, Modes of diagenesis responsible for observed successions of potash evaporites in the Salado Formation, Delaware Basin, New Mexico: Journal Sedimentary Petrology, v. 56, p. 648-656.

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

Holser, W. T., 1966, Diagenetic polyhalite in recent salt from Baja California: American Mineralogist, v. 51, p. 99-109.

Hovorka, S. D., 1992, Halite pseudomorphs after gypsum in bedded anhydrite; clue to gypsum-anhydrite relationships: Journal of Sedimentary Petrology, v. 62, p. 1098-1111.

Hryniv, S., J. Parafiniuk, and T. M. Peryt, 2007, Sulphur isotopic composition of K Mg sulphates of the Miocene evaporites of the Carpathian Foredeep, Ukraine: Geological Society, London, Special Publications, v. 285, p. 265-273.

Jones, C. L., 1972, Permian Basin potash deposits, south-western United States, in G. Richter-Bernberg, ed., Geology of Saline Deposits, v. 7: Paris, UNESCO Earth Science Series, p. 191-201.

Kemp, S. J., F. W. Smith, D. Wagner, I. Mounteney, C. P. Bell, C. J. Milne, C. J. B. Gowing, and C. J. B. Pottas, 2016, An Improved Approach to Characterize Potash-Bearing Evaporite Deposits, Evidenced in North Yorkshire, United Kingdom: Economic Geology, v. 111, p. 719-742.

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

Koriń, S. S., 1994, Geological outline of Miocene salt-bearing formations of the Ukrainian fore-Carpathian area (In Polish, English summary): Przegląd Geologiczny, v. 42, p. 744-747.

Lowenstein, T. K., 1988, Origin of depositional cycles in a Permian ''saline giant''; the Salado (McNutt Zone) evaporites of New Mexico and Texas: Geological Society of America Bulletin, v. 100, p. 592-608.

Lowenstein, T. K., L. A. Hardie, M. N. Timofeeff, and R. V. Demicco, 2003, Secular variation in seawater chemistry and the origin of calcium chloride basinal brines: Geology, v. 31, p. 857-860.

Lugli, S., 1999, Geology of the Realmonte salt deposit, a desiccated Messinian Basin (Agrigento, Sicily): Memorie della Societá Geologica Italiana, v. 54, p. 75-81.

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

Ma, L., Q. Tang, B. Li, Y. Hu, and W. Shang, 2016, Sediment characteristics and mineralogy of salt mounds linked to underground spring activity in the Lop Nor playa, Western China: Chemie der Erde - Geochemistry, v. 77, p. 383-390.

Mello, S. d. C., F. J. Pierce, R. Tonhati, G. S. Almeida, D. D. Neto, and K. Pavuluri, 2018, Potato Response to Polyhalite as a Potassium Source Fertilizer in Brazil: Yield and Quality: HortScience, v. 53, p. 373-379.

Ordóñez, S., J. P. Calvo, M. A. García del Cura, A. M. Alonso-Zarza, and M. Hoyos, 1991, Sedimentology of Sodium Sulphate Deposits and Special Clays from the Tertiary Madrid Basin (Spain): Lacustrine Facies Analysis, Blackwell Publishing Ltd., 39-55 p.

Pavuluri, K., Z. Malley, M. K. Mzimbiri, T. D. Lewis, and R. Meakin, 2017, Evaluation of polyhalite in comparison to muriate of potash for corn grain yield in the Southern Highlands of Tanzania: African Journal of Agronomy, v. 5, p. 325-332.

Peryt, T. M., C. Pierre, and S. P. Gryniv, 1998, Origin of polyhalite deposits in the Zechstein (Upper Permian) Zdrada Platform (northern Poland): Sedimentology, v. 45, p. 565-578.

Peryt, T. M., H. Tomassi-Morawiec, G. Czapowski, S. P. Hryniv, J. J. Pueyo, C. J. Eastoe, and S. Vovnyuk, 2005, Polyhalite occurrence in the Werra (Zechstein, Upper Permian) Peribaltic Basin of Poland and Russia: evaporite facies constraints: Carbonates and Evaporites, v. 20, p. 182-194.

Pierre, C., 1983, Polyhalite replacement after gypsum at Ojo de Liebre Lagoon (Baja California, Mexico); an early diagenesis by mixing of marine brines and continental waters: Sixth international symposium on salt, v. 1, p. 257-265.

Salvany, J. M., J. Garcia-Veigas, and F. Orti, 2007, Glauberite-halite association of the Zaragoza Gypsum Formation (Lower Miocene, Ebro Basin, NE Spain): Sedimentology, v. 54, p. 443-467.

Salvany, J. M., and F. Orti, 1994, Miocene glauberite deposits of Alcanadre, Ebro Basin, Spain: sedimentary and diagenetic processes, in R. W. Renaut, and W. M. Last, eds., Sedimentology and geochemistry of modern and ancient saline lakes, v. 50, SEPM/Society for Sedimentary Geology Special Publication, p. 203-215.

Schaller, W. T., and E. P. Henderson, 1932, Mineralogy of drill cores from the potash field of New Mexico and Texas, Bull. U.S. Geol. Surv., No. 833, 1-124, 39 pis., Washington, D.C., p. 171.

Smith, D. B., 1974, The origin of the Permian Middle and Upper Potash deposits of Yorkshire; an alternative hypothesis [with discussion]: Yorkshire Geol. Soc., Proc., v. 39.

Smith, D. B., G. M. Harwood, J. Pattison, and T. H. Pettigrew, 1986, A revised nomenclature for Upper Permian strata in eastern England, in G. M. Harwood, and D. B. Smith, eds., The English Zechstein and related topics, Geological Society of London Special Publication no 22.

Smith, F. W., J. P. L. Dearlove, S. J. Kemp, C. P. Bell, C. J. Milne, and T. L. Pottas, 2014, Potash – Recent exploration developments in North Yorkshire, in E. Hnger, T. J. Brown, and G. Lucas, eds., Proceedings of the 17th Extractive Industry Geology Conference, EIG Conferences Ltd.( 202pp), p. 45-50.

Stewart, F. H., 1949, The petrology of the evaporites of the Eskdale no. 2 boring, east Yorkshire; part 1, The lower evaporite bed: Miner. Mag., v. 28, p. 621-675.

Stewart, F. H., 1963, The Permian lower evaporites of Fordon in Yorkshire: Proceedings of the Yorkshire Geological Society,, v. 34, p. 1-44.

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

 

Stable isotopes in evaporite systems: Part III - 18O (Oxygen)

John Warren - Sunday, July 01, 2018

 

Introduction

Oxygen isotope determinations in evaporitic sediments are typically based on: 1) using oxygen held in the water molecule itself (H2O); 2) oxygen in the carbonate anion held in evaporitic dolomites or limestones or; 3) in sulphate from evaporitically precipitated gypsum or anhydrite. Oxygen measures on the water molecule can be co-associated with deuterium (D - heavy hydrogen) determinations. So isotopic sampling of evaporitic limestone and dolomites means carbon isotope values can be co-determined from the same mineral phase (CO3 source). Likewise, with the calcium sulphates, the sulphur isotope is always available for co-study (SO4 source in gypsum or anhydrite).

We have already discussed sulphur and carbon isotope variations in evaporitic settings in the previous two articles (30 April 2018 and 31 May 2018, respectively). So, in this article, we shall look at how oxygen isotope values vary with the co-associated deuterium, carbon and sulphur isotope phases. We focus on three sources for isotope samples (water molecules in a brine, evaporitic carbonate minerals, calcium sulphate minerals) and show that when oxygen values are co-plotted against deuterium, carbon or sulphur isotope values, it becomes a handy tool in defining depositional and diagenetic evolution in a range of evaporitic settings.


Oxygen isotope fractionation in water molecules in evaporating brines

The stable isotope community has long known of the potentially extreme effects of evaporation on the isotopic composition of liquids and the residual enrichment of the heavier isotope in the remaining brine. After all, Urey himself applied this knowledge when he demonstrated the existence of deuterium through evaporative enrichment of liquid hydrogen (Urey et al., 1932). Enrichment in heavier isotopes in the residual brine is documented in settings as diverse as evaporating Dead Sea brines (Gat, 1984) and degassing epithermal systems (Zheng, 1990).

As any water (brine) evaporates there is a commensurate preferential escape of the lighter 16O water molecules, this leaves behind an increasing proportion of heavier water molecules containing 18O. Hence, with increasing degrees of evaporation the δ18O signature in the remaining water mass becomes increasingly positive (Figure 1). Co-variance of deuterium with increasing oxygen isotope values in a concentring brine is a long-established observation (Figure 2; Cappa et al., 2003), and defines a type of Raleigh fractionation or distillation.


There is another factor involved in the degree of enrichment of the heavier isotopes of oxygen or deuterium, and that is the humidity of the air above the evaporating brine. Humidity controls the extent of evaporative concentration, and there is a differential level of isotope enrichment in the residual brine tied to changing humidity (Figure 2). It is a response to the lowering of the evaporation rate with increasing humidity. The humidity effect in evaporative settings is documented both experimentally and in natural settings such as modern sabkhas and salinas (Chapter 2 in Warren 2016, for a summary of literature). As a general rule, the lower the humidity, the greater the degree of enrichment of the heavier isotope. Horton et al., (2016) show that δ18OSMOW values of saline lake waters from are often shifted by >+10‰ relative to source waters discharging into the lake (Figure 3, especially 3c).


Up until February 1979, the Dead Sea was a permanently stratified hypersaline water body (see Warren 2016, Chapter 4 for hydrological and sedimentological details). Both the upper and lower water masses were moderately enriched in δ18OSMOW (Figure 4: Gat 1984). After the overturn and mixing the surface waters, the degree of enrichment in δ18O in the surface waters constitutes a balance between the dilution by freshwater influx and the isotope fractionation (enrichment) which accompanies evaporative water loss and vapour exchange with the atmospheric moisture. Gat's modelling of the seasonal cycle and long-term trends of δ18OSMOW in response to the changes in the environmental parameters, shows that the dominant control on isotope enrichment in the surface waters, post overturn, is exercised by the salinity of the surface waters, through its effect on the vapour pressure gradient between the lake's surface and the atmosphere. Interestingly, before the overturn event the upper water mass was more homogenous in terms of salinity and isotope enrichment and its enriched isotope values mostly tracked those of the much more stable and somewhat more saline lower water mass.


Deuterium-oxygen isotope plots of water molecules can also be useful in studying the origin of hydrated salts such as gypsum, but only if there has been minimal postdepositional alternation of the primary precipitate. A classic paper focusing on the composition of structural water held in the gypsum lattice of Messinian (Late Miocene) evaporites of Sicily was published by Bellanca et al., 1986. In Sicily, there are two main types of texture in gypsum-dominated outcrops in the Messinian sub-basins of Sicily (laminated and massive). The laminar gypsum, locally known as balatino, is a shallow-water saltern deposit, the other is a massive form of gypsum typically interpreted as a diagenetic replacement of either primary gypsum of anhydrite.

The different isotopic compositions of hydration water in the two gypsum lithotypes are shown in Figure 6. Laminar gypsum shows a predominance of positive values for both oxygen (range-1.59‰ < δ18OSMOW < +6.02‰) and deuterium (range -7.3‰ < δD < +22.7‰), while both oxygen and deuterium ranges in the massive gypsum are negative (-4.21‰ < δ18OSMOW > -2.23‰; -40.9‰ < δD < -34.4‰).


In Figure 5 the majority of points representative of the laminar gypsum mother waters fall to the right of the meteoric water line of Craig ( 1961) and lie on a path characterised by a positive slope (δD = 3.97δ18OSMOW - 0.59) and includes the SMOW point. Such a distribution is consistent with an origin of the gypsum by direct pre­cipitation from an evaporating solution saturated with respect to gypsum and is close to those of mother waters in recent gypsum samples precipitated in Mediterranean salinas and, there­fore, suggest that the solutions from which the laminar gypsum precipitated were marine waters concentrated by evaporation. A few other examples show δ18O and δD values shifted towards negative values, which indicate stages of dilution with large masses of continental waters poured into the deposition basin during the crystallisation of gyp­sum (Bellanca et al., 1986).

In contrast, the waters from Massive Gypsum plot along a line with a negative slope (δD = -2.66 δ18OSMOW -46.73). Clearly, these structural waters have a different origin. Bellanca (op. cit) argues these distinctive signatures are indicative of rehydration from anhydrite; others argue massive gypsum is a result of subsurface recrystallisation of primary gypsum without an intervening anhydrite stage (see Testa and Lugli (2000) for the detailed discussion of this topic)

Carbon and oxygen isotope co-variations in evaporitic carbonates

The isotopic makeup of residual water molecules evolving into a brine is not the only phase affected by the chemical consequences of evaporation (Horton et al., 2016; Warren 2016). As any natural water evaporates, its chemistry changes, as concentrating dissolved phases and increasing alkalinity force changes in equilibrium conditions. One of the most obvious consequences of evaporation is the formation of sedimentary evaporites, including brine pool carbonates (e.g. calcite, aragonite, dolomite, trona). The coupled δ18O and δ13C enrichment during evaporation, and the precipitation of endogenic Holocene carbonates is documented and discussed at some length in a number of review papers (Horton et al., 2016; Pierre, 1988).


Horton et al. (2016) document a general tendency for calcites precipitated in lakes located in somewhat less humid climates to show enrichment in the heavier isotope. The observed average lake carbonate δ18OPDB values from the 57 lakes plotted in Figure 6 are more positive than the modelled summer month meteoric water derived calcite δ18O values (Horton et al., 2016). Lake calcites precipitating in humid environments generally plot closer to the 1:1 line, suggesting lakes in these environments are less impacted by evaporative modification. Yet, 46 of the 57 lake records analysed (i.e. 81%) plot to the right of the 1:1 line consistent with evaporative modification of lake water δ18O. Forty-two percent of the lake carbonate δ18O records are >5‰ shifted towards more positive δ18O values than would be expected for summer-month carbonate precipitates derived from unmodified local meteoric water. Although many lakes with vastly different modern aridity index values show similar offsets between modelled and observed δ18O, lakes from currently arid and semi-arid environments have a much larger average δ18O offset (5.4‰) than sub-humid and humid environment lakes (2.0‰).


The dolomite forming lakes of the Coorong region show a similar set of enrichment in both oxygen and carbon isotopes within that type of Holocene dolomite precipitating directly from evaporating surface brines (dolomite Type-A; Rosen et al., 1989; Warren 1990, 2000). The other type of Holocene dolomite in the Coorong lakes (dolomite-B) shows no noticeable C-O covariant trend related to Raleigh distillation (Figure 7a). Type-A dolomite has a heavier oxygen isotope signature than type-B and is 3 - 6‰ heavier in 13C (Figure 7a). Type-A dolomite also has distinct unit cell dimensions (Rosen et al., 1989).

Type A tends to be magnesium-rich with up to 3-mole percent excess MgCO3, while type-B is near stoichiometric or calcian-rich. Type-A dolomite typically occurs in association with magnesite and hydromagnesite, Type B with Mg-calcite. Transmission electron microscopy (TEM) shows that Type A dolomites have a heterogeneous microstructure due to closely spaced random defects, while type B dolomites exhibit a more homogeneous microstructure implying excess calcium ions are more evenly distributed throughout the lattice. TEM studies show that the two types of Coorong dolomite are distinct and are not intermixed with other mineral phases; they are primary precipitates, and not replacements and are not transitional (Miser et al., 1987).

Within the lake stratigraphy the dolomites occupy two distinct positions, Type A dolomites occur as surficial 'yoghurt' textured gels that in each water-filled winter season are washed and blown across the lake surface. By late spring and through summer these surface waters have dried up (summer salinities ≈ 120‰), and the lake sediment surface is a mud-cracked interval of massive carbonate (Warren, 1990; 2016). Type B dolomites occur in the laminated unit that underlies the laminated with signatures implying precipitation from waters with bicarbonates, perhaps showing a stronger strong input from organic materials and are especially prevalent in the more marginward part of the laminated fille where meteoric groundwaters are continually flowing into the edges of the lakes and mixing with lake pore brines.

Figure 7b places these two Coorong dolomites in the context of other areas of primary dolomite accumulations within Holocene carbonate depositional settings. Today sulphate-reducing bacteria or archeal methanogens have been called upon to explain the primary precipitation of dolomite in bacterial biofilms in almost all these other settings. It is not my intention to question the importance of bacterial metabolism in these other dolomite-accumulating settings, only to point out the bicarbonate from which the Coorong type A dolomites have precipitated show a positive and co-variant enrichment in both carbon and oxygen valued that are more typical of evaporative concentration. Evaporative enrichment in carbon values tied CO2 degassing in highly saline waters was documented in the Dead Sea by Stiller et al., 1985 and discussed in last month's article (31 May 2018).

Evaporitic carbonates especially when interbedded with calcium sulphate beds can also dissolve and alter (Warren, 2016; Chapter 7). Evaporite-derived dedolomites are often associated with evaporite dissolution breccias, which indicates the stratigraphic position of the now dissolved calcium sulphate bed that supplied the excess calcium needed to dedolomitise (Lee, 1994; Fu et al., 2008). Dedolomite under this scenario forms via the reaction of calcium sulphate-rich solutions with pre-existing dolomite to produce calcite with magnesium sulphate as a possible byproduct. The latter is rarely preserved, as it is highly soluble, and either remains as dissolved ions in the escaping waters or is quickly redissolved and flushed by through-flowing groundwaters (Shearman et al., 1961). The CaSO4 dissolution process is often driven by meteoric flushing of nearsurface oxidising waters and former ferroan dolomites are preferentially replaced. The resulting calcitised dolomites are outlined by intervals stained red with iron oxides and hydroxides.


With uplift-related (telogenetic) dedolomites the distribution and isotopic composition of dedolomite can reflect variations in the regional hydrology. This can be seen in the dedolomites of the Lower Cretaceous Edwards Group in the Balcones fault zone area of south-central Texas (Ellis, 1985, 1986). The Edwards Group consists of 120-180 metres of porous limestone and dolomite that accumulated on the Comanche shelf in shallow-water subtidal, intertidal, and supratidal marine environments. During early burial diagenesis, carbonate mud neomorphosed to calcitic micrite, aragonite and Mg-calcitic allochems were altered to calcite or were leached, and evaporites formed in tidal-flat sediments. Each of these phases had a characteristic stable isotope signature (Figure 8). Dolomite is widespread and formed in environments ranging from hypersaline to fresh-water as shown by the two isotope clusters in the Edwards dolomite (meteoric versus evaporitic reflux).

Late Tertiary faulting along the Balcones fault zone, tied to Jurassic salt withdrawal, initiated a circulating, fresh-water aquifer system to the west and north of a fairly distinct “bad-water line,” which roughly parallels the Balcones fault zone. To the south of the bad-water line, interstitial fluids remained relatively stagnant and contain over 1000 mg/l dissolved solids. Because of the differences in the chemistry of the interstitial fluids, post-faulting diagenesis in the two zones has been very different.

Water in the bad-water zone can be saturated with respect to calcite, dolomite, gypsum, celestite, strontianite, and fluorite, whereas water in the fresh-water zone is saturated only with respect to calcite. Due to the change in water chemistry, rocks in the fresh-water zone have been extensively recrystallised to coarse microspar and pseudospar, extensive dedolomitization has occurred, and late sparry calcite cements have precipitated. This creates a suite of covariant isotope trends and clusters with the dedolomite showing a distinctive set of carbon and oxygen values relate to soil water influences indicated by calcites with more negative carbon values (Figure 8 indicated by brown shading). In contrast, rocks in the bad-water zone retain fabrics associated with pre-Miocene diagenesis, and there is little or no evidence of widespread dedolomite, indicated by pink shading in Figure 8.

The importance of meteoric diagenesis in the formation of dedolomite in shallow, subsurface telogenetic environments is illustrated by the fact that the Edwards Group had a stable mineralogy of calcite and dolomite before the circulation of fresh water began and drove the precipitation of meteoric spar, microspar and dedolomite. Isotopic values for the dedolomites follow a similar trend to those of the microspars and pseudospars. As with the microspars and pseudospars formed by the entry of telogenetic water, it can be shown that dedolomites are in isotopic equilibrium with Edwards water on a regional scale, which supports the contention that the dedolomites are still forming from crossflows of present-day formation-water (Ellis, 1985).

Sulphur and oxygen relationships in calcium sulphate

Modern seawater sulphate has a homogeneous and well-defined isotopic composition for both sulphur and oxygen:

34SSO4 = +20 ± 0.5‰ CDT

18OSO4 = +9.5 ± 0.5‰ SMOW

Likewise, the fractionation of sulphur and oxygen, which occurs during the transition from aqueous to the solid state of sulphate is also near constant at earth surface temperatures. For gypsum, the mean values of the isotope enrichment factor are (Pierre, 1988):

δ34Sgypsum—SO4 = 1.65‰

and,

δ18Ogypsum—SO4 = 3.5‰

Thus the δ34S and δ18O values of sulphate evaporites are directly related to the state of the aqueous sulphate reservoir wherever precipitation occurred. A plot of ancient marine CaSO4 evaporites shows the sulphur curve for seawater has varied across time from +30‰ in the Cambrian, to around +10‰ in the Permian and that it increased irregularly into the Mesozoic to its present value of +20‰. Oceanic oxygen isotope values show much less variability. Sulphur is largely resistant to isotopic fractionation during the increasing temperatures associated with burial alteration and transformation (Worden et al., 1997). All of these aspects are discussed in detail in the April 30, 2018 article.

With this knowledge of the relative lack of fractional in the subsurface compared to the much greater susceptibility of oxygen isotopes in the mesogenetic and telogenetic realms let us now look in more detail at the significance of oxygen variation in a variety of sulphate entraining settings.


Isotopically, the effects of dissolution and brine recycling in fracture-filling fibrous gypsum cements of various ages emplaced in a formation's burial evolution can be used define the sequential development of the superimposed diagenetic textures in the original gypsum unit (Figures 9, 10; Moragas et al., 2013). The upper Burdigalian Vilobí Gypsum Unit, located in the Vallès Penedès half-graben (NE Spain) and consists of a 60-m thick succession of laminated-to-banded primary and secondary gypsum. The unit is variably affected by Neogene extension in the western part of Mediterranean Sea. Tertiary extensional events are recorded in the evaporitic gypsum unit as six fracture sets and fills (faults and joints - S1 - S5), which can be linked with basin-scale deformation stages.

Combined structural, petrological and isotopic study of the unit by Moragas et al. (2013) established a chronology of fracture formation and infilling, from oldest to youngest as: (i) S1 and S2 normal faults sets with formation and precipitation of sigmoidal gypsum fibres; (ii) S3 joint sets with perpendicular fibres; (iii) S4 inverse fault sets, infilled by oblique gypsum fibres and associated with thrust-driven deformation of the previous fillings; and (iv) S5 and S6 joint sets tied to later dissolution processes and infilled by macrocrystalline gypsum cements likely related to the telogenetic realm. The fractures provided ongoing pathways for focused fluid circulation within the Vilobí Unit. The oxygen, sulphur and strontium isotope compositions of the original host rock and the various precipitates in the fractures imply ongoing convective recycling processes across the host-sulphates to the fracture infillings, as recorded by a general enrichment trend toward heavier S–O isotopes, from the oldest precipitates (sigmoidal fibres) to the youngest (macrocrystalline cements). The marine strontium signal is mostly preserved in the various postdepositional infillings, unlike the oxygen and to a lesser extent the sulphur isotope signals, which are evolving with the origin and temperature of the waters flowing in the fracture sets (Figure 10).


In any ancient silicified anhydrite nodule or bedded silicified succession, not all silica-replacing anhydrite in a particular region need come from the same source or be emplaced by the same set of processes. Silicified nodules within middle-upper Campanian (Cretaceous) carbonate sediments from the Laño and Tubilla del Agua sections of the Basque-Cantabrian Basin, northern Spain preserve cauliflower morphologies, together with anhydrite laths enclosed in megaquartz crystals and spherulitic fibrous quartz (quartzine-lutecite). All this shows that they formed by ongoing silica replacement of nodular anhydrite (Figures 10, 11; Gómez-Alday et al., 2002).

Anhydrite nodules at Laño were produced by the percolation of saline marine brines, during a period corresponding to a depositional hiatus. They have δ34S and δ180 mean values of +18.8‰ and +13.6‰, respectively, both consistent with Upper Cretaceous seawater sulphate values. Higher δ34S and δ180 (mean values of + 21.2‰ and 21.8‰ characterise nodules in the Tubilla del Agua section and are interpreted as indicating a partial bacterial sulphate reduction process in a more restricted marine environment (Figure 11a). Later calcite replacement and precipitation of geode-filling calcite in the siliceous nodules occurred in both sections, with δ13C and δ180 values indicating the participation of meteoric waters in both regions (Figure 11b). Synsedimentary activity of the Penacerrada diapir (Kueper salt - Triassic), which lies close to the Laño section, played a significant role in driving the local shallowing of the basin and in the formation of the silica in the nodules. In contrast, eustatic shallowing of the inner marine series in the Tubilla del Agua section led to the generation of morphologically similar quartz geodes, but from waters not influenced by brines derived from the groundwater halo of a diapir.


Conclusion

This and the previous two articles have underlined the utility of stable isotope samples of brine or precipitates in better understanding the origin of a range of brines and their associated precipitates. But other than the sampling of water molecules in modern brines, the interpretation of all isotope values is equivocal without a petrographic understanding of how and when the sampled textures formed. Stable isotopes of evaporitic minerals with sulphur, carbon and oxygen are the mainstays of isotope work in the study of most evaporite basins, both modern and ancient. Other isotopes that may be useful are 11B and 37Cl, and we shall look at their application to evaporitic sediments in a later blog.

References

Bellanca, A., and R. Neri, 1986, Evaporite carbonate cycles of the Messinian, Sicily; stable isotopes, mineralogy, textural features, and environmental implications: Journal of Sedimentary Petrology, v. 56, p. 614-621.

Cappa Christopher, D., B. Hendricks Melissa, J. DePaolo Donald, and C. Cohen Ronald, 2003, Isotopic fractionation of water during evaporation: Journal of Geophysical Research: Atmospheres, v. 108.

Ellis, P. M., 1986, Post-Miocene carbonate diagenesis of the Lower Cretaceous Edwards Group in the Balcones fault zone area, south-central Texa, in P. L. Abbott, and C. M. Woodruff, eds., The Balcones escarpment, geology, hydrology, ecology and social development in central Texas, Geological Society of America, p. 101-114.

Fu, Q. L., H. R. Qing, K. M. Bergman, and C. Yang, 2008, Dedolomitization and calcite cementation in the Middle Devonian Winnipegosis Formation in Central Saskatchewan, Canada: Sedimentology, v. 55, p. 1623-1642.

Gat, J. R., 1984, The stable isotope composition of Dead Sea waters: Earth and Planetary Science Letters, v. 71, p. 361-376.

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

Horton, T. W., W. F. Defliese, A. K. Tripati, and C. Oze, 2016, Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects: Quaternary Science Reviews, v. 131, p. 365-379.

Lee, M. R., 1994, Emplacement and diagenesis of gypsum and anhydrite in the late Permian Raisby Formation, north-east England: Proceedings - Yorkshire Geological Society, v. 50, p. 143-155.

Miser, D. E., J. S. Swinnea, and H. Steinfink, 1987, TEM observations and X-ray structure refinement of a twiined dolomite microstructure: American Mineralogist, v. 72, p. 188-193.

Moragas, M., C. Martínez, V. Baqués, E. Playà, A. Travé, G. Alías, and I. Cantarero, 2013, Diagenetic evolution of a fractured evaporite deposit (Vilobí Gypsum Unit, Miocene, NE Spain): Geofluids, v. 13, p. 180-193.

Pierre, C., 1988, Application of stable isotope geochemistry to the study of evaporites, in B. C. Schreiber, ed., Evaporites and hydrocarbons: New York, Columbia University Press, p. 300-344.

Rosen, M. R., D. E. Miser, M. A. Starcher, and J. K. Warren, 1989, Formation of dolomite in the Coorong region, South Australia: Geochimica et Cosmochimica Acta, v. 53, p. 661-669.

Shearman, D. J., J. Khouri, and S. Taha, 1961, On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura: Proceedings Geological Association of London, v. 72, p. 1-12.

Shearman, D. J., J. Khouri, and S. Taha, 1961, On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura: Proceedings Geological Association of London, v. 72, p. 1-12.

Stiller, M., J. S. Rounick, and S. Shasha, 1985, Extreme carbon-isotope enrichments in evaporating brines: Nature, v. 316, p. 434.

Testa, G., and S. Lugli, 2000, Gypsum-anhydrite transformations in Messinian evaporites of central Tuscany (Italy): Sedimentary Geology, v. 130, p. 249-268.

Urey, H. C., F. G. Brickwedde, and G. M. Murphy, 1932, A Hydrogen Isotope of Mass 2: Phys. Rev., v. 39, p. 164.

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

Warren, J. K., 2000, Dolomite: Occurrence, evolution and economically important associations: Earth Science Reviews, v. 52, p. 1-81.

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

Worden, R. H., P. C. Smalley, and A. E. Fallick, 1997, Sulfur cycle in buried evaporites: Geology, v. 25, p. 643-646.

 

Stable isotopes in evaporite systems: Part II - 13C (Carbon)

John Warren - Thursday, May 31, 2018

 

Introduction

13C interpretation in most ancient basins focuses on carbonate sediment first deposited/precipitated in the marine realm. Accordingly we shall first look here at the significance of variations in 13C over time in marine carbonates and then move our focus into the hypersaline portions of modern and ancient salty geosystems. In doing so we shall utilize broad assumptions of homogeneity as to the initial distribution of 13C (and 18O) in the marine realm, but these are perhaps oversimplifications and associated limitations need to be recognized (Swart, 2015)

In the next article we shall look at the utility of crossplots of carbon and oxygen isotopes. Stable oxygen isotope values (d18O) crossplotted with respect to carbon isotope values (d13C) from, the same sample creates one of the most widely applied proxies used to infer palaeo-environmental conditions (depositional and diagenetic) in Holocene and ancient carbonate sediments. This is in large part due to kinetic fractionations that occur during evaporation (Leng and Marshall, 2004). It has long been known that as any liquid evaporates, the residual fluid becomes enriched in the less abundant heavy isotope(s) (see Horton et al., 2016 for detailed discussion).

Interpreting 13C

Over the Phanerozoic the standard paradigm for interpreting variations in variations in 13C values from modern and ancient marine carbonate is based on an integration of our understanding of the carbon cycle with the following arguments. Most of the carbon in Earth’s near-surface systems is stored in sedimentary rocks with only about 0.1% in living organisms and the atmosphere-hydrosphere (Figure 1). Oxidized carbon occurs primarily as marine carbonates and reduced carbon as organic matter in sediments. In the carbon cycle, CO2 from the oceans and atmosphere is transferred into sediments as carbonate carbon (Ccarb) or organic carbon (Corg), the former of which monitors the composition of the oceans (Figure 1). The cycle is completed by uplift and weathering of sedimentary rocks and by volcanism, both of which return CO2 to the atmosphere.


There are two stable carbon isotopes, carbon 12 (6 protons and 6 neutrons) and carbon 13 (6 protons and 7 neutrons). Photosynthetic organisms incorporate disproportionately more CO2 containing the lighter carbon 12 than the heavier carbon 13 (the lighter molecules move faster and therefore diffuse more easily into cells where photosynthesis takes place). During periods of high biological productivity, more light carbon 12 is locked up in living organisms and in resulting organic matter that is being buried and preserved in contemporary sediments. Consequently, due the metabolic (mostly photosynthetic) activities of a wide variety of plants, bacteria and archaea, the atmosphere and oceans and their sediments become depleted in carbon 12 and enriched in carbon 13 (Figure 2)


It is assumed that the carbon isotopic ratio in calcareous shells of marine organisms is in equilibrium with that of seawater. So as more carbon 12 is held in biomass during times of high primary productivity, and increased burial of organic carbon, calcareous (CaCO3) skeletal materials become enriched in carbon 13. In contrast during periods of low biological productivity and decreased burial of organic carbon, for example following mass extinctions, marine calcareous skeletal materials become enriched in carbon 12.

Hence plotting variations in carbon isotopes in marine carbonates and organic matter over time offers a way to trace the growth of the crustal reservoir of reduced carbon (Des Marais, 1997). That is, the relative abundance of carbon isotopes is controlled chiefly by: 1) equilibrium isotopic effects among inorganic carbon species, 2) fractionation associated with the biochemistry of organic matter, and 3) the relative rates of burial of carbonate and organic carbon in sediments (Condie 2016).

Because organic matter preferentially incorporates 12C over 13C, there should be an increase in the 13C/12C ratio (as measured by δ13C) in buried carbon with time, and indeed this is what is observed (Des Marais, 1997; Worsley & Nance, 1989). δ13Corg increases from values < -40‰ in the Archaean to modern values of -20 to -30‰. On the other hand, seawater carbon as tracked with δ13Ccarb remains roughly constant with time, with δ13Ccarb averaging about 0%.

Variation in fluxes over time within the carbon cycle can be monitored by an isotopic mass balance (Des Marais, 1997), whereby;

δin = fcarbδ13Ccarb + forgδ13Corg

δin represents the isotopic composition of carbon entering the global surface environment comprised of the atmosphere, hydrosphere, and biosphere. The right side of the equation represents the weighted-average isotopic composition of carbonate (δ13Ccarb) and organic (δ13Corg) carbon buried in sediments, and fcarb and forg are the fractions of carbon buried in each form (fcarb = 1 - forg). For timescales longer than 100 Myr, δin = -5‰, the average value for crustal and mantle carbon (Holser et al., 1988). Thus, where values of sedimentary δ13Ccarb and δ13Corg can be measured, it may be possible to determine forg for ancient carbon cycles. Higher values of δ13Ccarb indicate either a higher value of forg or a greater negativity of average δ13Corg.


During the Phanerozoic, there are several peaks in δ13Ccarb, the largest at about 110, 280, 300, 400, and 530 Ma (Figure 3). These peaks are widely interpreted to reflect an increase in burial rate of organic carbon (Des Marais et al., 1992; Frakes et al., 1992). This is because organic matter selectively enriched in 12C depletes seawater in this isotope, raising the δ13C values of seawater. In the late Paleozoic (300-250 Ma), the maxima in δ13Ccarb correspond to the rise and spread of vascular land plants, which provided a new source of organic debris for burial (Condie 2106, Berner, 1987, 20 01). Also conducive to preservation of organic remains at this time were the vast lowlands on Pangea, which appear to have been sites of widespread swamps where bacterial decay of organic matter is minimized. The drop in δ13Ccarb at the end of the Permian is not understood. Perhaps, large amounts of photosynthetic O2 generated by Carboniferous forests led to extensive forest fires that destroyed large numbers of land plants in the Late Permian (Condie, 2016). However, the reasons for the oscillations in δ13Ccarb are not yet unequivocally resolved and, as in all sciences, the tenet "...perceived correlation does not necessarily equate to causation"must always be at the forefront in the scientific mindset.


Across the Precambrian and the Phanerozoic, the initiation of glaciation on a global scale, as in the Cryogenian ‘Snowball Earth’, has been interpreted to be dependent on parameters like the latitudinal extent of continents and oceanic circulations (Figure 4; Condie, 2016). The main drive for an onset of global glaciation is believed to be the lowering of atmospheric CO2. It likely also requires a continental landmass to be covering one of the earth's polar positions. More recently, cooling related to an increase in the earth's albedo due to widespread evaporites (saline giants) has been added to the list of possible drivers to the onset of glaciation.

Climate modelling studies imply that CO2 concentrations as low as 100–150 ppm are required to initiate global glaciation (e.g. Liu et al., 2013; Feulner and Kienert, 2014). One potential cause of lowered CO2 is drawdown of CO2 during intense silicate weathering in equatorial regions (Hoffman and Schrag, 2002; Goddéris et al., 2003). Photosynthesis provides another mechanism for CO2 drawdown, via conversion of CO2 to O2 and rapid burial of organic carbon, which is reflected in a positive δ13C excursion for carbonates (Pierrehumbert et al., 2011). Additionally, long term cloud cover (Feulner et al., 2015), fluctuations in atmospheric-ocean heat transport, the earth's albedo, or solar luminosity (Pierrehumbert et al. (2011) are also proposed as potential causes of the onset of glaciation (ice-house mode climate).

In a recent paper, Schmid 2017 focused on the cause of the Bitter Springs carbon isotope anomaly, she argues the cause of the pre-glacial, globally recognised, carbon and oxygen isotope variations in carbonate sediments tied to the Bitter Springs anomaly is a response to widespread fractional evaporation of dissolved CO2. This carbon isotope anomaly ties to a well defined correlation with the distribution of Neoproterozoic evaporite basins. She also shows volcanism occurred during the onset of the Bitter Springs Stage (811–788 Ma) and associated widespread evaporite distribution across Australia.


Schmid (op. cit.) argues that the albedo effect began with of the widespread deposition of Rodinian supercontinent evaporites in very shallow marine to epicontinental sedimentary successions beginning ≈810 Ma, increased siliciclastic redbed weathering. This and continuing evaporite deposition and exposure between ≈780 and 720 Ma drove a worldscale increase in Earth's albedo. Such highly reflective salt deposits defined a saline giant across an area that today covers one-third of the Australia continent. Thus, this and other penecontemporaneous saline giants over the Rodinian supercontinent played a potentially significant role in the onset of atmospheric cooling via a significant increase in albedo (Figure 5). These salt beds occur in periods that typify the onset of local (750 Ma) and then global glaciation (720 Ma).


Schmid (2017) goes on to note that the degree of evaporation in the Bitter Springs group sediments is related to the δ13C signature in variably concentrated waters (Figure 6). That is the Tonian Bitter Springs Group (≈830–750 Ma), within the Amadeus Basin in central Australia consists of thick halite and sulphate evaporite accumulations and associated carbonates. The deposition of halite occurred in shallow marine, lagoon (salina) environment (Gillen Formation), and developed into sulphate-dominated supratidal sabkha during sea level regression (Johnnys Creek Formation). The overall regression was interrupted by a transgressive phase lasting at least 20 Ma and leading to deposition of basin-wide stromatolitic dolostone (Loves Creek Formation). The salinity and high evaporation is reflected in positive δ13C in the intercalated carbonates (+4 to +6‰ VPDB) of the evaporitic units, while the shallow marine stromatolitic incursion of the Loves Creek Formation (−2‰ δ13C) show typical marine carbonate isotopic values (Figure 7).

This salinity controlled isotopic separation supports the observations of Stiller et al. (1985) who noted extreme enrichment of 13C in the dissolved inorganic carbon pool in evaporating brines up with δ13C values of up to + 16.5‰ under natural abiotic, oxic conditions in Dead Sea evaporation ponds (Figure 7). The systematic increase in 13C values in highly evaporated waters from the various bittern ponds of the Dead Sea Saltworks is thought to result from a nonequilibrium gas-transfer isotope fractionation. The process of ongoing evaporation leads to CO2 loss within the evaporative brine as less and less gas can held in solution (see Warren 2016, Chapter 9). CO2 exchange in a concentrating surface brine occurs directly between the water column and air, resulting in direct CO2 loss through evaporation. In a sabkha environment. CO2 is released from the hypersaline groundwater through sediments before being released to air as evaporites may form intrasediment precipitates. Overall, atmospheric CO2 uptake in hypersaline settings fed by shallow marine water is diminished compared to the normal marine settings.


Precipitated carbonates modern salinas and sabkhas are mainly aragonite, and formed in association with such evaporative brine, are consistently13C enriched, as seen in nearby Solar Lake and Sabkha Gavish (Figure 2; Stiller et al., 1985; Schidlowski et al., 1984). In a similar fashion, Palaeoproterozoic interbedded shallow marine carbonates, redbeds and evaporites have values up to δ13C + 17.2‰ (Melezhik et al., 1999). Permian and Triassic (Schmid et al., 2006a) redbeds and evaporite sequences also have 13C-rich carbonates (up to +7‰) and enrichment is partly attributed to evaporation and associated CO2 loss (Beauchamp et al., 1987). In modern oceans, atmospheric CO2 is consumed by biological activity and carbonate production originates from mainly marine organisms, leading to near atmospheric to organic negative δ13C signatures in the precipitated sediment(Andersson, 2013).

If increasing salinity leads to unfavourable conditions for photosynthesising organisms to survive (Lazar and Erez, 1992), carbonate through to bittern precipitation becomes increasingly abiotic and evaporation driven, especially at the upper end of the evaporation series. The loss of Ca during evaporation of a brine, via aragonite and calcium sulphate precipitation, leads to an increase in Mg/Ca ratio and an increase in residual brine density. This can result in primary dolomite precipitation or widespread reflux dolomitisation (Schmid et al., 2006, Warren 2000, 2016.

In summary, the typical δ13C signature in normal marine carbonate sediment across much of geological time centres around 0 ‰ and ranges between a few parts per mille on either side of the zero line reflecting precipitation by calcifying and photosynthesising organisms (e.g. algae), while abiotic, evaporation induced carbonates tend to have δ13C values above +1‰. More positive δ13C values (+4 to +6‰) tend to typify dominantly abiotic carbonates (and local methanogenic carbonates with even more positive values) and support the notion of evaporation-driven 13C-enrichement in times of widespread evaporitic epeiric and basinwide carbonates. In the Precambrian, widespread marine stromatolitic units such as, algal Loves Creek Formation reflects δ13C values for biogenic carbonate precipitation under shallow marine, non-hypersaline conditions. The change from a shallow hypersaline lagoon towards evaporitic mudflats and salterns suggests an increase in aridity and continentality/hydrographic isolation, with associated more positive δ13C values.

Implications for some types of 13C anomaly

The Bitter Springs Group chemostratigraphy has been correlated globally and the negative excursion was named previously after this unit (Bitter Springs Stage anomaly). However, the mechanism of evaporation-driven fractionation of δ13C is different from the commonly proposed inorganic-organic carbon fractionation, and challenges the views on interpreting global chemostratigraphic anomalies or excursion and their cause. Evaporite basins covered vast regions worldwide prior to the Sturtian glaciation, e.g. the Australian evaporites would have covered a third of the continent. The light surface of evaporites and associated carbonates would have had a high albedo and effectively cause less surface heat absorption. This subsequently would have triggered temperature decrease on a continental and possibly global scale. The Schmid paper hypothesises that the deposition of evaporites worldwide would have contributed to global cooling starting ≈100 Ma prior to Snowball Earth and would have played an important role in the onset of global glaciation.

References

Andersson, A.J., 2013. The oceanic CaCO3 cycle. In: T. Holland (Editor), Treatise on Geochemistry, 2nd ed. Elsevier, pp. 519-542.

Beauchamp, B., Oldershaw, A.E. and Krouse, H.R., 1987. Upper Carboniferous to Upper Permian 13C-enriched primary carbonates in the Sverdrup Basin, Canadian Arctic: comparisons to coeval western North American ocean margins. Chem. Geol. , 65: 391-413.

Berner, R.A., 1987. Models for carbon and sulfur cycles and atmospheric oxygen; application to Paleozoic geologic history. American Journal of Science, 287: 177-196.

Berner, R.A., 2001. Modeling atmospheric O2 over Phanerozoic time. Geochimica et Cosmochimica Acta, 65: 685-694.

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

Des Marais, D.J., 1997. Isotopic evolution of the biogeochemical carbon cycle during the Proterozoic Eon. Organic Geochemistry, 27(5): 185-193.

Des Marais, D.J., Strauss, H., Summons, R.E. and Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature, 359: 605-609.

Feulner, G., Hallmann, C. and Kienert, H., 2015. Snowball cooling after algal rise. Nat. Geosci. , 8: 659-662.

Feulner, G. and Kienert, H., 2014. Climate simulations of Neoproterozoic snowball Earth events: similar critical carbon dioxide levels for the Sturtian and Marinoan glaciations. Earth Planet. Sci. Lett., 404: 200-205.

Frakes, L.A., Francis, J.E. and Syktus, J.L., 1992. Climate modes of the Phanerozoic. Cambridge University Press, New York, 274 pp.

Goddéris, Y., Donnadieu, Y., Nédélec, A., Dupré, B., Dessert, C., Grard, A., Ramstein, G. and François, L.M., 2003. The Sturtian ‘snowball’ glaciation: fire and ice. Earth Planet. Sci. Lett. , 211: 1-12.

Hoffman, P.F. and Schrag, D.P., 2002. The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14: 129-155.

Holser, W.T., Schidlowski, M., Mackenzie, F.T. and Maynard, J.B., 1988. Geochemical cycles of carbon and sulfur. In: C.B. Gregor, R.M. Garrels, F.T. Mackenzie and J.B. Maynard (Editors), Chemical cycles in the evolution of the earth. John Wiley, New York, pp. 105–173.

Horton, T.W., Defliese, W.F., Tripati, A.K. and Oze, C., 2016. Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects. Quaternary Science Reviews, 131: 365-379.

Lazar, B. and Erez, J., 1992. Carbon geochemistry of marine-derived brines: I. 13C depletions due to intense photosynthesis. Geochim. Cosmochim. Acta, 56: 335-345.

Leng, M.J. and Marshall, J.D., 2004. Paleoclimate interpretation of stable isotope data from lake sediment archives. Quaternary Science Reviews, 23(811-831).

Liu, Y., Peltier, W.R., Yang, J. and Vettoretti, G., 2013. The initiation of Neoproterozoic ‘‘snowball” climates in CCSM3: the influence of paleocontinental configuration. Climate Past, 9: 2555-2577.

Melezhik, V.A., Fallick, A.E., Medvedev, P.V. and Makarikhin, V.V., 1999. Extreme 13Ccarb enrichment in ca. 2.0 Ga magnesite-stromatolite-dolomite-‘red beds’ association in a global context: a case for the world-wide signal enhanced by a local environment. Earth Sci. Rev., 48: 71-120.

Pierrehumbert, R.T., Abott, D.S., Voigt, A. and Koll, D., 2011. Climate of the neoproterozoic. Annu. Rev. Earth Planet. Sci., 39: 417-460.

Schidlowski, M., Matzigkeit, U. and Krumbein, W.E., 1984. Superheavy organic carbon from hypersaline microbial mats; Assimilatory Pathway and Geochemical Implications. Naturwissenschaften, 71(6): 303-308.

Schmid, S., 2017. Neoproterozoic evaporites and their role in carbon isotope chemostratigraphy (Amadeus Basin, Australia). Precambrian Research, 290: 16-31.

Schmid, S., Worden, R.H. and Fisher, Q., 2006. Carbon isotope stratigraphy using carbonate cements in the Triassic Sherwood Sandstone Group: Corrib Field, west of Ireland. Chem. Geol., 225: 137-155.

Stiller, M., Rounick, J.S. and Shasha, S., 1985. Extreme carbon-isotope enrichments in evaporating brines. Nature, 316: 434.

Swart, P.K., 2015. The geochemistry of carbonate diagenesis: The past, present and future. Sedimentology, 62(5): 1233-1304.

Warren, J.K., 2000. Dolomite: Occurrence, evolution and economically important associations. Earth Science Reviews, 52(1-3): 1-81.

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

Worsley, T.R. and Nance, R.D., 1989. Carbon redox and climate control through Earth history: A speculative reconstruction. Paleogeography, Paleoclimatology, Paleoecology, 75: 259-282.

 

Stable isotopes in evaporite systems: Part I: Sulphur

John Warren - Monday, April 30, 2018

 

Introduction

The sulphur isotopic composition of sulphate dissolved in modern seawater (SW), and the relationship with the associated modern and ancient sulphate precipitates, has been studied for more than five decades. An understanding of the controlling factors is fundamental in any interpretation of the origin of modern and ancient sedimentary calcium sulphates.

So, we shall look at the significance of sulphur isotopes, first by reviwing what is known in terms of the isotopic evolution of marine sulphate salts across the evaporation series from gypsum to the bitterns, and then across a time perspective via the evolution of oceanic sulphate and sulphide signatures from the Archean to the present.


Sulphur isotopes across the bittern series

The accepted d34S value of modern seawater-derived calcium sulphate (gypsum) is + 20.0 ±0.2‰ (Sasaki, 1972; Zak et al., 1980 and references therein). This is a average value, based on numerous analyses across the range ( +19.3 to +21.4‰). Notably, Rees et al. (1978) obtained a mean of +20.99 ± 0.09‰, using the SF6 method, which has a better reproducibility than the conventional S02 method. Mediterranean seawater gave a d34S value of +20.5‰ (Nielsen, 1978).

Measured values in natural gypsum from seawater show initial precipitates have a d34S value slightly higher than that of its source brine (Figure 1). The highest isotope differential for gypsum naturally precipitated from seawater, as recorded in the literature, is +4.2‰ (Laguna Madre, Texas, U.S.A.; Thode, 1964). Most reported d34Sgypsum-sw differentials lie in range from 0 to + 2.4‰ (Ault and Kulp, 1959; Thode et al., 1961; Thode and Monster, 1965; Holser and Kaplan, 1966).

Prior to Raab and Spirto (Figure 1; 1991), laboratory experiment data on d34Sgypsum-solution are scarce, especially for solutions mimicking initial precipitation of gypsum from natural seawater and passing into halite saturation. Harrison (1956) measured a d34Sgypsum-solution value of ~ + 2‰ for gypsum precipitated from an artificial solution, that was saturated with respect to gypsum. Thode and Monster (1965) calculated a K-value [ (32S/34S)solution/ (32S/34S)gypsum] of 1.00165 from a measured a d34Sgypsum-solution value of + 1.65‰ for a CaSO4.2H2O -saturated solution, evaporated under reduced pressure and allowed to age and equilibrate for 24 months at room temperature. An experiment using natural seawater was carried out by Holser and Kaplan (1966), who sampled the products of evaporating seawater in a tank with continuous refilling (green circles in Figure 1). The results show “only a small difference between brine and gypsum precipitated” (Holser and Kaplan, 1966, p.97), resulting in a mean value of d34Sgypsum-seawater = +1.7‰ (+19.4 to +21.1‰). Harrison (1956) calculated from experimental vibrational frequencies for S04 in solution and in crystalline CaSO4.2H2O, a constant K = 1.001 for the reaction:

(Ca34S04.2H2O + 32SO4)SOLID = (Ca32S04.2H2O + 34SO4)SOLUTION

which means a 1‰ increase of d34S in the solid fraction. Nielsen (1978, p. 16-B-20), using Rayleigh-type fractionation curves indicates that, “...the gypsum/anhydrite of the sulphate facies should be slightly enriched in 34S with respect to the unaffected seawater sulphate”

In the geological record the evaporites of the later Mg- and K-Mg- sulphate bittern facies are depleted in 34S relative to the earlier, basal Ca-sulphates, as rseen in the geological record. Nielsen and Ricke (1964, p.582) give a mean value of +2‰ for the depletion in 34S in later bittern evaporite sulphates relative to the basal Ca-sulphates in the Upper Permian Zechstein Series (Hattorf and Reyershausen, Germany) whereas Holser and Kaplan (1966, pp. 116 and 117) give a value of -1.0±0.8‰ (their d34Spotash-magnesia facies sulphates - d34Sgypsum/anhydrite facies) for the Zechstein Basin (Germany) and -0.8±0.5‰ for the Upper Permian Delaware Basin (U.S.A.) evaporites (green circles in Figure 1).

Theoretical calculations of the behaviour of the sulphur isotopic fractionation during the late evaporation stages were made by Holser and Kaplan (1966, pp. 116 and 117, fig. 4) and by Nielsen (1978, p. 16-B-20, fig. 16-B-12) applying the Rayleigh distillation equation and using the same fractionation factor calculated from the initial gypsum (1.00165). Their curves are thus in a continuous line with those calculated for the Ca-sulphates. These show an increasing degree of depletion in 34S in the sulphates precipitated in the course of the progressive evaporation in a closed basin, relative to the first Ca-sulphate precipitated, up to the end of the carnallite facies. They explain it by the continuous depletion in 34S in the brines. Thus their calculated d34Scrystal-initial gypsum at the end of the halite facies is ~ -0.6‰, at the end of the Mg-sulphate facies -1.0‰, and at the beginning of the carnallite facies -3.8‰, and relative to the original seawater (their 34SC) the differences are +1.0, +0.4 and -2.2‰, respectively. Nielsen (1978) also plotted an extrapolated fractionation curve for the residual brines in a closed reservoir, indicating that the brine is constantly depleted by 1.65‰ relative to the associated precipitate.

Prior to the laboratory work of Raab and Spiro (1991), no experimental data pertaining to the isotopic behaviour of sulphate sulphur in the late evaporative stages of seawater was available in the literature. Raab and Spiro evaporated seawater, stepwise and isothermally at 23.5°C, for 73 days, up to a degree of evaporation of 138x by H2O weight. At various stages of evaporation the precipitate was totally removed from the brine and the brine was allowed to evaporate further. The sulfur isotopic compositions of the precipitates and related brines showed the following characteristics (Figure 1) where the initial d34S of the original seawater is +20‰. The d34S of both precipitates and associated brines decrease gradually across the gypsum field nd aup to the end of the halite field, where d34Sprecipitate = + 19.09‰ and d34Sbrine = + 18.40‰. The precipitates are always enriched in 34S relative to the associated brines in these fields, but the enrichment becomes smaller towards the end of the halite field. A crossover, where the d34S value of the brines becomes higher than those of the precipitates, occurs at the beginning of the Mg-sulfate field. The d34Sprecipitate increases from + 19.09‰ at the end of the halite field through +19.35‰ in the Mg-sulfate field to + 19.85‰ in the K-Mg-sulfate field, whereas the d34Sbrine increased from +18.40‰, through +20.91‰ to +20.94‰, respectively.

This evolution implies different values of fractionation factors (a) for the minerals precipitated in the late halite, Mg-sulphate and K-Mg-sulphate fields, other than that for gypsum (1.00165). The value of aprecipitate-residual brine would then be very slightly >1 in the late halite field and >1 in the two later fields.

The experimental pattern of evolution of the d34S-values of the precipitates from their experiment is in good agreement with data for natural anhydrites interbedded in halites, where d34S-values are lower relative to basal gypsum (and secondary anhydrite), and of primary minerals of the Mg- and K-Mg-sulfate facies, reported in evaporitic sequences, such as those of the Delaware (U.S.A.) and of the Zechstein (Germany) basins and so can be used to better interpret a marine origin of the sulphate bitterns.


Ancient oceanic sulphate

The element sulphur is an important constituent of the Earth’s exogenic cycle. During the sulphur cycle, 34S is fractionated from 32S, with the largest fractionation occurring during bacterial reduction of marine sulphate to sulphide. Isotopic fractionation is expressed as d34S, in a manner similar to that used for carbon isotopes and the longterm carbon curves related to the sulphur isotope curve across deep time (see next article). Sedimentary sulphates (mostly measured on anhydrite, but also baryte) typically are used to record the isotopic composition of sulphur in seawater (Figure 2). Mantle d34S is near 0‰, and bacterial reduction of sulphate to sulphides (mostly as pyrite) strongly prefers 32S, thus reducing d34S in organic sulphides to negative values (≈ -18‰), so leaving oxidized sulphur species with approximately equivalent positive values (+17‰; Figure 3).


Historically, the sulphur cycle has been interpreted as being largely controlled by the biosphere and in particular by sulphate-reducing bacteria that inhabit shallow marine waters (Strauss, 1997). Typically, sulphur occurs in its oxidized form as dissolved sulphate in seawater or as evaporitic sulphate and in its reduced form as sedimentary pyrite. The isotopic compositions of both redox states are sensitive indicators for changes of the geological, marine geochemical or biological environments in the past (Figure 2). The isotope record of marine sedimentary sulphate through time has been used successfully to determine global variations of the composition of seawater sulphate.

The isotopic composition of sedimentary (biogenic) pyrite reflects geochemical conditions during its formation via bacterial sulphate reduction. Sedimentary pyrite is, thus, an important record of evolutionary (microbial) processes of life on Earth. Both time records (anhydrite and pyrite) have been combined in an isotope mass balance calculation, and changes in burial rates of oxidized vs. reduced sulphur can be determined (Strauss, 1997). This, in turn, yields important information for the overall exogenic cycle (i.e. the earth's oxygen budget as discussed in the next article).

And so, values preserved in ancient marine sulphate evaporites are part of the broader world sulphur cycle across deep time that includes movements in and out of marine sulphides (dominantly pyrite) and marine baryte precipitates (Figure 2). Values based on evaporitic CaSO4 are consistent with the ranges seen in modern gypsum (Figure 3). A plot of ancient marine CaSO4 evaporites shows the oxxidised sulphur curve for seawater has varied across time from +30‰ in the Cambrian, to around +10‰ in the Permian and that it increased irregularly in the Mesozoic to its present value of +20‰ (Figure 4). Oxygen values show much less variability and will be discussed in more detail in the next article in this series. Time-consistent variations are reflected in all major marine sulphate evaporite deposits and were most likely controlled by major input or removal of sulphides from the oceanic reservoirs during changes driven by longterm variations in tectonic activity and weathering rates.

Historically, simple removal of oceanic sulphate via an increase in the volume of megasulphate deposition in a saline giant was not thought to be accompanied by dramatic isotopic effects. Rather, variations within the global sulphur cycle were thought to be controlled by a redox balance with stored sulphides and organics in more reducing environments, which are also linked to the carbon cycle and the atmospheric oxygen budget.

In this scenario the oxidative part of the global sulphur cycle is largely governed by continental weathering (especially of marine black shale), riverine transport and evaporite deposition, while the reduced part of the sulphur cycle is controlled by levels of fixation of reduced sulphur-bearing compounds in the sediment column, mostly as pyrite via bacterial sulphate reduction (Figure 2.). The latter process preferentially removes isotopically light sulphur from seawater and so increases the d34S value in the ocean, and any consequent precipitate.

However, more recent work question aspects of this older sulphur cycle/pyrite/organics model. As just discussed, variations in d34Ssulphate across the Phanerozoic are traditionally interpreted to reflect changes in the total amount of sulphur buried as pyrite in ocean sediments — a parameter referred to as fpyr and defined as (Hurtgen, 2012);

fpyr = [(pyrite Sburial)/(pyrite Sburial + evaporite S burial)].

However, Wortmann and Paytan (2012) conclude that the 5‰ negative d34Ssulphate shift in ~120-million- year-old rocks was caused by massive seawater sulphate removal, which accompanied large-scale evaporite deposition during the opening of the South Atlantic Ocean (Figure 4). In their model, the negative d34Ssulphate shift is driven by lower pyrite burial rates that result from substantially reduced marine sulphate levels in the world ocean, tied to megasulphate precipitation. The authors attribute a 5‰ positive d34Ssulphate shift in the world’s oceans about 50 million years ago to an abrupt increase in marine sulphate concentrations as a result of large-scale dissolution of freshly exposed evaporites; they argue that the higher sulphate concentrations in the ocean in turn led to more pyrite burial.


Likewise, Halevy et al. (2012 ) studied past sulphur fluxes to and from the ocean, but over a longer time-frame (the Phanerozoic). They quantified sulphate evaporite burial rates through time, then scaled these rates to obtain a global estimate of variation in sulphur flux. Their results indicate that sulphate burial rates were higher than previously estimated, but also greatly variable. When Halevy et al. (2012) integrated these improved evaporite burial fluxes with seawater sulphate concentration estimates and sulphur isotope constraints, their calculations implied that Phanerozoic fpyr values (fpyr = fraction of sulphur removed from the oceans as pyrite) were ~100% higher on average than previously recognized. These surprisingly high and constant pyrite burial outputs must have been balanced by equally high and constant inputs of sulphate to the ocean via sulphide oxidation (weathering). These relatively high and constant rates of pyrite weathering and burial over the Phanerozoic, as identified by Halevy et al. (2012, suggest that the consumption and production of oxygen via these processes played a larger role in regulating Phanerozoic atmospheric oxygen levels than previously recognized, perhaps by as much as 50%.

Both studies recognize the importance of episodic evaporite burial on the sulphur cycle, while Wortmann and Paytan (2012) clearly show that large-scale deposition and dissolution of sulphate evaporites over relatively short geologic time scales can have an enormous impact on marine sulphate concentrations, pyrite burial rates, and the carbon cycle and so probably play a more important role than previously recognised in regulating the chemistry of the ocean atmosphere system.

The 18O content in seawater sulphate fluctuates less than sulphur values over geologic time (see next article for detailed discussion). The isotopic composition of sulphate minerals varied only slightly from the Neoproterozoic to the Palaeozoic decreasing from +17 to +14‰ (Figure 4). Values then rose during the Devonian to reach +17‰ during the Early Carboniferous (Mississippian). Values then fell to =+10‰ during the Permian, mimicked by a similar decline in sulphur values in the Late Permian to Early Triassic. Since the rise to +15‰ in the Early Triassic, values of marine sulphate minerals have remained close to +14‰ (add 3.5‰ to mineral determined value to give ambient seawater value). Overall, oxygen values show little correlation with marine sulphate variation and are perhaps are more controlled by sulphide weathering reactions.

What is also significant is that, given the now well established sulphur isotope age curve, a comparison of a measured d34S value from an anhydrite or gypsum of known geological age to the curve allows an interpretation of a possible marine origin to the salt. A value which differs from the marine signature does not necessarily mean a nonmarine origin, but, at the least, it does mean diagenetic reworking or, more likely, a groundwater-induced recycling of sulphate ions into a nonmarine saline lake (Pierre, 1988). Such oxygen and sulphur isotopic crossplots have been used to establish the continental (nonmarine) origin of the Eocene gypsum of the Paris Basin and the upper Miocene gypsum of the Granada basin, with sulphate derived from weathering of uplifted Mesozoic marine evaporites (Fontes and Letolle, 1976; Rouchy and Pierre, 1979; Pierre, 1982).

Sulphur is largely resistant to isotopic fractionation during burial alteration and transformation of gypsum to anhydrite (Figure 5; Worden et al., 1997). For example, primary marine stratigraphic sulphur isotope variation is preserved in anhydrites of the Permian Khuff Formation, despite subsequent dehydration to anhydrite during burial (≈1,000m) and initial precipitation as gypsum from Permian and Triassic seawater. Gypsum dehydration to anhydrite did not involve significant isotopic fractionation or diagenetic redistribution of material in the subsurface. At depths greater than 4300 m, the same sulphur isotope variation across the Permian-Triassic boundary is still present in elemental sulphur and H2S, both products of the reaction of anhydrite with hydrocarbons via thermochemical sulphate reduction (Figure 5). Clearly, thermochemical sulphate reduction did not lead to sulphur isotope fractionation. Worden et al. also argues that significant mass transfer has not occurred in the system, at least in the vicinity of the Permian-Triassic boundary, even though elemental sulphur and H2S are both fluid phases at depths greater than 4300 m. Primary differences in sulphur isotopes have been preserved in the rocks and fluids, despite two major diagenetic overprints that converted the sulphur in the original gypsum into elemental sulphur and H2S by 4300 m burial and the potentially mobile nature of some of the reaction products. That is, all reactions occurred must have occurred in situ; there was no significant sulphur isotope fractionation, and only negligible sulphur was added, subtracted, or moved internally within the system.


The resistance to fractionation of sulphur isotopes in subsurface pore waters can also be utilised to determine the origin of saline thermal pore waters. In a study of sulphur isotopic compositions of waters in saline thermal springs, Risacher et al. (2011) came to the interesting conclusion that dissolution of continental sedimentary gypsum from the Tertiary-age Salt Cordillera was the dominant supplier of sulphate (Figure 6). The sulphate in the springs was not supplied by the reworking of volcanic sulphur in this active volcanic terrain. d34S values from 3 to 11‰ in continental gypsum and this also encompasses the range of d34S in pedogenic gypsum (5 to 8‰) and in most surface waters (3.4 to 7.4‰) including salt lakes (Rech et al., 2003). Frutos and Cisternas (2003) found isotope ratios ranging from 1.5 to 10.8‰ in five native sulphur samples. Figure 6 presents the sulphur isotope ratio of dissolved sulphate in thermal waters sampled by Risacher et al. (2011) and references therein. The d34S of sulphate in northern thermal springs is within the range of salt lakes waters and continental gypsum. In an earlier paper Risacher et al. (2003) showed that salar brines leak through bottom sediments and are recycled in the hydrologic system. Deep circulating thermal waters are dissolving continental gypsum in sedimentary layers below the volcanics associated with the present day salars. The exception to this observation is the sulphur in Tatio springs where Cortecci et al. (2005) proposed a deep-seated source for the sulphate, related to magma degassing (Figure 6).


References

Cortecci, G., Boschetti, T., Mussi, M., Herrera Lameli, C., Mucchino, C. and Barbieri, M., 2005. New chemical and original isotopic data on waters from El Tatio geothermal field, northern Chile. Geochemical Journal 39: 547-571.

Fontes, J.C. and Letolle, R., 1976. 18O and 34S in the upper Bartonian gypsum deposits of the Paris Basin. Chemical Geology, 18(4): 285-295.

Frutos, J. and Cisternas, M., 2003. Isotopic Differentiation in Volcanic-Epithermal Surface Sulfur Deposits of Northern Chile: d34S < 0‰ in “Fertile” Systems (Au-Ag-Cu Ore Deposits below), versus d34S ≥ 0‰ for “Barren” Systems. Short Papers - IV South American Symposium on Isotope Geology (Salvador, Brazil, 2003): 733-735.

Halevy, I., Peters, S.E. and Fischer, W.W., 2012. Sulfate Burial Constraints on the Phanerozoic Sulfur Cycle. Science, 337(6092): 331-334.

Holser, W.T. and Kaplan, I.R., 1966. Isotope geochemistry of sedimentary sulfates. Chemical Geology: 93-135.

Hurtgen, M.T., 2012. The Marine Sulfur Cycle, Revisited. Science, 337(6092): 305-306.

Nielsen, H., 1978. Sulfur isotopes in nature. In: K.H. Wedepohl (Editor), Handbook of Geochemistry Section 16B, pp. B1 - B40.

Nielsen, H. and Ricke, W., 1964. Schwefel-lsotopenverhältnissen von Evaporiten aus Deutschland; Ein Beitrag zur Kenntnis von d34S im Meerwasscr-Sulfat. Geochimica et Cosmochimica Act, 28: 577-591.

Pierre, C., 1982. Teneurs en isotopes stables (18O, 2H, 13C, 34S) et conditions de genese des evaporites marines; application a quelques milieux actuels et au Messinien de la Mediterranee. Doctoral Thesis, Orsay, Paris-Sud.

Raab, M. and Spiro, B., 1991. Sulfur isotopic variations during seawater evaporation with fractional crystallization. Chemical Geology: Isotope Geoscience section, 86(4): 323-333.

Rech, J.A., Quade, J. and Hart, W.S., 2003. Isotopic evidence for the source of Ca and S in soil gypsum, anhydrite and calcite in the Atacama Desert, Chile. Geochimica et Cosmochimica Acta 67(4): 575-586.

Rees, C.E., Jenkins, W.J. and Monster, J., 1978. The sulfur isotopic composition of ocean water sulphate. Geochimica et Cosmochimica Acta, 43: 377-381.

Risacher, F., Fritz, B. and Hauser, A., 2011. Origin of components in Chilean thermal waters. Journal of South American Earth Sciences, 31(1): 153-170.

Rouchy, J.M. and Pierre, C., 1979. Donnees sedimentologiques et isotopiques sur les gypses des series evaporitiques messiniennes d'Espagne meridionale et de Chypre. Rev. Geogr. Phys. Geol. Dyn., 21(4): 267-280.

Sasaki, A., 1971. Variation in sulfur isotope composition of oceanic sulfate. 14th Int. Geol. Congr. Sect. 1: 342-345.

Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time. Palaeogeography Palaeoclimatology Palaeoecology, 132: 97-118.

Thode, H.D., 1964. Stable isotopes a key to our understanding of natural processes. Bulletin Canadian Petroleum Geologists, 12: 246-261.

Thode, H.G. and Monster, J., 1965. Sulfur-Isotope Geochemistry of Petroleum, Evaporites, and Ancient Seas, Fluids in Subsurface Environments. AAPG Memoir 4, pp. 367-377.

Worden, R.H., Smalley, P.C. and Fallick, A.E., 1997. Sulfur cycle in buried evaporites. Geology, 25(7): 643-646.

Wortmann, U.G. and Paytan, A., 2012. Rapid Variability of Seawater Chemistry Over the Past 130 Million Years. Science, 337(6092): 334-336.

Zak, I., Sakai, H. and Kaplan, R., 1980. Factors controlling the 18O/16O and 34S/32S isotopic ratios of ocean sulfates and interstitial sulfates from modern deep sea sediments. In: E.D. Goldberg, Y. Horibe and K. Saruhaki (Editors), Isotope Marine Chemistry. Geochem. Res. Assoc, Tokyo, pp. 339-373.


 

Well (wireline) log interpretation of evaporites: An overview

John Warren - Saturday, March 31, 2018

Introduction

Often, when I run an evaporite training program for a client in the hydrocarbon or the potash industries, I am asked to add a short training module on the identification of evaporites in a set of conventional wireline log outputs. This blog is an overview of what I discuss in such a module. But every evaporite basin has its own set of mineralogies and problems and a generalised discussion, as in this blog, must be refined to meet the needs of the drilling or mining program in a particular evaporite basin.

Significant thickness units of evaporites are rarely cored in oil and gas drilling, unless in error, while when drilling rock chips of the more soluble salt minerals are quickly dissolved in most drill muds; so only a small portion of any subsurface evaporite bed can be studied directly. The situation is somewhat different in Salt and potash mining where cores are commonly collected ahead of the mine face to ascertain ore extent and thickness. Increasingly core calibrated well logs are replacing the need for extensive coring when ascertaining and predicting ore quality.

Many evaporite properties can be ascertained by examining a suite of conventional wireline logs. Many evaporite beds contain only one or two dominant saline minerals, they lack free pore fluids and have negligible porosity. This dramatically simplifies log interpretation and enhances the reliability of inferences with respect mineralogy. Thick clean evaporites will show the same characteristic set of log responses, not only locally but according to some authors worldwide (e.g., Serra, 1984, p. 173; Warren, 2006, Chapter 10). The most commonly available logs for the study of evaporites are the logs measuring hole diameter, electrical properties, bulk density, neutron porosity logs, sonic logs and, if significant levels of potash salts are present, both gamma and multispectral gamma logs.


Evaporites as seen in well logs

Well-logs are a continuous recording of a geophysical paragf meter along a borehole, where the value of the measurement is continuously plotted against depth in the borehole. Currently, the well logging industry is transitioning from wireline or cable-based well logging tools (Figure 1) to the increasing use of well-log tools designed for use in directional drilling. Wireline or cable tools can only be utilised in vertical to steeply inclined wells. The same set of conventional well log measurements are now increasingly collected using MWD (measurement while drilling) and LWD (logging while drilling) tools. With MWD/LWD, measurements are made by a suite of well-logging tools that reside immediately behind the advancing drill-bit. Part of the data collected by these tools is sent to the surface in real time (MWD) by mud pulsing or some other method of telemetry. The remaining portion of the collected data (LWD) is stored on a hard disk and recovered typically when a worn drill-bit is pulled to the surface to be replaced.

Although there are numerous well-logging tools and measurements that can be used in the study of evaporites, this section deals with only a few of the more conventional logging methods. Rider (1996) is an excellent overview from a geological, not petrophysical, perspective of the general principles of well-log interpretation. For a more comprehensive discussion of the geological applications of well-logs, there are many logging-company manuals, as well as excellent books and articles such as Kruger (2014), Crain (2010), Ellis and Singer (2007), Nelson (2007), Rider (1996), Nurmi (1978), and Alger and Crain (1966), Crain and Anderson (1966).


Electrical properties

Electrical resistivity, the reciprocal of electrical conductivity, is the degree with which a formation opposes the flow of electrical current. Onshore, a log of the spontaneous potential of a formation is run at the same time as a resistivity log. In reality, the measured resistivity is dependent on the combined resistivity of both the rock matrix and any contained fluids. Most solid rock materials are insulators, while their enclosed fluids are conductors. Hydrocarbons are the exception to fluid conductivity; they are infinitely resistive, and this is the basis for the quick look-identification of hydrocarbons and the use of Archies Law to determine water saturation levels in potential hydrocarbon reservoirs. In terms of evaporite identification, most evaporite units contain little if any pores or free water and so have very high resistivities compared to other more porous units (Table 1).

When the evaporite unit is relatively pure and monomineralogic, it creates a distinctive blocky log shape, whereas when it entrains beds of thin more porous lithologies (mudstones, shale, sands, limestone, dolomite) or perhaps contains brine-filled cavities and vugs, a muh spikier log is seen across a saline interval. The actual wireline log signature depends on the content of brine, sand, clay, bitumen and other variables. Within a local area in a basin, an elevated resistivity signature, although it does not allow a first indication of the presence of evaporite salts, can subsequently confirm it. When bitumens and salts are present in the same interval (as in salt-encased EoCambrian carbonate-slither reservoirs in the South Oman Salt Basin), the co-occurrence of halite cement, anhydrite cement and bitumen complicates a reliable interpretation of movable hydrocarbons.

Total & Spectral gamma-ray logs

The gamma-ray or gamma log is a record of the formation’s radioactivity. The radiation emanates from uranium, thorium and potassium which occur naturally in the formation. A simple gamma-ray log measures the radioactivity of the three radiogenic elements (U, K, Th) combined, while the spectral gamma log shows the amount of each radiogenic element contributing to a formation's radioactivity.

As a first indicator of lithology in non-evaporitic intervals, the gamma log is extremely useful in suggesting where shale may be expected in a formation; worldwide, elevated gamma readings in a sandstone-mudstone succession are typically used to indicate shaliness of the formation (Vclay). Clays can contain high levels of potassium-containing minerals, thorium (another radiogenic mineral) tends to be fixed in shales (c.f. sands), and that clays typically “fix” marine uranium into the sediment in three main ways i) chemical precipitation in acid (pH 2.5 - 4.0) or reducing environments, ii) adsorption by organic matter in the clays, iii) adsorption by phosphates in the clays. Uranium can also be mobilised and “re-fixed in the subsurface across redox interfaces. High gamma readings can also be due to the elevated potassium content in glauconite-rich sands, or the secondary movement of uranium to form “hot” cement and fissure fills in a number of Middle East reservoirs.


More importantly for our purposes, high gamma signatures can be associated with those evaporites which contain high proportions of the potassium salts such as sylvite, carnallite, and polyhalite (Table 1; Figure 2). In the potash-entraining salts, there is between 10% and 50% potassium by weight. When it is considered that the average shale contains only 2.7% potassium, the very strong radioactivity indicative of the potassium salts in an evaporite suite is understandable and means potash beds can be distinguished from the somewhat-elevate uranium-derived kicks of marine shale and the low radiogenic content of any adjacent halite, anhydrite or carbonate beds.

In contrast to halite units containing potassium salts, the more common evaporites, such as halite and anhydrite, give very low readings on the gamma log scale (Figure 2). Once an initial tie-back to core-determined assay values is done, it is possible to reliably estimate the percentage of K2O from the gamma response. As a general "rule of thumb," Edwards et al. (1967) showed that for a 6.25-inch, liquid-filled hole there was a correlation of 12.6 API units per 1% K2O.


Bulk Density Log

The bulk density or density log is related to the electron density of a formation and is the near-numerical equivalent of the formation's specific gravity (gm/cc); that is it is considered to measure variations in the average total density of the formation. A tool-measured value includes the density of the solid rock matrix and the density of fluids enclosed in the pores. The bulk density log is a measure of the degree of scattering or attenuation of gamma rays by electrons in the formation (Compton scattering). The electron density of a formation (electrons/cc) is closely related to the common density (gm/cc) and is typically used as a direct indicator of common density. Unfortunately, some minerals, including halite and sylvite, have electron densities that are not directly proportional to their specific gravities. Such minerals require the use of apparent bulk density for interpretation. Fortunately, many of the evaporite minerals have sufficient differences in bulk density to be recognised especiallry when crossplotred against Pef or NPHI (neutron) values (Figure 3, 4 and 5).

Many evaporite units are relatively pure and often mono- or bi-mineralic. Because of this, their lithological composition can be suspected, if not positively identified from the density log. However, when impure, the densities will fluctuate. Fortunately, most relatively pure evaporites thicker than a metre tend to give intervals of constant density with only minor variation. When this occurs, densities near the expected value in a clean evaporite unit can be easily identified and correlated to mineralogy using the bulk density log.

Neutron Logs

The neutron porosity index or neutron log provides a continuous record of a formation’s reaction to fast neutron bombardment. It is primarily a measurement of the hydrogen concentration in the formation, whether from the water of hydration, as in the case of hydrated salts such as gypsum and carnallite, or from water or oil in the more commonly understood non-evaporite situation. Quantitatively, the neutron log is used to measure porosity (in limestone-equivalent porosity units), qualitatively, it is a good discriminator between oil and gas in intervals without hydrated salts or other minerals. Geologically, it can be used to identify gross lithology, and so define evaporites (negative porosity values), hydrated minerals, and volcanic rocks and zeolites. A crossplot of formation bulk density versus neutron-log measurement is an extremely valuable tool for identifying various subsurface evaporite lithologies (Figure 2).

For example, in thick evaporite successions, a neutron log can distinguish between various evaporite salts on the basis of water of crystallisation (Rider, 1996). Gypsum is the most common of the evaporites containing water of crystallisation. However, carnallite, polyhalite, and kainite also contain the water radical (Table 1). In a neutron-density (NPHI-RHOB) crossplot, all these hydrated salts have high neutron-log values and characteristic tightly-clustered apparent bulk densities, which separates them from other anhydrous evaporites such as salt or anhydrite, which contain no water and hence have NPHI values near zero (Figure 4).


Sonic or Acoustic Logs

The sonic or acoustic log shows a formation’s interval transit time, designated ∆t, measured in microseconds/ft or microseconds/m (∆t is the reciprocal of sonic velocity * 1000). It is a measure of a formation’s capacity to transmit sound waves. Geologically this capacity typically varies with lithology and rock texture, notably porosity. Once again, because most subsurface evaporites have extremely low porosities and are often relatively pure, the sonic log can be used to reliably identify evaporites, once an initial identification has been made by some other means (Table 1). The seeming precision of the figures given in Table 1 are illusory as the actual transit times in thick evaporites can be strongly influenced by compositional variation, temperature and confining pressure.

Rock salt is formed mainly composed of the mineral halite and is a lithology whose density is effectively constant with depth (Warren, 2016; Chapter 1). Since density is probably the most critical factor in determining acoustic velocity, the ∆tma of a thick halite unit tends to be relatively constant over a wide depth range. For pure halite, the interval transit time is 68 µsec/ft (14,625 ft/sec). However, many halite units contain varying levels of impurities, usually anhydrite, either as interbeds or disseminated throughout the sequence. Anhydrite has an interval transit time of 50 µsec/ft (20,000 ft/sec). The velocity variations in a binary system of halite and anhydrite are related linearly (either by weight or volume) to the densities of the various mixtures of the end member values allowing semiquantitative determinations of the purity of the units. Sonic logs are widely used in the oil industry for correlation and the construction of synthetic seismograms. When considering representative velocities in interpreting seismic lines, it must be remembered that the presence of bedded anhydrite and carbonate units within the total rock salt interval can have an appreciable effect on the average seismic velocity through a salt interval.

Basic identification conventional wireline log outputs

The gamma log ()aka as the lithology log) measures the natural or spontaneous radioactivity of a formation. In a sand-shale basin, the measured gamma values are used to infer clay contet. In an evaporite basin, the gamma log (especially the spectral gamma log) is a reliable indication of the presence or absence of potash salts.

In a classic quick-look analysis of any potential hydrocarbon reservoir, the sonic, density and neutron logs are used both individually and in combination to estimate the porosity of likely reservoir strata. These three logs are referred to as the “porosity logs”. Although they are typically used to indirectly infer porosity, they actually reflect variations in rock properties related to the passage of sound, induced gamma radiation and high energy neutron bombardment. The fact there is negligible porosity in most subsurface evaporites means the “porosity logs” in combination with each other, or with a spectral gamma log, can be used to identify evaporite mineralogies.


With conventional wireline log suites in bedded and halokinetic sequences worldwide I use following quick-look procedure to identify various major anhydrite, halite and potash units (refer to Table 1 and figure 5):

1) Tentatively define evaporite intervals as zones dominated by lowest gamma-ray values (some carbonates also show very low, but typically slightly higher gamma values). Remember potash beds encased in halite, or less often anhydrite, will have high gamma values.

2) Confirm pure anhydrite intervals (thicker than a metre -tool resolution dependent) using

a) Sonic - ∆t ≈ 50 microseconds

b) Bulk Density - Log value of 2.98 gm/cc. Anhydrite densities in log curve greater than 2.95 typically indicate anhydrite (but be aware of possible metal sulphides (pyrite, galena) and barite cement in some evaporite masses, especially in the caprock to halokinetic structures. NPHI porosities of anhydrite tend to hover at zero or on the negative side (in standard limestone porosity units)

3) Confirm pure halite intervals (thicker than a metre) using a combination of density and NPHI (neutron) logs. Halite-dominated zones show a consistent combination of bulk densities around 2.1 gm/cc, negative NPHI porosities and sonic (delta T) values around 64 - 70 µs/ft.

4) Use the caliper - a curve tracking the nominal bit size to confirm anhydrite versus halite (±potash salts). A caliper value much larger than bit-size indicates borehole washout, it is probably due to the intersection of salt, not the less-soluble anhydrite. Anhydrite beds tend to show an "in-gauge" caliper profiles and also tend to be slower-drilled units compared to halite (penetration data can be seen in a mudlog or well completion report). However, carbonate intrasalt beds can also show slow drill penetration rates and an "in-gauge" profile

5) Use the resistivity log - Salt like anhydrite has high resistance to current flow (Table 1), but, due to wash-out, salt units often shows lower apparent resistivity values, especially in the microresistivity and shallow reading curve outputs.

6) If there are high gamma (K-rich) intervals within thick halite beds or high-density values (>2.8gm/cc) adjacent to anhydrite consider these intervals to be possible zones with elevated levels of salts such as sylvite, carnallite or polyhalite.

7) Zones of very-low apparent bulk densities (<1.4 gm/cc) and low gamma values within a thick halite may indicate beds dominated by a non-potash evaporite minerals such as bischofite.


8) An overlaid combination of a density and a gamma log in the same track can also useful in separating what are often two co-associated high density subsurface sulphate salts, namely anhydrite and polyhalite (Figure 6). Polyhalite is a potash salt with elevated density (2.8 gm/cc) and elevated gamma values. Anhydrite has an even higher density but lacks potash and hence exhibits relatively low Gr values. This diffence in potash response in what are both characteristically high-density minerals allows for their differentiation.

9) A lack of porosity and the near linear response of the spectral gamma log to mineral proportions means GR outputs when tied to assay values can be reliably used to infer K2O ore grades (Figure 7).

In summary, most nonporous, thick relatively monomineralic evaporite units are readily identified using wireline logs, and often the proportions of minerals can be reliably determined using relevant crossplots. However, any mineralogical interpretation based on a well log outputs is just that, an interpretation, and whenever possible should be checked against rock evidence such as chips or core. When confirming a log suite interpretation of an evaporite interval, you should keep in mind that chips composed of the more soluble evaporite minerals are often completely dissolved in the drilling mud before they reach the shale shaker. In this case, the wireline logs can give a better indication of actual mineralogy than the mud chips.


References

Alger, R.P. and Crain, E.R., 1966. Defining evaporite deposits with electrical well logs. In: L.L. Raymer, W.R. Hoyle and M.P. Tixier (Editors), Second Syposium on Salt. North Ohio Geol. Soc. , pp. 116-130.

Biehl, B.C., Reuning, L., Strozyk, F. and Kukla, P.A., 2014. Origin and deformation of intra-salt sulphate layers: an example from the Dutch Zechstein (Late Permian). International Journal of Earth Sciences, 103(3): 697-712.

Crain, E.R., 2010. Potash redux. InSite CWLS Magazibe, 29(2): 17-26.

Crain, E.R. and Anderson, W.B., 1966. Quantitative log evaluation of the Prairie Evaporite Formation in Saskatchewan. J. Can. Pet. Technol., 5(3): 145-152.

Ellis, D.V. and Singer, J.M., 2007. Well logging for Earth Scientists (Second Edition). Elsevier.

Fuzesy, L.M., 1982. Petrology of potash ore in the Esterhazy Member of the Middle Devonian Prairie Evaporite in southeastern Saskatchewan. NDGS/SKGS-AAPG; Fourth International Williston Basin Symposium, October 5-7, 1982: 67-73.

Hill, D.G., 1993. Multiple log potash assay. Journal of Applied Geophysics, 30(4): 281-295.

Kruger, N., 2014. The Potash Members of the Prairie Formation in North Dakota. Report of Investigation - North Dakopta Geologica; Survey No. 113.

Nelson, P.H., 2007. Evaluation of potash grade with gamma-ray logs. U.S. Geological Survey Open-File Report 2007-1292, 14 p.

Nurmi, R.D., 1978. Use of well logs in evaporite sequences. In: W.E. Dean and B.C. Schreiber (Editors), Marine evaporites. . Soc. Econ. Paleontol. Mineral., Short Course notes, Tulsa, OK, pp. 144-176.

Rider, M., 1996. The geological interpretation of well-logs, second edition. Whittles Publishing, Caithness, 268 pp.

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


Salt Dissolution (5 of 5): Metals and saltflow-focused fluids

John Warren - Wednesday, February 28, 2018

 

Introduction

Most subsurface evaporites ultimately dissolve and, through their ongoing dissolution and alteration, can create conditions suitable for metal enrichment and entrapment in subsurface settings ranging from the burial diagenetic through to the metamorphic and igneous realms. This article looks at a few examples tied to halokinesis, a more comprehensive set of examples and more detailed discussion is given in Chapters 15 and 16 in Warren (2016). Because most, if not all, of any precursor salt mass that helped form these metalliferous deposits via dissolution, has gone, the resulting metal and other accumulations tend to be at or near the edges of salt basins, or in areas where most or all of the actual salts are long gone (typically via complete subsurface dissolution or metamorphic transformation, so that only breccias, weld and indicator mineral suites remain).

A lack of a direct co-occurrence with evaporite salts is perhaps why the metal-evaporite association is not recognised by some in the economic geology community. The significance of disappearing salt masses in focusing and enhancing metal precipitation, via the creation of chloride-rich and sulphate-rich brines, may not be evident without the conceptual tools needed to recognise the former presence of evaporites, post-salt halokinetic structural geometries, and meta-evaporite mineral associations.

The various ore tonnage-grade plots in Warren (2016), shows that many metal accumulations with an evaporite association tend to plot at the larger end of their respective deposit groupings.



Evaporite dissolution helps create "prepared ground."

I am not saying all large metal accumulations require evaporites or the highly-saline subsurface fluids that they can generate. Although, some recent papers do argue for a widespread role of evaporites in a Pb-Zn association (Fusswinkel et al., 2013; Wilkinson et al., 2009) and in sedimentary redbed copper deposits (Rose, 1976; Hitzman et al., 2010).

Typically, the conceptualisation of an evaporite in the economic geology literature is as a bedded evaporite and brine source (Figure 1). Likewise, this article and the relevant chapters in Warren (2016) detail a number of megagiant ore deposits where dissolving evaporite bodies have contributed in some way to a metal accumulation (Table 1). However this current article, like Warren 2016) focuses on the mechanisms and indicators tied to a halokinetic-ore association. Halokinesis is an aspect of evaporites that is not widely discussed in the field of ore deposit models.


Not all sediment-associated ore deposits are associated with evaporites. Only in those ore deposits classified as anorogenic and/or continental margin can subsurface evaporite masses can be involved in the same unusual concentration and alteration conditions that lead to the creation of metalliferous ore deposits (evaporite associations are indicated by E in Figure 2). At other times and locations hydrothermal mineral salts, especially anhydrite (CaSO4)which can supply sulphur as it dissolves, can be an integral part of the ore accumulation, but their occurrence may be unrelated to aridity. Hydrothermal anhydrite and other burial/magmatic hydrothermal salts tend to form in high salinity conditions inherent to the ore-forming environment and not necessarily to the presence of precursor evaporites; as in the formation of carbonatites (e.g. Afrikanda and Bayan Obo; Wu, 2008), or pegmatites and some IOCG deposits (hydrothermal anhydrite is indicated by HA in Figure 2). In some such hot subsurface settings the role of any nearby buried true” evaporite may be, via its dissolution or alteration, to aid in the creation of highly-saline high-temperature basinal brines (Chapter 16; Warren, 2016). According to whether the resulting brines are chloride or sulphate-rich, they can act as either enhanced metal carriers or fixers.

The role of evaporites creating metalliferous ores is two-fold; 1) In solution (halite-dominant precursor) they can act as chloride-rich metal carriers and 2) Locally, asCaSO4 beds or masses alter and disintegrate, their dissolution products, especially if trapped, can supply sulphur (mostly via bacteriogenic or thermogenic H2S). Dissolutional interfaces set up chemical interfaces that act as foci during brine mixing so manufacturing conditions suitable for precipitation of metal sulphides or native elements. As a consequence, most evaporite-associated ore systems tend to epigenetic, rather than syngenetic. Subsurface salt beds and masses are merely the solid part of a sizeable ionic recycling system, dissolved metals are another part, and zones of mixing between the two are typically sites where metal sulphides tend to gather.

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

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

Modelling

Conceptually, this evaporite-related notion of regional fluid flow in a sedimentary/metasedimentary host is somewhat different to the internal process and local mineralised halo models that dominate our understanding of those world-class ore deposits related to the interior workings of igneous systems. The latter is known as an orthomagmatic system where internal igneous processes of fractional crystallisation and liquid immiscibility largely control ore formation. Ores are deposited in an evolving framework of world-scale tectonics and magmatism across time, from Archaean greenstones to those of present-day sialic plate tectonism. Examples, where buried evaporites have been assimilated into a magma chamber, are discussed chapter 16 in Warren, 2016. Then there are the various ore deposits that are external to (paramagmatic) or unrelated to the emplacement of igneous bodies (nonmagmatic). In both cases, the mineralisation is typically part of an ongoing long-term sedimentary burial history, tied to dissolving and flowing salt masses and associated hydrothermal circulation.

Evidence for hydrothermally-induced low-moderate temperature mineralisation is often best preserved in textures in the hydrothermally altered rock matrix, typically located outside the actual ore deposit (in its hydrothermal alteration halo). From the hydrothermal fluid perspective, one should see the role of evaporites and metal sulphides as each contributing its part to a larger scale “mineral systems” paradigm; much in the same way as, in a petroleum system, the integration of concepts of source, carrier, seal and trap are fundamental requirements to understand and predict economic oil and gas accumulations.

This holistic ore systems approach is not fully encompassed in some economic geology studies that use sequence stratigraphic sedimentological approaches for ore deposit prediction in greenschist terrains (Ruffel et al. 1998; Wilkinson and Dunster, 1996). In my opinion, this approach can shift the interpretation paradigm too far into the depositional realm. The problem with classic sequence stratigraphic criteria, when trying to understand ore genesis, is that sequence stratigraphy does not handle well the concept of a mobile ephemeral subsurface salt body that climbs the stratigraphy via autochthonous and allochthonous process sets (halokinesis). As the salt flows, it dissolves and so brings with it the associated epigenetic influences of brine-driven diagenesis and metasomatism.

Current sequence stratigraphic paradigms in the economic geology realm are dominated by the assumption that the geometry of the units in the depositional system, and associated fault characteristics, are relatively static within the buried sediment prism. Yet, in terms of most sediment-hosted hypersaline ore deposits, what is most important in understanding the metal-evaporite association is the understanding of; 1) Evaporite dissolution and halokinesis, 2) Migration of subsurface fluids, 3) Creation of shallower or lateral-flow redox fronts along with, 4) Opening and closing of fault/shear focused fluid conduits, typically tied to, the coming and going of bedded and halokinetic salts. These factors, rather than primary sediment wedge geometries, are the dominant controls as the mineralising system passes from the diagenetic into the metamorphic realm.

It is interesting that in a benchmark paper, discussing and classifying the world’s ore deposits in a plate-tectonic-time framework, Groves et al., (2007) list almost all the major ore categories shown in Figure 2b as belonging to the group of “...sediment-hosted deposits of non-diagnostic or variable geodynamic setting.” Into this category, they place all stratiform to stratabound sediment-hosted deposits with variable proportions of Pb, Zn, Cu (including Zambian Copper belt, Kupferschiefer and SedEx deposits). They go on to note (p. 26) that, although there is general agreement that the majority of these various deposits formed during active crustal extension, either in intracratonic rift basins or passive margin sediment hosts, there is considerable controversy concerning their broader scale tectonic setting at the time of mineralisation and the driving force for hydrothermal fluid flow at the time of their mineralisation.

Perhaps this lack of model specificity in the varied interpretations of sediment-hosted deposits reflects the fact that one piece of significant information is missing from many ore genesis models. Namely, that the greater majority of these poorly classified sediment-hosted deposits sat atop, or adjacent to, or beneath, what were once thick evaporite sequences (Table 1). In many cases, the salt mass is long gone. It was the dissolution of these salt masses, either bedded or halokinetic-allochthonous, that focused much of the ore-fluid flow in the sedimentary-diagenetic realm. The loss of salt as the basin sediments passed from the low temperature diagenetic into the metamorphic realm, and as the metalliferous fluid flow was focused into permeable conduits about, below or above the dissolving and retreating, or flowing salt edges, is how salt-related ore deposits form.

This is why the majority of these salt-aided deposits tend to occur outside salt basins that retain substantial salt masses still in the diagenetic realm. The deposits are a response to the dissolution and flow of evaporites, or the residual seawater bitterns created in underlying and subjacent settings as the salt beds were deposited, not to the presence of actual undissolved primary evaporite masses. As we see in Proterozoic and Archaean meta-evaporites and most Precambrian evaporite associations, the original salt mass is long gone from the hosting succession, via varying combinations of halokinesis, dissolution and metasomatism (Warren, 2016, Chapter 13; Salt Matters blog, August 28, 2016).

Ore deposits of Precambrian tend to be linked to evaporite alteration products and residues and rarely preserve actual sedimentary salts (other than local remains of minor hydrothermal anhydrite). In younger Phanerozoic deposits, such as Kupferschiefer, the Atlantis II deep and Dzhezkazgan, portions of actual salt (brine source) can remain in the more deeply buried parts of the basin.

Metal sulphide precipitates are not rare or unique in the subsurface diagenetic fluid milieu, what is essential in the prediction of ore-grade levels of metal sulphide buildups is understanding where and why the metal precipitation system is focused into particular structurally-controlled positions and encompass time frame/fluid volumes sufficient to build an ore deposit.

That is, evaporite-associated ore deposits are no more than ancient subsurface hydrology-specific associations where the precipitation system was stable enough, for long enough, to allow higher, ore-grade levels of metals sulphides to accumulate from carrier brines at particularly favourable and stable chemical and temperature interfaces. As such, metal precipitation sites are part of an ore-forming process set, spread across the epigenetic and syngenetic realms (Table 1). They are part of the regional evolution of the fluid plumbing from the time of deposition, into burial, and on into the realm of metamorphic transformation. This means to understand the ore system tied an evaporite-entraining system holistically; one must integrate local ore paragenesis with various aspects of the basin-scale geology, sedimentology, sequence stratigraphy, diagenetic-metamorphic-igneous facies, fluid flow conduits and structural evolution of the evaporitic basin.

Metals with a halokinetic focus

To illustrate the importance of salt dissolution tied to halokinetic fluid focusing I have chosen two well-known deposits, one is a stratiform redbed copper association (Corocoro deposit), the other a SedEx style Pb-Zn deposit (McArthur River or HYC deposit)

Corocoro and other sandstone-hosted deposits of the Central Andes

Stratabound deposits of copper (±Ag), hosted by variably-dipping continental clastic sedimentary rocks, occur in Central Andean intermontane basins and are known to postdate compressive deformation/uplift events in the region (Flint, 1986, 1989). The deposits are relatively small with variable host-rock depositional ages and include; Negra Huanusha, central Peru (Permo-Triassic); Caleta Coloso, northern Chile (Lower Cretaceous); Corocoro, northwestern Bolivia (Oligo-Miocene); San Bartolo, northern Chile (Oligo-Miocene); and Yasyamayo, northwestern Argentina (Miocene-Pliocene).


The Corocoro area has produced the largest amount of copper in these Andean examples, something like 7.8 million tonnes of copper at a grade of 7.1% (Cox et al., 2007). The location of mineralisation is controlled by structurally-focused redox fronts in bedded sediment hosts, which abut a steeply-dipping translatent thrust fault (Figure 3). Deposits are irregular, usually elongate lenses of native metal, sulphides, and their oxidation products. Typically, deposits are hosted in alluvial fan and playa sandstones or conglomerate facies that also contain abundant gypsum and lesser halite. The undersides of some copper sheets at Corocoro even preserve mudcrack polygons and bed-parallel burrow traces (Savrda et al., 2006). Ore mimicry of mudcracks is not a feature controlled by on-for-one-replacement of organic material deposited in a sandstone; rather it is following pre-existing permeability/redox contrasts.

Corocoro deposits have been mined sporadically since they were first exploited by the local Indians, prior to the Spanish invasion in the 16th century and were largely exhausted in half a century of more intense mining operations that began in 1873 (Figure 3). Sandstone and conglomerate matrices show evidence of bleaching and leaching of the original redbed host with numerous red-greybed redox interfaces visible in the mined sequences. Ore minerals (dominantly native copper) are secondary fills within secondary intergranular pores created by the dissolution of earlier carbonate and sulphate masses and intergranular cement. Twelve grey sandstone beds, which were host to the long worked-out native copper ores, occur within a stratigraphic thickness of 60 m, in a unit known as the Ramos Member that still hosts abundant CaSO4 as gypsum (Figure 3).

Ores are stratabound, but not necessarily stratiform, and the larger masses of native copper are typically shallower and present as vein fills. Sometimes the copper pseudomorphs large orthogonal-ended aragonite prisms, which can be several centimetres across. There are two main styles of mineralization; 1) Ore minerals as a matrix to stratiform detrital silicates, typically low dipping and commonly highlighting primary sedimentary structures, such as cross stratification, 2) Ores in stacked channelized sand bodies, that show steep dips in structurally complex and folded zones with local brecciation (Figure 3). Native copper commonly fills thin laterally extensive sheets in tectonic fractures in the limbs of tight folds. Ljunggren and Meyer (1964) interpreted these folded diagenetic sheets of copper as a remobilization products precipitated during deformation of earlier matrix-pore filling copper.

Critical factors in Corocoro ore genesis include (Flint, 1989; Aliva-Salinas, 1990): 1) Stratigraphic association of evaporites, organic-rich lacustrine mudstones, clastic reservoir rocks, and orogenic, igneous provenance areas for both basin-fill sediments and metals; and 2) Intrabasinal evolution of metal-mobilising saline brines derived from the buried and dissolving lacustrine evaporites that flush volcaniclastics, volcanics and feldspathic sediments. The same saline diagenetic fluids also caused the dissolution of early, framework-supporting cement and large aragonite prisms, all now pseudomorphed by native copper. Avila-Salinas (1990) notes the presence of a salt-cored décollement and its likely tie to some of the highly saline sodium chloride brines found at depth in the vicinity of the Toledo Mine (Figure 3b)

The ore-hosting clastic horizons are consistently located in the highly gypsiferous Vetas Member of the Ramos Formation, which was deposited as redbeds in braidplains or fluviodeltaic playa margins centripetal to the edges of saline evaporitic lakes that were accumulating gypsum and halite (Figure 4; Flint, 1989). Abundant gypsum is still present in the Ramos Member as nodules and satinspar vein fills. Both are secondary evaporite textures likely implying the dissolution of previously more voluminous CaSO4 and NaCl beds and masses. Gypsum along with celestite are the most common gangue minerals associated with native copper veins in all the Corocoro deposits (Singewald and Berry, 1922). In the geological analysis of the first two decades of last century, the copper-bearing beds of the westerly-dipping series were called "vetas" and those of the easterly-dipping beds "ramos" and, as a matter of convenience, the names became attached to the rocks themselves. The term "veta" is Spanish for vein and "ramo" the Spanish for branch (native copper). The 1922 paper by Singewald and Berry noted that the veta horizons were traceable continuously for over 5 km in outcrop, but they found no apparent primary trends related to ramos outcrops (Figure 3).

Six mineralised layers of each kind were in exploited in mining during the first two decades of last century, the thicknesses of which varied from a few centimetres to 7 meters (Figure 3). Sheets and masses of native copper, called charque, were up to 600 pounds in weight, but more significant volumes of copper were extracted from vetas sandstones where copper was found as diffuse minute grains, pellets, or granular masses of the native metal. Associated with the enriched copper zones were more oxidised minerals as malachite, chrysocolla, azurite, domeykite, and chalcocite. Singewald and Berry (1922) noted gypsum and salt were the principal gangue minerals, while silver minerals were rare. The vetas sediment hosts tended to be coarser grained, often conglomeratic; whereas the ramos sediment hosts were finer-grained with copper present as smaller particles and masses.


The currently accepted interpretation of the Corocoro copper is that it formed during early diagenesis within a saline playa depositional environment, and in combination with dissolution of the adjacent bedded lacustrine evaporites (Figure 4). This bedded combination is thought to have controlled the formation, transport and precipitation of the copper ore (Flint, 1989). Playa sandstones, sealed between impervious evaporitic mudstone layers, created the plumbing for focused metalliferous fluid migration toward the basin margin. It is argued that the carbonaceous material at Corocoro was likely concentrated in the sandstones and conglomerates and not in the shalier members of the sedimentary sequence (Eugster, 1989).

The organics were considered strata-entrained as primary plant matter (e.g. spores) preferentially in the sandstones, along with later possible catagenic/hydrothermally cracked products migrating as hydrocarbons out of the basin. This created locally reducing pore environments in the aquifers wherever these reduced fluids met with somewhat more oxidising updip pore waters. This updip migration of saline reducing waters, in combination with sulphur supplied as H2S from the adjacent dissolving calcium sulphate beds and nodules, as well as from dissolving intergranular sulphate cement, precipitated copper in the newly created secondary porosity. The pore water chemistry and flow hydrology of this sandstone-hosted Cu system is thought to show many affinities with diagenetic uranium-redox precipitating systems, as defined by Shockey and Renfro (1974).

However, there is, in my mind, a possible anomaly in this model, which assumes organics were deposited in fluvial sandstones at the time of deposition. It is highly unusual to have higher plant material accumulating in large volumes in sandstone in a setting that is sufficiently arid and oxidising to precipitate ongoing interbeds of halite and gypsum. Such settings are typically too dry to allow abundant higher plant growth. Also, groundwaters that are flowing basinward through bajada sandstones in Neogene sediments of the Andes are ephemeral or too oxidising to facilitate the long-term reducing conditions needed to preserve significant volumes of high plant remains in the sandstone aquifers.

What is also interesting in this sedimentological/diagenetic model of Tertiary age cupriferous redbeds deposits in the Andes, centred on Corocoro, but not considered in any detail in the published literature base, is the question..., What controlled the folding, and the associated brecciation and perhaps even subsurface brine interfaces responsible for the Cu precipitation? All the stratabound Bolivian Cu deposits accumulated in sediment hosts that were deposited in fault-bound intermontane groundwater sumps. All are located in hydrologic lows in the crustal shortening tectonic scenario that typifies the Tertiary history of the Andes.

The variable ages of the host sediments and the predominance of evaporite indicators including gypsum in outcrop (often as diagenetic residues, not primary, features in the fluvial hosts) and all intimately tied to the Corocoro ore forces the question...., “was the fluid focusing driving the Cu precipitation a response to compression-driven halokinesis in an evolving salt-lubricated thrust belt?” Did this on-ground scenario occur in a halokinetic hydrology, that was possibly related to a combination of thrust-driven telogenesis, redox setup, evaporite dissolution and aquifer focusing of brines with dissolution aiding local slumping? This, along with associated strike-slip prisms, could better explain the stability of redox interfaces in sandstone aquifers across timeframes needed to accumulate significant native copper volumes. After all, most of the ore textures are passive precipitates, mainly in pre-existing porosity. If so, perhaps these deposits are not a variation on a roll-front uranium theme, which is predicated on dispersed primary organic material in the host sandstones (Shockey and Renfro, 1974).


When one plots the position of Corocoro and other redbed copper across the region, the 1000-lb gorilla that has been standing in the corner of the room for the past century becomes obvious. The Corocoro redbed copper deposit is located on a salt-cored fault system linked across less than a kilometre to an outcropping gypsum-capped remnant of a salt diapir which crosscuts the anticlinal axis of a saline redbed/greybed Corocoro sequence and ties to the saline decollement of the Corocoro Fault (Figure 5). The same tie to salt-cored decollement and diapir proximity is true of other nearby redbed copper deposits to the south-southeast, such as Veta Verde and Callapa. It is highly likely that the saline fluid interfaces forming the redbed Cu deposits of Corocoro, Veta Verde and Callapa were halokinetically focused. A similar-salt lubricated set of thrusts and strike-slip faults typifies halokinetic anticline outcrops in Central Iran.

It is highly likely that much of the structuration that is controlling Corocoro ore positioning is a response to salt flow related uplift, brine conduits and fracture creation. Metal precipitation occurred at redox interfaces induced and controlled by regional salt-lubricated compressional tectonism, and the associated salt-structuration has driven the brine-interface redox hydrology.

Work by Rutland (1966) did make an observation that the Corocoro ore deposits are related to an unconformity between the Ramos and Vetas Formations. Previously, the unconformity was interpreted as directly due to the outcrop of the Corocoro Fault. He noted that the fault and the unconformity were one and the same. In the 1960s there was no notion of a salt weld but it was nonetheless a highly astute observation by Roy Rutland. He went on to note a similar unconformity is tied to the growth of the Chuquichambi salt diapir, some 100 km southeast of Corocoro. Unfortunately, the halokinetic implications of Rutland's work were not considered 20 years later in Flint's key 1989, paper inferring a mostly clastic sedimentological origin for the Corocoro and other similar SSC deposits.

A possible halokinetic/weld association also leads to the question... Were the salt lakes, that are considered an integral part of the depositional and saline ore-precipitation systems at Corocoro by Flint, also a response to dissolution of the same nearby diapiric structures, when they were active in the mid to late Tertiary? This tie, between diapir/weld brines sourced in the drainage hinterland and bedded evaporite - lacustrine mud interbeds accumulating in the groundwater outflow sumps, is the case with groundwater inflow for the Salar de Atacama infill, as it is in other Quaternary salt lakes in the region. The are many diapir remnants across the Andes region. It seems that the Corocoro style of Cu mineralisation is perhaps another example of suprasalt redox focusing in a halokinetic setting.

Whether the halokinetic scenario, or the currently accepted non-halokinetic bedded arid-lacustrine evaporite scenario, explains the Cu mineralisation Corocoro is yet to be tested. But in terms of future copper exploration for similar deposits, it probably requires an answer. A halokinetic association offers an exploration targeting mechanism, utilising satellite imagery and aerial/gravimetric data, prior to the acquisition of on-ground land positions and geochemical surveys.

McArthur River (HYC), Ridge II and Cooley II deposits, Australia

This material on the HYC deposit will be expanded upon in an upcoming paper by Lees and Warren (in prep.). Before mining, the McArthur River (or HYC) Pb-Zn-Ag deposit, contained 227 million tonnes of 9.2% Zn, 4.1% Pb, 0.2% Cu and 41 ppm Ag (Logan et al., 1990; Pirajno, and Bagas, 2008). The deposit is hosted in the HYC Pyritic Shale member and lies adjacent to the Emu Fault in the McArthur Basin and adjacent to what are currently sub-economic base metal deposits in the Emu Fault zone known as the Cooley II and the Ridge II deposits (Figure 6a). Across all these deposits, major ore sulphides are pyrite, sphalerite and galena, with lesser chalcopyrite, arsenopyrite and marcasite. The mineralised region has an area of two km2 and averages 55 m in thickness (Figure 6b). It is elongated parallel to the major Emu growth Fault, which lies 1.5 km to the east, but is separated from the main ore mass by carbonate breccias of the Cooley Dolostone Member (Figure 6a-d).


The sequence at McArthur River comprises dolomites of the Emmerugga Dolostone (with the Mara Dolostone and Mitchell Yard members), overlain by the Teena Dolostone with abundant aragonite splays indicative of a normal-marine tropical Proterozoic carbonate. Overlying the Teena Dolostone in the vicinity of the HYC deposit is the somewhat deeper water Barney Creek Formation and its equivalents, containing the W-Fold Shale member, while the ore is hosted in carbonaceous shales, with multiple lenses of fine-grained galena-sphalerite-pyrite, separated by inter-ore sedimentary breccias (Large et al., 1998). This unit contains numerous sedimentary features indicative of a deeper-water anoxic setting. For example, comparison with d13C values from isolated kerogen in the HYC laminites confirms that n-alkanes in Bitumen II are indigenous to HYC, indicating that the deposit formed under euxinic conditions. This supports a generally-held model for Sedex deposits the region, whereby lead and zinc reacted in a stratified water column with sulphide produced by bacterial sulphate reduction (Holman et al., 2014).

The ore-hosting organic-rich 1,643-Ma HYC Pyritic Shale Member of the Barney Creek Formation is much thicker in the HYC sub-basin than elsewhere in the Batten Trough Fault Zone (e.g., Glyde River Basin) and consists mainly of dolomitic carbonaceous siltstones (Figure 7; Davidson and Dashlooty, 1993; Bull 1998). I would argue this thickening reflects a combination of long-term local basinfloor subsidence, related to salt withdrawal, and brine stratification due to ongoing salt dissolution and focused outflow. Indicators of former salt allochthon tiers are widespread in the vicinity of the HYC deposit, but are absent in the Glyde River Basin.


Breccias in and around HYC

In the HYC mine area, the ore interval is overlain by the HYC pyritic shale member and made up of pyritic bituminous and dolomitic shales and polymict breccias (Figure 7). Importantly, when contacts are walked out in outcrop, the polymict breccias are significantly transgressive to bedding, while drilled intersections in the vicinity of the HYC deposit and in the mine itself show the breccias are stratabound. Another interesting feature of these breccias is that they can contain mineralised clasts. More broadly, a variety of sedimentary breccias occur throughout the Barney Creek Formation stratigraphy, especially along the eastern margin of the HYC half graben and tend to pass updip into the breccias of the Cooley Dolostone (Figure 6a).

Williams 1976, defined three breccia types (I, II and III) in the HYC area. Type I breccia beds occur in the lower half of the HYC Pyritic Shale Member and contain clasts characteristic of lithologies in formations of the McArthur Group below the Barney Creek Formation (Table 2). In the northern end of the sub-basin, the breccias are of a chaotic nature with no sorting and minor grading of clasts (Figure 6b). The underlying shale beds are frequently contorted and squeezed between the breccia fragments, which reach a maximum size of approximately 10 m. Toward the south, the thickness and maximum clast size of individual breccia beds decrease (Figure 6b). All breccia units are thickest adjacent to the Emu Fault Zone and likely record sediment sinks controlled by rapid fault-controlled basin subsidence during Barney Creek time. Inter-ore breccias amalgamate and thicken to the north-north-east of HYC, and occupy a position toward the foot of what is interpreted as a more substantial breccia lens, dominated by sediment gravity flow deposits (Figure 6d; Logan et al., 2001).


In a subsequent study, Ireland et al. (2004a) identified four distinct sedimentary breccia styles within Type I breccias: framework-supported polymictic boulder breccia; matrix-supported pebble breccia; and gravel-rich and sand-rich graded turbidite beds (Table 2). The boulder breccias can be weakly reverse-graded and show rapid lateral transition into the other facies, all of which are interpreted as more distal manifestations of the same sedimentary events. The flow geometry and relationships between these breccia styles are interpreted by Ireland et al. (2004a) to reflect mass-flow initiation as clast-rich debris flows, with transformation via the elutriation of fines into a subsequent turbulent flow from which the turbidite and matrix-supported breccia facies were deposited.

All the Type 1 mass-flow facies contain clasts of the common and minor components of the in-situ laminated base-metal mineralised siltstone. Texturally these clasts are identical to their in-situ counterparts and are distinct from other sulphidic clasts that are of unequivocal replacement origin. In the boulder breccias, intraclasts may be the dominant clast type, and the matrix may contain abundant fine-grained sphalerite and pyrite. Dark-coloured sphaleritic and pyritic breccia matrices are distinct from pale carbonate-siliciclastic matrices, are associated with a high abundance of sulphidic clasts, and systematically occupy the lower parts of breccia units. Consequently, clasts that resemble in-situ ore facies are confirmed as genuine intraclasts incorporated into erosive mass flows before complete consolidation. Disaggregation and assimilation of sulphidic sediment in the flow contributed to the sulphide component of the dark breccia matrices. The presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows. That is, the presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows (Ireland et al., 2004a).

Type II breccia beds occur throughout the HYC Pyritic Shale Member but are most common in the upper half of the Member. Clasts are predominantly grey dololutite which occasionally contain radiating clusters of acicular crystal pseudomorphs (“coxcos”) indicative of tropical Proterozoic shelf carbonates. The clasts are similar to lithologies in the Emmerugga and Teena Dolomites and are considered to have been derived from these formations. A characteristic of this breccia type, which differentiates it from Type I and III breccias is the absence of green and red clasts, signifying that clasts in Type II breccias were not derived from the Tooganinie or lower formations, but mostly derived by erosion and collapsed of updip shallow-water cemented shelf carbonate layers. Type II breccias lack the well-developed grading seen in Type I breccias. Isopach maps (Figure 6c) and maximum clast-size plots of individual breccia beds show a close correlation and indicate the type II breccias dominate in the southeast of the HYC subbasin.

Type III breccia beds are confined to the uppermost breccia unit of the HYC Pyritic Shale Member in the HYC sub-basin and are equivalent to the Upper Breccia of Murray (1975). This unit consists exclusively of Type III breccias with the exception of several shale beds near the base. The top of the Upper Breccia is not exposed in the sub-basin, and the unit reaches a maximum known thickness of 210 m. Clasts within the breccias are completely chaotic, and there is no recognisable grading or sorting. Clasts range in size from a few millimetres up to several tens of metres. The fragment lithologies are identical to those in the Type I breccias with the notable exception that they also contain clasts of sandstone, quartzite and potash-metasomatized quartz dolerite—lithologies that are characteristic of the underlying Masterton Formation. The fragments are therefore considered to be derived from the McArthur Group (below the Barney Creek Formation) and the Masterton Formation. According to Walker et al. (1977), the most likely source of the clasts from the Masterton Formation is erosion uplifts and horsts in the Emu Fault Zone. But the same authors also state the exact source area and the direction of movement of the clasts could not be identified. In my opinion, Type III breccias are salt-ablation derived and so contain a variety of clasts lithologies plucked by the rising salt as it rose toward the surface to feed an at-seafloor allochthon.

More broadly, breccias of the updip Cooley Dolostone member, that interfinger and also overlie the HYC deposit (Figure 6a) are usually regarded as part of the Barney Creek Formation. The Cooley Dolostone is interpreted, historically, as a talus slope breccia (Walker et al. 1977, Logan 1979), containing clasts eroded from the Teena and Emmerugga Dolostones. Hinman (1995) regarded the Cooley Dolostone as a tectonic breccia, formed along reverse faults within the steep to overturned, brittle dolomitic lithologies of Teena, Mitchell Yard and Mara Dolostones(members of the Emmerugga Dolostone) as they were overthrust against and over Barney Creek Formation lithologies. Perkins & Bell (1998) interpret the Cooley Dolostone as an in situ alteration body, contiguous with, and derived from, the HYC sequence, rather than being separated from it by a thrust fault. I interpret much of the Cooley as a salt allochthon breccia derived from a salt-cored basin edge fault system, now evolved into a salt weld (Table 2).

Brine haloes and mineralisation

Regional-scale potassic alteration of Tawallah Group dolerites and sediments were documented by Cooke et al. (1998), Davidson (1998, 1999). These authors describe fluids responsible for this alteration as oxidised, low-temperature (100˚C), saline (> 20wt % NaCl equiv), Na-K-Ca-Mg-rich brines, and argue that the high salinities and the presence of hydrocarbons are consistent with brine derivation from nearby evaporitic carbonates during diagenesis.

I suggest that saline fluids feeding these haloes came not from the dissolution of evaporites in adjacent bedded carbonate hosts, but from the decay of former fault-fed thick salt allochthon tongues in positions that now are indicated by salt allochthon breccias. These breccias tie back to what were salt-lubricated fault and salt welds. The presence of salt and diagenetic haloes in these features focused tectonic movement and fluid supply in both initial extensional and subsequent compressional stages. As such, this interpretation supports a salt dissolution origin of the brine origins proposed by both Logan (1979) and Hinman (1995). The difference with their interpretations is that I envisage the brine being derived during salt flow emplacement and dissolution, tied to focused fault conduits in a mobile, suprasalt fault complex, atop or adjacent to the now-dissolved flowing and tiered salt mass. I do not think the nearby platform carbonates (with coxcos and smooth-walled cherts) ever contained significant volumes of primary evaporites.

Worldwide and across deep time, most halokinetic basinwide evaporite associations are typified by an initial extensional and loaded set of diapirs evolving into salt-cored fault welds, with subsequent reactivation of these features in compression (Warren, 2016; Chapter 6). Such a framework typifies long-term salt tectonics with inherently changing structural foci across most Phanerozoic halokinetic salt realms, as in the North Sea, the Persian (Arabian) Gulf and most circum-Atlantic salt basins. It is indicative of continental plate-edge evaporites caught up in the Wilson cycle (Warren, 2010).

Near the HYC deposit, Mn-enrichment, particularly of dolomite and ankerite in the W-fold Shale beneath the ore zone, is considered to be related to exhalation of Mn-bearing brines, associated with rifting and basin deepening, before the onset of zinc-lead mineralisation (Large et al. 1998). This too, is consistent with the salt-focused mineralisation hydrology of diagenetic ferroan and Mn-bearing hydrologies of the modern Red Sea halokinetic deeps (Schmidt et al., 2015) and the Danakhil depression in the Quaternary, when it was a marine-fed saline system (Bonatti et al., 1972).

Ridge and Cooley deposits

In the area to the east of to McArthur River HYC basin, a number of currently sub-economic Zn-Pb-Cu deposits occur, typified by the nearby Ridge and Cooley deposits (Figure 6a; Walker et al. 1977; Williams 1978). Both are similar to the Coxco deposit, being described as MVT deposits mainly hosted by dolomitic breccias, but with minor, shale-hosted concordant mineralisation in the Ridge II deposit (Figure 8; Williams 1978). Likewise, the Coxco deposit contains several million tonnes at 2.5% Zn and 0.5% Pb, in coarse-grained, stratabound galena-sphalerite-pyrite-marcasite, hosted by dolomitic breccias containing clasts of the Mara Dolostone Member, Reward Dolostone, and the Lynott Formation of the McArthur Group, within the Emu Fault Zone (Walker et al. 1977, Walker et al. 1983). Mineralisation comprises veins, “karst” and dissolution breccia fill likened to Mississippi Valley Type (MVT) mineralisation (Walker et al. 1977).

According to Williams (1978), the Emmerugga Dolostone hosts the discordant mineralisation of Cooley II deposit, while Cooley Dolostone breccias contain the Ridge II deposit (Figure 8). The Emmerugga Dolostone at Cooley II consists of massive to laminated dolostone and contains carbonaceous matter, stromatolites, oncolites, and ooids, indicating that it was deposited in a shallow-water normal-marine environment with high biologic productivity. Similarly, the Cooley Dolostone host at Ridge II is a breccia composed of randomly oriented dolostone clasts varying in diameter from a few millimetres up to several tens of metres. Some clasts have near-identical lithologies to those comprising the Emmerugga Dolostone, whereas others contain coxcos and were likely derived from the fragmentation of Teena Dolostone. The Cooley Dolostone breccia contains little depositional matrix. Clast boundaries are marked by sudden changes in features such as dolostone type and bedding-core angles, indicating that the breccia was mostly clast-supported at the time of formation. Most interestingly, drilling in the vicinity of the deposit (DDHR210) intersected a large clast of “out of sequence” dolerite (Figure 8a). Similar large salt-buoyed clasts (up to 100’s meters across) composed of Eocene dolerite occur in the salt allochthon breccias at Kuh-e-Namak-Qom (Salty Matters blog, March 10, 2015).


Two major phases of crosscutting brecciation in the area are recognised by Williams (1978) in drill core samples of discordant mineralisation from both the Emmerugga and Cooley Dolostone hosts. First generation breccias, formed during the earlier phase of brecciation, consist of angular clasts of dolostone (< 1 mm to at least 1 m in diameter) in a dark colored matrix of tiny ( < 1 µm to 20µm) anhedral dolomite grains, disseminated euhedral pyrite crystals (<50 µm in diameter) and reddish brown carbonaceous matter). The identical nature of the first generation breccias in both the Emmerugga and Cooley Dolostone hosts suggests that brecciation occurred simultaneously in both, via the same mechanism (Williams, 1978). At the time this interpretation was made, there was no “data” (paradigm) available to determine whether the brecciation in the Cooley Dolostone occurred in situ or whether it took place in the dolostone before its removal from the Western Fault Block. Today, we would likely interpret these features as reworked salt ablation breccias on the deep seafloor with infiltrated suspension clays and early-diagenetic pyrite.

Second generation breccias, formed during a later phase of brecciation, consist of angular clasts of first-generation breccias (< 1 mm to at least 10 cm in diameter) in a matrix of either veins filled with sulphide minerals and dolomite, or fine-grained (10 µm to 100 µm in diameter) anhedral dolomite grains, disseminated to massive sulfide minerals, small (on the average 500 µm x 20 µm) interlocking laths of barite or dolomite pseudomorphs after barite, and brown carbonaceous matter (Williams, 1978). Second generation breccias, although coincident with the first generation breccias, are less widespread than the earlier breccias. Again, according to Williams (op. cit.), the similarity of the second generation breccias in both the Emmerugga and Cooley Dolostones suggests a common origin. Again, they concluded there was no “data” (paradigm) available to establish the time of this brecciation relative to the deposition of the Cooley Dolostone. I would argue these “second generation” breccias represent a less distally reworked salt ablation breccia, possibly with interspace anhydrite and gypsum at the time they formed. These calcium sulphate phases facilitated the shallow subsurface emplacement of metal sulphides via bacterial or thermochemical sulphate reduction, in a way not too dissimilar to the mechanisms emplacing Pb-Zn at Cadjebut or Bou Grine ores in Tunisia (Warren and Kempton et al., 1997; Warren 2016; Chapter 15).

Allochthon Interpretation

The origin of the HYC deposit and adjacent subeconomic mineralised accumulations is still somewhat controversial and equivocal (Figure 6a; Ireland et al. 2004a,b; Perkins and Bell, 1998; Logan, 1979; Walker et al., 1977). Large et al. 1998 summarised the alternative models: 1) a sedimentary-exhalative (‘sedex’) model was proposed by Croxford 1968 and Large et al. 1998; while, 2) a syndiagenetic subsurface replacement model was introduced by Williams 1978; Williams & Logan 1986; Hinman 1995 and Eldridge et al. 1993, the latter based on sulphur isotopes. In my opinion, a third factor, namely a now-dissolved salt allochthon system, should be considered in interpretations of ore genesis and associated breccias. I interpret ore-hosting laminites of HYC deposit as DHAL laminites, and the Ridge II and Cooley II were hosted in updip regions once dominated by salt tongues and salt ablation breccias within a fault-fed salt allochthon complex surrounded by updip normal-marine shoal-water platform carbonates (Figure 9).

That is, all three deposits are related to the ongoing and time-transgressive dissolution of shallow halokinetic salt tiers. The salt tongues periodically shed mass flow deposits, triggered by seafloor instability created by the interactions of salt flow, salt withdrawal and the dynamic nature of salt and fault welds. In my opinion, the lack of equivalent breccias, DHAL laminites and halo evidence in otherwise similar deepwater sediment in Barney Creek Formation in the Glyde River Basin, some 80 km to the south-east of HYC, is why this basin lacks economic levels of base metal mineralisation (Figure 7).


Assuming that the first and second generation breccias in Type 1 and III breccias in all of the stratigraphically discordant deposits (allochthon and weld breccia), first defined by Walker et al., 1977 (Table 2) had shared salty origins, the wider distribution of the first generation breccias suggests that they formed via seafloor reworking processes acting across the whole region as a rim to discordant mineralisation (Williams 1978). Therefore, Williams (op cit.) argued geologically reasonable causes of the brecciation in the Cooley Dolostone include; movement on the Western and Emu faults, slumping of debris off the Western Fault Block, and stratal collapse due to the dissolution of evaporite minerals. I would argue for all of the above, but add that the whole Cooley Dolostone breccia system at the time the first generation breccias formed was a massive salt-flow fault-feeder system that was salt-allochthon cored and salt-lubricated. Situated at and just below the deep seafloor, salt tongue dissolution created salt-ablation breccias, while the halokinetic-induced seafloor instability instigated periodic mass flows into a metalliferous brine lake; as occurs today in the modern Red Sea deeps, the Orca basin in the Gulf of Mexico and the various brine lakes (DHAL's) of the Mediterranean Ridges (Table 2).

Breccia textures in a halokinetic salt ablation system are always two stage (Warren, 2016); the first stage of brecciation occurs as the salt tongue is inflated and spreading over the surrounds, even as its edges dissolve into ablation breccias reworked by further salt tongue movements and accumulations of contemporary salt-carapace materials (Figure 9). This first stage is typified by mass wasting piles related to the debris rims accumulating about the salt tongue edges, as debris slides downslope across the top of a continuously resupplied salt mass. The friction along the underside of the expanding salt sheets drives overturn, contortion, and brecciation of the underlying deep seafloor bed, this ultimately creates subsalt thrust overfolds (known as gumbo zones beneath the salt allochthons of the Gulf of Mexico). The second stage of brecciation is related to the dissolution of the salt itself once the salt supply is cut off by salt withdrawal and overburden touchdown.

Because allochthons are set up in the expansion stage of salt movement across the seafloor, Stage 1 breccias tend to be more widespread at the landsurface than stage 2 breccias. Stage 2 breccias form once the mother salt supply to the salt tongue or tier is cut off, the salt tongue then dissolves and final brecciation occurs, often with significant roof collapse features in any overburden layers. Similar two-stage allochthon breccias outcrop and subcrop in salt namakier provinces across Iran (Warren 2016, Chapter 7). However, unlike Iran the HYC laminites and associated breccias accumulated in a local deeper marine anoxic sump within a dominant subaqueous normal-marine carbonate shelf setting. There are also partial analogies with salt-cored Jurassic shelf carbonates and allochthon breccias in the paleo Gulf of Mexico, or the Cretaceous mineralised and ferruginised shelf-to-slope halokinetic-cored depositional system that now outcrops in the Domes Region of North Africa (Warren, 2008; Mohr et al. 2007).

Based on the sedimentology of the HYC ore host (Figure 9), I conclude that the HYC deposit accumulated as classic DHAL deposit in a salt allochthon-floored sump. Initial ore accumulation took place as metalliferous laminites in a local salt withdrawal basin. The anoxic brine-filled DHAL sump sat atop a deflating salt allochthon sheet with one of the tiers indicted by salt dissolution breccias at the Myrtle-Mara contact.

The following observations further support this conclusion; 1) the scale and deepwater setting of the deposit, 2) the fault-bound brine-fed margin to the deposit, 3) the rapid local subsidence of the sediments in the deeper water anoxic portion that constitutes the Barney Creek Fm host (HYC Pyrite member), 4) the syndepositional nature of the inter-ore polymict mass flow breccias, 5) the presence of syndepositional barite and Mn haloes from a diagenetically imposed oxidised saline set of pore waters hosted in what were formerly normal-marine sediment pore fluids.

Salt flowing from an allochthon sheet into salt risers in the Emu-Western fault region drove fault-bound rapid subsidence that created local deeper-water anoxic brine-filled sumps in an otherwise healthy marine carbonate shelf (see Salty Matters blog, April 29, 2016, for a salt-controlled structural analogy in the Red Sea). The fault-controlled salt risers allowed brine to escape onto the seafloor at Barney Creek time and to flow across the seafloor into the large DHAL sump that is today the HYC deposit (Figure 9). With time, the salt risers evolved in salt welds and ultimately into fault welds with salt-ablation breccia textures.

The characteristic Fe-Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the fault-conduit aquifer focus to the suprasalt brine flow and the level of hypersaline brine intersections. There are also transitions into more-typical more-oxidised marine pond and pore water masses in the upper levels atop the DHAL waters and around the edge of its brine curtain.

Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree and would argue that a later DHAL mineralisation focus, during the creation of a later generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as the mother salt supply was lost (Table 2).

Williams (op. cit.) noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite, presumably brought in from an outside source, in the matrix of the second generation breccias suggest that the later breccias formed by solution collapse following the introduction of mineralizing solutions into the porous, first generation breccias. I am in complete agreement with this conclusion. In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession that we classify as the sedex brine pool stage that is forming the HYC deposit. With time and salt dissolution/source depletion, we pass to the next generation of breccias, which are linked to a fault weld, evaporite-collapse sub-economic set of MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).

In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones. Salt withdrawal from allochthon sheets emplaced below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits, as defined by thickness and mineralogical/ colour changes in the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor depression, which was the HYC DHAL sump, it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault-weld). The stratigraphic level of the withdrawal is indicated by the allochthon collapse breccia seen at the top of the Myrtle Shale.

The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride rich brines circulating in the buried sediments of the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALs, some of the same metal-bearing brines exploited the presence of fractionally dissolved interclast calcium sulphate within diapir collapse breccias. So a similar set of redox interfaces drove discordant mineralisation in second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and salt collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks breccia and peripheral Cretaceous seafloor DHAL laminites in the Bahloul Formation of Northern Africa (see Warren 2016; Chapter 15).

The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after the main episode of laminite Pb-Zn ore formation. This interpretation of both inter-ore “sedimentary” and Cooley Dolostone member breccias across the region reconciles what were seen as previously conflicting primary versus time-transgressive relationships (e.g., Williams 1978; Perkins & Bell 1988).

The characteristic Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the aquifer and the level on hypersaline brine intersections with the more typical more oxidised marine water mass and pores water at levels atop the brine lake.

Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree, and would argue that the later mineralisation focus, during the creation of the second generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as any ongoing mother salt supply was lost. Williams (op. cit.) in a discussion of the Ridge and Cooley deposits noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite in the matrix of the second generation breccias, presumably brought in via fluids with an outside source. He suggests that later breccias formed by solution collapse following the introduction of mineralising solutions into the porous, first generation breccias. I agree also with this conclusion but would also place it in the typical saline baryte ore association seen in many salt diapir provinces such as the Walton-Magnet Cove region of Nova Scotia, or the Oraparinna Diapir in the Flinders Ranges, South Australia (see Warren 2016, Chapter 7 for detail on theses and other similar baryte deposits).

In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession we classify as the sedex brine pool that is the HYC deposit, to the next generation of breccias that are linked to a fault weld, evaporite-collapse sub-economic set of smaller scale MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).

In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones were deposited. Salt withdrawal below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits defined by the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor that was the HYC DHAL sump it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault- weld) with the stratigraphic level of the withdrawal indicated by the allochthon collapse breccia at the top of the Myrtle Shale.

The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride-rich brines circulating in the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALS, some of the same metal-bearing brines exploited diapir collapse breccias and drove discordant mineralisation and second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks and Cretaceous seafloor laminites of the Bahloul Formation of Northern Africa (see Warren 2016 Chapter 15).

The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after ore formation.

References

Avila-Salinas, W., 1990. Origin of the Copper Ores at Corocoro, Bolivia. In: L. Fontboté, G.C. Amstutz, M. Cardozo, E. Cedillo and J. Frutos (Editors), Stratabound Ore Deposits in the Andes. Special Publication No. 8 of the Society for Geology Applied to Mineral Deposits. Springer Berlin Heidelberg, pp. 659-670.

Bonatti, E., Fisher, D.E., Joensuu, O., Rydell, H.S. and Beyth, M., 1972. Iron-manganese-barium deposit from the north Afar rift (Ethiopia). Economic Geology, 67(6): 717-730.

Bull, S.W., 1998. Sedimentology of the Palaeoproterozoic Barney Creek formation in DDH BMR McArthur 2, southern McArthur basin, northern territory. Australian Journal of Earth Sciences: An International Geoscience Journal of the Geological Society of Australia, 45(1): 21-31.

Cooke, D.R., Bull, S.W., Donovan, S. and Rogers, J.R., 1998. K-metasomatism and base metal depletion in volcanic rocks from the McArthur basin, northern territory - Implications for base metal mineralization. Economic Geology, 93(8): 1237-1263.

Cox, D.P., Lindsey, D.A., Singer, D.A., Moring, B.C. and Diggles, M.F., 2007. Sediment-Hosted Copper Deposits of the World: Deposit Models and Database. USGS Open File Report 03-107, Version 1.3 (Available online at http://pubs.usgs.gov/of/2003/of03-107/.

Davidson, G.J., 1998. Alkali alteration styles and mechanisms, and their implications for a brine factory source of base metals in the rift-related McArthur Group, Australia. Australian Journal of Earth Sciences, 45(1): 33-49.

Davidson, G.J., 1999. Feldspar metasomatism along a Proterozoic rift-basin margin - "Smoke" around a base-metal "fire" (HYC deposit, Australia) or a product of background diagenesis? Geological Society of America Bulletin, 111(5): 663-673.

Davidson, G.J. and Dashlooty, S.A., 1993. The Glyde Sub-basin - A volcaniclastic-bearing pull-apart basin coeval with the McArthur River base-metal deposit, Northern Territory. Australian Journal of Earth Sciences, 40(6): 527-543.

Eldridge, C.S., Williams, N. and Walshe, J.L., 1993. Sulfur isotope variability in sediment-hosted massive sulfide deposits as determined using the ion microprobe SHRIMP: II. A study of the H.Y.C. deposit at McArthur River, Northern Territory, Australia. Economic Geology, 88(1): 1-26.

Entwistle, L.P. and Gouin, L.O., 1955. The chalcocite-ore deposits at Corocoro, Bolivia. Economic Geology, 50(6): 555-570.

Eugster, H.P., 1989. Geochemical environments of sediment-hosted Cu-Pb-Zn deposits. In: R.W. Boyle, A.C. Brown, C.W. Jefferson and E.C. Jowett (Editors), Sediment hosted stratiform copper deposits. Geological Association of Canada, Special Paper, pp. 111-126.

Flint, S., 1986. Sedimentary and diagenetic controls on red-bed ore genesis; the middle Tertiary San Bartolo copper deposit, Antofagasta Province, Chile. Economic Geology, 81(4): 761-778.

Flint, S.S., 1989. Sediment-hosted stratabound copper deposits of the Central Andes. Geological Association of Canada Special Paper, 36: 371-398.

Fusswinkel, T., Wagner, T., Wälle, M., Wenzel, T., Heinrich, C.A. and Markl, G., 2013. Fluid mixing forms basement-hosted Pb-Zn deposits: Insight from metal and halogen geochemistry of individual fluid inclusions. Geology, 41(6): 679-682.

Groves, D., I. and Bierlein, F., P. , 2007. Geodynamic settings of mineral deposit systems. Journal of the Geological Society, 164: 19-30.

Groves, D.I., Condie, K.C., Goldfarb, R.J., Hronsky, J.M.A. and Vielreicher, R.M., 2005. Secular changes in global tectonic processes and their influence on the temporal distribution of gold-bearing mineral deposits. Economic Geology, 100(2): 203-224.

Harrison, H. and Patton, B., 1995. Translation of salt sheets by basal shear. Proceedings of GCCSEPM Foundation 16th Annual Research Conference, Salt Sediment and Hydrocarbons, Dec 3-6, 1995: 99-107.

Hinman, M., 1995. Base metal mineralisation at McArthur River: structure and kinematics of the HYC-Cooley zone at McArthur River. Australian Geological Survey Organisation, Record,1995/5.

Hitzman, M.W., Selley, D. and Bull, S., 2010. Formation of Sedimentary Rock-Hosted Stratiform Copper Deposits through Earth History. Economic Geology, 105(3): 627-639.

Holman, A.I., Grice, K., Jaraula, C.M.B. and Schimmelmann, A., 2014. Bitumen II from the Paleoproterozoic Here’s Your Chance Pb/Zn/Ag deposit: Implications for the analysis of depositional environment and thermal maturity of hydrothermally-altered sediments. Geochimica et Cosmochimica Acta, 139: 98-109.

Ireland, T., Bull, S.W. and Large, R.R., 2004a. Mass flow sedimentology within the HYC Zn-Pb-Ag deposit, Northern Territory, Australia: evidence for syn-sedimentary ore genesis. Mineralium Deposita, 39(2): 143-158.

Ireland, T., Large, R.R., McGoldrick, P. and Blake, M., 2004b. Spatial distribution patterns of sulfur isotopes, nodular carbonate, and ore textures in the McArthur River (HYC) Zn-Pb-Ag deposit, northern territory, Australia. Economic Geology, 99(8): 1687-1709.

Large, R.R., Bull, S.W., Cooke, D.R. and McGoldrick, P.J., 1998. A genetic model for the HYC deposit, Australia: Based on regional sedimentology, geochemistry, and sulfide-sediment relationships. Economic Geology, 93(8): 1345-1368.

Ljunggren, P. and Meyer, H.C., 1964. The copper mineralization in the Corocoro basin, Bolivia. Economic Geology, 59: 110-125.

Logan, G.A., Hinman, M.C., Walter, M.R. and Summons, R.E., 2001. Biogeochemistry of the 1640 Ma McArthur River (HYC) lead–zinc ore and host sediments, Northern Territory, Australia. Geochimica Cosmochimica Acta, 65: 2317-2336.

Logan, R.G., 1979. The Geology and mineralogical zoning of the HYC Ag-Pb-Zn deposit, McArthur River, NT Masters Thesis, Australian National University, Cannberra, Australia.

Logan, R.G., Murray, W.J. and Williams, N., 1990. HYC silver-lead-zinc deposit, McArthur River. In: F.E. Hughes (Editor), Geology of the mineral deposits of Australia and Papua New

Guinea. Monograph Series - Australasian Institute of Mining and Metallurgy, pp. 907-911.

McConachie, B.A. and Dunster, J.N., 1998. Regional stratigraphic correlations and stratiform sediment-hosted base-metal mineralisation in the northern Mt Isa Basin. Australian Journal of Earth Sciences: An International Geoscience Journal of the Geological Society of Australia, 45(1): 83-88.

Meyer, C., 1988. Ore deposits as guides to geologic history of the Earth. In: G.W. Wetherill and et al. (Editors), Annual review of earth and planetary sciences. Vol. 16. Annual Reviews Inc., pp. 147-171.

Mohr, M., Warren, J.K., Kukla, P.A., Urai, J.L. and Irmen, A., 2007. Subsurface seismic record of salt glaciers in an extensional intracontinental setting (Late Triassic of northwestern Germany). Geology, 35(11): 963-966.

Murray, W.J., 1975. McArthur River HYC lead-zinc-silver and related deposits. In: C.L. Knight (Editor), Economic geology of Australia and Papua New Guinea—metals. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 329-338.

Perkins, W.G. and Bell, T.H., 1998. Stratiform replacement lead-zinc deposits: A comparison between Mount Isa, Hilton, and McArthur River. Economic Geology, 93(8): 1090-1212.

Pirajno, F. and Bagas, L., 2008. A review of Australia's Proterozoic mineral systems and genetic models. Precambrian Research, 166(1-4): 54-80.

Rose, A.W., 1976. The effect of cuprous chloride complexes in the origin of red-bed copper and related deposits. Economic Geology, 71(6): 1036-1048.

Ruffell, A.H., Moles, N.R. and Parnell, J., 1998. Characterisation and prediction of sediment-hosted ore deposits using sequence stratigraphy. Ore Geology Reviews, 12(4): 207-223.

Rutland, R.W.R.R., 1966. An unconformity in the Corocoro basin, Bolivia, and its relation to the copper mineralization. Economic Geology, 61: 962-964.

Savrda, C.E., Cook, R.B. and Petrov, A., 2006. Trace Fossil Preservation by Native Copper, Corocoro, Bolivia. Rocks & Minerals, 81(5): 362-363.

Schmidt, M., Al-Farawati, R. and Botz, R., 2015. Geochemical classification of brine-filled Red Sea Deeps. In: N.M.A. Rasul and I.C.F. Stewart (Editors), The Red Sea. Springer, Berlin, pp. 219-233.

Shockey, P.N. and Renfro, A.R., 1974. Copper-silver solution fronts at Paoli, Oklahoma. Economic Geology, 69: 266-268.

Singewald, J.T. and Berry, E.W., 1922. The geology of the Corocoro copper district of Bolivia. Johns Hopkins University studies in geology -- No. 1.. 117 p.

Walker, R.N., Gulson, B. and Smith, J., 1983. The Coxco deposit - a Proterozoic Mississippi Valley-type deposit in the McArthur River district, Northern Territory, Australia. Economic Geology, 78(2): 214 - 249.

Walker, R.N., Logan, R.G. and Binnekamp, J.G., 1977. Recent geological advances concerning the H.Y.C. and associated deposits, McArthur river, N.Y. Journal of the Geological Society of Australia, 24(7-8): 365-380.

Warren, J.K., 2000. Evaporites, brines and base metals: low-temperature ore emplacement controlled by evaporite diagenesis. Australian Journal of Earth Sciences, 47(2): 179-208.

Warren, J.K., 2010. Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits. Earth-Science Reviews, 98(3-4): 217-268.

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

Warren, J.K. and Kempton, R.H., 1997. Evaporite Sedimentology and the Origin of Evaporite-Associated Mississippi Valley-type Sulfides in the Cadjebut Mine Area, Lennard Shelf, Canning Basin, Western Australia. In: I.P. Montanez, J.M. Gregg and K.L. Shelton (Editors), Basinwide diagenetic patterns: Integrated petrologic, geochemical, and hydrologic considerations. SEPM Special Publication, Tulsa OK, pp. 183-205.

Wilkinson, J.J., Stoffell, B., Wilkinson, C.C., Jeffries, T.E. and Appold, M.S., 2009. Anomalously Metal-Rich Fluids Form Hydrothermal Ore Deposits. Science, 323(5915): 764-767.

Williams, N., 1978. Studies of base metal sulphide deposits at McArthur River, Northern Territory, Australia I. The Cooley and Ridge Deposits. Economic Geology, 73(6): 1005 - 1035.

Williams, N. and Logan, R.G., 1986. Geology and evolution of the H.Y.C. stratiform Pb-Zn orebodies, Australia. Stanford Univ.Pub. Geol. Sci., 20: 57-60.

Wu, C., 2008. Bayan Obo Controversy: Carbonatites versus Iron Oxide-Cu-Au-(REE-U). Resource Geology, 58(4): 348-354.

 

Salt Dissolution (4 of 5): Anthropogenically-enhanced geohazards

John Warren - Thursday, November 30, 2017

 

Introduction

As we saw in the previous article the dissolution and collapse of nearsurface and at-surface salt is a natural and ongoing process. When salt bodies experiencing natural dissolution and alteration are penetrated by drilling or parts of the salt mass are extracted in a poorly supervised fashion, the resulting disturbance can speed up natural solution and collapse, sometimes with unexpected environmental consequences (Figure 1; Table 1). To minimise the likelihood of unexpected environmental consequences tied to enhanced rates of salt dissolution in the vicinity of engineered structures the same rule applies as is applied to safe salt mining, namely "Stay in the salt" (Warren, 2016) In this context, here is a quote from Zuber et al. (2000) in a paper dealing with flooding and collapse events in Polish salt mines.


“Catastrophic inflows to salt mines, though quite frequent, are seldom described in the literature, and consequently students of mining and mine managers remain, to a high degree, ignorant in this respect. Contrary to common opinion, inflows are seldom caused by unavoidable forces of nature. Though some errors were unavoidable in the past, modern geophysical methods are, most probably, quite sufficient to solve the majority of problems (e.g., to determine a close presence of the salt boundary). Detailed study of the recent catastrophic floods, which happened in Polish salt mines, shows that they usually occur, or have strong negative impacts, due to human errors. Most probably similar human errors caused catastrophic inflows to salt mines in other countries. It seems that a knowledge of the real history of catastrophes, better education of mine engineers and the application of modern geophysical methods could lead to the reduction of floods in salt mines.”

This article, the 4th of 5 in Salty Matters, focuses on the anthropogenic enhancement of salt dissolution, however, it discusses only a few of the many documented examples of problems that result from enhanced dissolution brought about by human activity. For a more comprehensive documentation of relevant case histories and an expanded discussion that includes mine collapse, brinefield subsidence and collapse and industrial accidents associated with salt cavity storage, the interested reader should refer to Chapters 7 and 13 in Warren (2016).


As wells as mine floods and mine losses, poorly monitored then abandoned brinefield and solution cavities can be areas with major environmental problems especially where old extraction wells are undocumented and unmonitored. Like areas of natural solution collapse; they can become zones of catastrophic ground failure. This is especially problematic if located near cities or towns (Tables 1, 2). Some of the most outstanding examples of how not to solution mine a resource and how not to control ground subsidence effects are to be found in east European countries that are still trying to deal with the environmental outcomes of being former satellite states of the Soviet Union. But caving problems have been tied also to mines and boreholes in the United States and Canada, where once again human error, greed or ignorance has created many of these problem structures (Table 1 and Figure 1).


Ocnele Mari Brinefield, Romania

SALROM, a government-owned company, solution mines Badenian (Miocene) halite in the Valcea Prefecture of Romania (Figure 2). Production from Field 2 was shut down on March 5, 1991, after significant earth vibrations were noted by SALROM workers. A subsequent sonar survey showed that poorly monitored brinefield leaching between 1971 and 1991 had created a gigantic merged cavern as salt pillars separating adjoining caverns were inadvertently dissolved (Figure 2); the upper parts of the captive boreholes 363, 364, 365, 366, 367 and 369 had merged into a common cavern. The cavern was filled with some 4.5 million cubic metres of brine, was less than a hundred metres below the landsurface, was more than 350 metres across and was overlain by loosely consolidated sandy marls (von Tryller, 2002; Zamfirescu et al., 2003). The cavern was overlain by a bowl of subsidence and, even though mining operations in the region of the cavern completely stopped in 1993, the ground above continued to subside to a maximum of 2.2 metres prior to the September 2001 collapse. In the period 1993 - 2001 the cavern continued to enlarge and the northern part of the cavity expanded by some 25-35 m (see inset in Figure 3 showing sonar surveys 1995 -2001).


The 1993 sonic survey of the cavern showed it was so large and shallow that its roof must ultimately collapse, predicting when was the unknown. The fear was that if it collapsed catastrophically, it could release a flood of at least a million cubic metres of brine. Brine could be released either over several hours (least damaging scenario) or instantly forming a wave of escaping water several metres high that would flood the nearby Sarat River valley for many kilometres downstream. There were 22 homes on top of the cavern that would be immediately affected; also at risk were the hundreds and perhaps thousands of residents in the area of the saltworks and river valley below, as well as the local ecology and the civil/industrial infrastructure. Ongoing brine delivery from other nearby operational SALROM caverns to the large Oltchem and Govora chemical plants would also be disrupted. When collapse did ultimately occur (twice in the period 2001 - 2004), each episode took place over a number of hours and so a catastrophic wave of water did not eventuate.

According to Solution Mining Research Institute, there was no good engineering solution to prevent the Ocnele cavern roof from collapsing. Prior to collapse, brine in the cavern exerted pressure against the roof helping to hold it up; removing even a small amount of brine would remove hydraulic pressure that could possibly trigger a catastrophic collapse. Nor was it practical to fill the cavern with sand and industrial wastes as suggested by the Romanian government. It would take too long and it was unsafe to place men and equipment on top of the still expanding cavern. Even if it could be done, only the areas directly below the injection wells would be filled. The best solution was to construct a dam as close to the cavern as possible, and then perhaps trigger a controlled collapse by pumping out the cavern brine.

A partial collapse of the cavern roof atop well 377 occurred on 12 September of 2001, the ensuing brine flood killed a child and injured an older man. The collapse began at 7 pm with brine spilling out of wells 365 and 367. Southward of well 377 a collapse cone some 10 metres across started to form and fill with brine. The cone continued to expand and fill with water until its southern lip was breached. Water spilled out of the cone and down the hill slope. The flow rate of the expelled flow reached a maximum of 17 m3/sec at 3 am on September 13th, with flow continuously exceeding 10 m3/sec for a period of more than 6 hours around that time. By 7 am flow was down to 4-5 m3/sec and by 7 pm, 24 hours after the flow began the flow rate was 0.4-0.5 m3/sec. Some 24 hours after the onset, roof collapse had formed a water-filled lake with an area of 2.4 ha (Figures 2, 3; Zamfirescu et al., 2003).


A second roof collapse occurred on July 13, 2004. Realitatea Romaneasca, Romania, reported that ground collapse occurred at 9:30 pm July 13, 2004, possibly triggered by heavy, recent rains. The wellhead of bore 365 was destroyed in this collapse, and many other well heads in Field 2 had been destroyed by ongoing subsidence in the period 2001-2004. This collapse was the culmination of a series of collapse-related events that began on Monday July 11. It was planned that a purpose-built earthen dam would contain the brine flood. But when the roof fall occurred at 9:30 pm, the dam was breached 30 minutes later by some 250,000 m3 of salt water. Flow rate through the breach reached a maximum of 6 m3/sec. Collapse did not occur until just after local authorities had evacuated people from 50 homes in the area around the cavern. Most of the brine forced out by the roof collapse escaped into the Sarat (Olt) River. The government attempted to dilute the effects of this flush of salt water by releasing fresh water into the river from nearby dams. Fortunately, unlike the 2001 event, there was no loss of life.


Induced collapse, Gellenoncourt saltworks, France

Gellenoncourt is one of three sites that exploit halite beds of the Lower Keuper in the eastern part of the Paris Basin. The deposits extend from Cezanne in the west to Nancy with a length of 250 km and a width of 50 to 70 km. Salt production utilising solution mining techniques is focused on three sites along the Meurthe river, which all lie immediately to the east of Nancy.

On March 4, 1998, a sinkhole more than 50 m across and 40 m deep formed atop the SG4 and SG5 brinefield caverns in the Gellenoncourt saltworks near Lorraine, France (Figure 4; Buffet, 1998). It was an induced collapse designed to prevent a possible uncontrolled future ground collapse. The problem started in 1967 with the beginning of the exploitation of Triassic salt layers in the Keuper Fm. In total the Keuper is more than 150 m thick with five salt layers at its base, passing up into variegated and poorly consolidated marls and sandstones and capped by the Dolomite du Beaumont. The top of the salt is some 220 m below the surface and divided into five beds numbered 1 to 5 from top to bottom, with the solution mining program designed to leach beds 1 through 3 in the region of SG4 and SG5 caverns (Figure 5). Five wells, SG1 to SG5, were joined using hydrofracturing in February 1967. Theoretically, the process was designed to leave a substantial salt pillow atop all the cavities and so separate the solution cavities from the overlying marls (Figure 5; 1967-1971). The SG4 and SG5 caverns unexpectedly joined in 1971. Brine injection to these two wells was stopped, but crossflowing brines flowing to the producing SG1 well continued to excavate these two caverns. By 1982 the salt cushion in the roof of the SG4-SG5 cavern had completely dissolved away, placing the cavity roof in direct contact with the base of the marls.


From 1982 until October 1992 there was no further upward growth of the cavern roof. Then a 25 m-thick section of the variegated marls in the roof broke free and fell to the cavern floor, leaving a large section of the cavity roof in direct contact with the brittle Dolomite de Beaumont. This stiff dolomite layer prevented any further immediate collapse of the roof and consequent propagation of the cavity to the surface. But continued growth of the roof span beneath the dolomite would mean a later, larger, perhaps catastrophic collapse. In 1995 the operator tried to induce a controlled collapse by placing a submerged pump in the cavity and pumping out brine to create an exposed upper face. But this didn’t work. The next approach was to further enlarge the roof span by injecting 300,000 m3 of freshwater into the cavity. Collapse occurred on March 4, 1998, forming a 50 m wide crater. To protect the surrounding countryside from any brineflood damage a dam was constructed to capture any brine overflow, but in the actual event it was not needed.

Retsof Mine, New York State, USA

The Retsof Mine collapse is perhaps one of the best documented examples of an operational mine being lost to dissolution features and consequent flooding, yet even here the exact causes of the mine loss are still argued. Was it because of the intersection of the expanding mine workings with a natural water-saturated salt dissolution and fracture system, or the insection of the mine workings with brine filled cavities formed by wild brining operations in the 1800s, or was it a result of mine roof instabilities related to a change in room and pillar sizes?

At the time it was operational, the 24 km2 area of subsurface workings in the AKZO-Retsof salt mine made it the largest underground salt mine in the USA, and the second largest salt mine in the world (Figure 6, 7).


Operational history

Retsof Mine started operating in 1885 after completion of the 3.7× 4.9-meter wide, 303.5-meter-deep Shaft #1. The mine claimed an initial 5,460-metric-tons per day hoist capacity (Goodman et al., 2009). Early main haulage ways were driven east and west while production headings were driven north (updip) for salt-tramming ease. Room heights in the 6-meter-thick salt bed were 2 to 4 meters, with salt left in both the floor and roof. Four-meter-high rooms were worked in two benches. Rooms were 9.2 meters wide and separated by 9.2-meter pillars. At that time there was no timbering, the mine was dry, mine air temperature was 17°C, and the mine was largely gas-free.

By 1958, the Retsof Mine was connected to the former Sterling Mine for ventilation and emergency escape purposes [Figure 6a; Gowan et al., 1999]. By the late 1960s, the mine had advanced beneath the modern Genesee River and Valley (Figure 6b). By the early 1970s, the Retsof Mine operators had installed an underground surge bin, fed by a new conveyor system, and the old rail-haulage system was eliminated. Mainline conveyors led to yard or panel belts feeding each mining section, where a Stamler feeder-breaker crushed salt delivered by diesel-powered Joy shuttle cars. In the early 1980s, the shuttle cars were gradually replaced by load-haul-dump vehicles (LHDs). In 1969, Netherlands-based Akzo Corporation acquired International Salt Company and operated the Retsof mine until its abandonment due to flooding in 1995.

During April 1975, an explosion occurred in the original Sterling B Shaft during efforts to control water inflow into the Retsof Mine from this abandoned and partially collapsed shaft (Goodman et al. 2009). The leaky B Shaft had not been used or maintained for years. By 1975, International Salt was concerned that freshwater inflow from the B Shaft could pose a salt dissolution, collapse, and flooding risk to the then-connected Retsof Mine. Removal of a partial shaft blockage of timber and rock debris was attempted as a means of regaining airflow needed to safely access, rehabilitate, and grout off the water inflow to the shaft and mine below. A maintenance crew attempted to dislodge the shaft obstruction by pushing a large boulder into the shaft that was to drop down and knock through the debris. A methane explosion occurred upon impact. The upward force of the explosion killed four people on the surface near the shaft collar and injured others. On November 19, 1990, a roof fall resulted in two fatalities. Deformation and fracture of roof salt can occur because of a concentration of stresses; i.e., punching of the roof by stiff pillars. After the fatalities, the mine tested smaller, yielding pillars to alleviate roof falls (Figure 6b). Positive test results led to the adoption of a yield-pillar design.

The Retsof mine was lost to water flooding in 1994-1995.Before abandonment, the mine had been in operation since 1885, exploiting the Silurian Salina Salt and prior to shut down was producing a little over 3 million tons of halite each year. At that time it supplied more than 50% of the total volume of salt used to de-ice roads across the United States.

Geology and hydrology in the vicinity of Retsof Mine

The Genesee Valley sediments preserve evidence of several complex geologic processes that include; (1) tectonic uplift of Palaeozoic sedimentary rocks and subsequent fluvial down cutting, (2) waxing and waning glacial events that drove erosion of bedrock and the subsequent deposition of as much as 750ft of glacial sediments; and (3) ongoing erosion and deposition by postglacial streams (Figure 7a; Yager, 2001; Young and Burr, 2006). The Genesee Valley spans through western New York north to south from Avon, NY to Dansville, NY, including the Canaseraga Creek up through its mergence with Genesee River. A detailed section from Palaeozoic rocks and younger have been recorded in the Genesee River Valley (Figure 7a, b); however, detailed analysis of glacial sediments and till are still somewhat scarce. The B6 salt bed (Retsof Bed) of the Vernon Formation was the salt unit extracted at the Retsof Mine (Figures 7b). Several other salt layers exist in the Salina Group both above and below the B6. These salt layers include two horizons in Unit D at the base of the Syracuse Formation approximately 50 m (160 feet) above the B6 salt level.

Quaternary-age sediment in the Genesee Valley consists mostly of unconsolidated glacial sediments ranging up to 750 feet thick. These sediments encompass gravel, sand, silt and clays that were deposited mostly during the middle and late Wisconsin deglaciation and filled the lower parts of the pre-existing glacial scour valley. End moraines consisting of glacial debris were deposited in lobes to the south of the slowly retreating glacier. As the glacier had scoured through the valley, carving out bedrock and accumulating sediment, steep-sided valley walls were cut and pro-glacial lakes formed. The glacial lake sediments are dominated by muds, but also include large boulders and cobbles carried to the lake depressions by glacial ice. Fluvial sediment from the Genesee River and Canaseraga Creek also drained into these glacial lakes. A final pro-glacial lake formed as the Fowlerville end moraine was deposited. The Fowlerville end moraine extends approximately 4.5 to 8 miles north of the Retsof collapse site (Figure 7a). The various glacial lakes and moraines disrupted the normal flow fluvial patterns of most local drainages and creeks in the valley. Alluvium is the uppermost layer of the surface and is variable in thickness throughout the valley, but normally ranges about fifty feet thick and is still being deposited across the Genesee River Valley floodplain (Yager, 2001).

The aquifer system is hosted within the glacial valley-fill and consists of three main aquifers separated by two confining layers. It is underlain by water-bearing zones in fractured Palaeozoic bedrock (Yager, 2001). The glacial aquifers are bounded laterally by the bedrock valley walls. The uppermost aquifer consists of alluvial sediments 20 to 60 ft thick (unit 1 in Figure 7b); the middle aquifer consists of glaciofluvial sand and gravel less than 10 ft thick (unit 3 in Figure 7b); and the lower aquifer consists of glaciofluvial sand and gravel about 25 ft thick overlying the bedrock valley floor (unit 5 in Figure 7b). These aquifers are separated by aquitards dominated by muds and clays (Units 2 and 4 in Figure 7b).

The now abandoned Retsof Mine lies 550 to 600 ft below the eroded valley floor (Figure 7b). Hence, the upper and middle aquifers are separated by an upper confining layer of lacustrine sediments and till as much as 250 ft thick, and the two confined aquifers are separated by a lower confining layer of undifferentiated glaciolacustrine sediments as much as 250 ft thick. The principal water-bearing zone in the bedrock overlying the mine consists of fractured carbonates and sands near the contact between the Onondaga and Bertie Limestones. The fractured aquifer that occurs at this level in the stratigraphy supplied a significant volume of the water that ultimately flooded the Retsof Mine. The glacial aquifers are hydraulically connected at the edges of the confining layers and in subcrop zones, where water-bearing zones in the bedrock intersect a fractured and karstified bedrock surface.

Ground water within the valley generally flowed northward and updip before the mine collapse (Yager, 2001). The hydraulic head distribution in the confined aquifers under natural (precollapse) conditions is assumed to have been similar to that in the upper aquifer before the collapse, but water levels in the confined aquifers were probably above the water table beneath the valley floor. Much of the ground water reservoired along the fractured Onondaga/Bertie Limestone contact also flowed northward to escape at the Bertie Limestone subcrop, now located in the valley north of the Fowlerville Moraine (Figure 7c).

Water influx tied to changes in room and pillar mining?

In 1993, ceiling falls began to occur in rooms in the deepest part of the Retsof mine near its southern boundary (Figure 6a; Yager et al., 2009). In response, the mine owner, AkzoNobel Salt Incorporated (ANSI), turned to an innovative “yielding pillar” mining technique that utilised many narrow (20 feet × 20 feet) pillars rather than few wide ones in the mined section (Figure 6b). Geotechnical analyses indicated that the resulting configuration would allow the salt pillars to slowly yield and create a “stress envelope” in the surrounding bedrock to support the entire mined room.

Closure monitoring was conducted in the yield-pillar test panels and the two full-scale panels during mining to measure panel behaviour and to see if the new design mitigated the floor and roof problems being experienced in the large pillar area of the mine (van Sambeek et al., 2000). Monitoring initially indicated that room closure rates were slightly greater than expected, but had an overall character (trend) of steadily decreasing rates, which is consistent with stable conditions. This trend changed dramatically to a rapid and unstable closure rates in the final weeks leading up to the inflow. The change in trend was initially obscured by fluctuating closure rates because salt extraction was occurring between the two yield-pillar panels as the monitored abutment pillar was isolated. Whereas the closure rates were expected to decrease after this mining was complete, they did not; in fact, they increased. This change in panel character later interpreted to indicate that a pressure surcharge existed or developed over two of the yield-pillar panels prior to the in- flow (Gowan et al., 1999).

 

Loss of roof stability and flooding

In November 1993, strain measurements in a yielding-pillar area within the mine indicated a larger than expected deformation of salt near the eastern wall of room 2 Yard South (Figure 6a, b). Mining in the area was halted as ceiling falls of salt continued during the next four months. On March 12, 1994, a magnitude 3.6 seismic event, caused by a large roof collapse, was detected by seismometers more than 300 miles away. Mine workers attempted to enter room 2 Yard South but found it was blocked by a pile of rock rubble within the formerly mined room and that saline water entering via fractures in the mine roof. Over the next several weeks, Akzo made concerted attempts to save the mine by pumping water out and drilling around the collapse area to inject cement grout so as to stabilise the collapsed room and prevent a further inflow of water. Meanwhile, unstable shale layers overlying room 2 Yard South sagged and collapsed to form a 300-foot-diameter zone of rock rubble that slowly propagated upward through overlying layers of shale (Figures 6b, 8a, b). This column of rock rubble is referred to as a rubble chimney.

The propagating rubble chimney eventually reached a layer of carbonate (limestone) rock that was strong enough to temporarily resist further collapse, stopping further the rubble chimney’s upward progression. At this point, the flow of water into the mine stabilised at about 5,500 gallons per minute. Water entering the mine was saline and probably a mixture of saline water from the shale and a prominent fracture zone aquifer within the Onondaga and Bertie Limestones (Figures 7c). By the end of March 1994, tons of cement grout had been injected into the mine and the rubble chimney through nearly 30 boreholes drilled in the collapse area, but these efforts failed to stem the rate of water flowing into the mine and the inflow was becoming increasingly less saline.


On April 6, 1994, the limestone rock layer collapsed, and 550 feet of unconsolidated sediments in the Genesee River valley quickly slumped downward into the resulting cavity, forming a sinkhole at the land surface, more than 15 feet deep and several hundred feet across (Figure 8, 9a b). The collapse of the limestone rock was like pulling the plug in a bathtub—it allowed groundwater from a fresh-water aquifer at the base of the unconsolidated glacial sediments (the lower confined aquifer (Figure10) to drain downwards through the rubble chimney and into the mine. By mid-April, a second collapse occurred in an adjacent room (11 Yard West; Figure 6b). On May 25, a drilling crew working above room 11 Yard West felt tremors and removed their drill rig, and themselves, just before this second sinkhole formed at the land surface. This one had a surface expression that was more than 50 feet deep and several hundred feet across (Figures 6a, 8a, 9a,b). The discharge from the aquifer through both rubble chimneys increased the flow of water into the mine to about 18,000 gallons per minute.

Water began to fill the southern end of the mine and then spread steadily northward, dissolving the bases of the salt pillars that supported the mine ceiling (Figure 6a). As the pillars gave way, the southern part of the mine began to collapse, causing the land surface above it to subside. The greatest subsidence (more than 15 feet) was beneath the two sinkholes, which altered the channel of Beards Creek, allowing surface water to fill the sinkholes. The surface water did not flow downward to the mine, however, because hundreds of feet of fine-grained sediments underlie the Genesee River valley. The instability also forced the closure of the U.S. Route 20A bridge over Beards Creek; the southern end of the bridge eventually subsided by 11 feet (Figure 9c, d). The bed of the Genesee River 1 mile north of the collapse areas subsided by as much as 5 feet and altered the pattern of sediment scour and deposition along a 1.5-mile reach downstream of Beards Creek.

Events indicating loss of mine

The eventual loss of the Retsof Salt Mine occurred in stages, driven first by “out of salt” roof breaches, followed by ongoing salt dissolution of the water-encased salt pillars in the flooded mine. It began in the early morning hours of March 12, 1994, with a magnitude 3.6 earthquake. The quake was caused by the catastrophic breakdown of a small mine pillar and panel section some 340 meters below the surface and was accompanied by the surface collapse of an area atop the mine that was some 180 by 180 meters across and 10 meters deep. This all occurred at the southern end of the mine near the town of Cuylerville. A month later, on April 18, an adjacent mine room collapsed to form a second collapse crater (Figure 6a, b) The initial March 12 collapse in the mine was accompanied by an inrush of brine and gas (methane) and by a sustained intense inflow of water at rates in excess of 70 m3/min, via the overlying now fractured limestone back (Gowan and Trader, 2000).

In a little more than a month, the two steep-sided circular collapse features, some 100 meters apart, had indented the landscape above the two collapsed mine rooms (Figures 6, 8, 9). The northernmost collapse feature, which was more than 200 meters across, included a central area that was about 60 meters wide and had subsided about 6 to 10 meters. The southernmost feature, which was about 270 meters in diameter, included a central area that was about 200 meters wide and had subsided about 20 meters (Figure 6b). Fractures extending up from the broken mine back created hydraulic connections between aquifers, which previously had been isolated from each and so provided new high volume flow routes for rapid migration of perched groundwaters into the mine level.

Water flooded the mine at rates that eventually exceeded 60,000 litres per minute and could not be controlled by pumping or in-mine grouting. By January 1996 the entire mine was flooded. Associated aquifer drawdown caused inadequate water supply to a number of local wells in the months following the collapse; the fall in the water table as ground waters drained into the mine in effect meant some water wells went dry (Figure 8c; Tepper et al., 1997).

Aside from the loss of the mine and its effect on the local economy, other immediate adverse effects included abandonment of four homes, damage to other homes (some as much as 1.5 kilometers from the sinkholes), the loss of a major highway and bridge, loss of water wells and prohibition of public access to the collapse area (Figure 9). Land subsidence, possibly related to compaction induced by aquifer drainage to the mine, even occurred near the town of Mt. Morris some 3 miles south-west of the collapse area. Longer term adverse effects are mostly related to increasing salinization of the lower parts of the Genessee Valley aquifer system in the vicinity of the mine (Figure 10; Yager, 2013).

 

What caused the loss of the mine?

Post-mortem examination of closure data from the two failed mine panels has been interpreted as indicating an anomalous buildup of fluid pressure above the panels in the period leading up to their collapse (Gowan et al., 1999). The initial influx of brine and gas following the first collapse coincided with the relief of this excess pressure.

Gowan and Trader (1999) argued for the existence of pre-collapse pressurised brine cavities and gas pools above the panels and related them to nineteenth-century solution mining operations. They document widespread natural gas and brine pools within Unit D of the Syracuse Formation approximately 160 ft above the mined horizon in the Retsof Mine. The satellite image also shows that collapse occurred in a pre-existing landscape low that defined the position of Beard Creek valley above the mine (Figure 6a). Brine accumulations likely formed in natural sinks, long before salt solution mining began in the valley. Salt in the shallow subsurface dissolved naturally, driven by the natural circulation and accumulation of meteoric waters along vertical discontinuities, which connected zones of dissolving salt to overlying fresh water aquifers (see Warren, 2016, Chapter 7 for a detailed documentation of this salt related hydrology and geomorphology).

Gowan and Trader (2003) argued that daylighting sinkholes had formed by the down-dropping of the bedrock and glacial sediments into pre-existing voids created by the dissolution of salt and the slaking of salt-bearing shale upon exposure to fresh water. It is likely that the extent of these brine filled voids was exacerbated by the “wild-brining” activities of salt solution miners in the 1800’s.

Nieto and Young (1998) argue that the transition to the yield pillar design was a contributing factor to the loss of mine roof integrity. Loss of mechanical integrity in the roof facilitated fracturing and the influx of water from anthropogenic “wild brine” cavities. The exact cause of the loss of roof integrity and subsequent mine flooding is still not clear. What is clear is that once the Retsof mine workings passed out of the salt mass, and into the adjacent non-salt strata, the likelihood of mine flooding greatly increased.

Even so, the loss of the Retsof salt mine to flooding was a total surprise to the operators (Van Sambeek et al., 2000). The mine had operated for 109 years with relatively minor and manageable incidents of structural instability, water inflow, and gas occurrences. A substantial database of geological information was also collected throughout the history of the mine. It was this relatively uneventful mine history and the rich technical database that provided support for pre-inflow opinions by mine staff that there was no significant potential for collapse and inundation of the mine. The Retsof collapse took place in a salt-glacial scour stratigraphy and hydrology near identical to that in the Cayuga Mine region.

 

Patience Lake Potash Mine flood

In the 1970s the Patience Lake potash mine operation, located on the eastern outskirts of Saskatoon, Canada, encountered open fractures tied to a natural collapse structure and was ultimately converted to a successful solution mining operation (Figure 11). Grouting managed to control the inflow and mining continued. Then, in January of 1986, the rate of water inflow began to increase dramatically from the same fractured interval (Figure 12; Gendzwill and Martin 1996).

At its worst, the fractures associated with the structure and cutting across the bedded ore zones were leaking 75 m3/min (680,000 bbl/day) of water into the mine. The water was traced back to the overlying Cretaceous Mannville and possibly the Duperow formations. Finally, in January 1987 the mine was abandoned. It took another six months for the mine to fill with water. Subsequent seismic shot over the offending structure suggested that the actual collapse wasn’t even penetrated; the mine had merely intersected a fracture within a marginal zone of partial collapse (Gendzwill and Martin 1996).

Part of the problem was that the water was undersaturated and quickly weakened pillars and supports, so compromising the structural integrity of the workings. The unexpected intersection of one simple fracture system resulted in the loss of a billion dollar conventional potash mine. Patience Lake mine now operates as a cryogenic solution mine by pumping warm KCl-rich brine from the flooded mine workings to the surface. Harvesting of the ponds takes place during winter after cryogenic precipitation of sylvite in at-surface potash ponds (Fig. 11).

 

Unlike the Patience Lake Mine flood, there was a similar episode of water inflow in the nearby Rocanville Potash Mine. But there a combination of grouting and bulkhead emplacement in succeeded in sealing off the inflow, thus saving the mine (see Warren 2016 for detail). Unlike Patience Lake, the brine from the breached structure in Rocanville was halite-saturated, so limiting the amount of dissolution damage in the mine walls. Different outcomes between the loss of the Patience Lake Mine and recovery from unexpected flooding in the Rocanville Mine likely reflects the difference between intersecting a natural brine-filled dissolution chimney that had made its way to the Cretaceous landsurface and is now overlain by a wide-draining set of aquifer sediments, versus crossing a blind dissolution chimney in a saline Devonian sediment surround that never broke out at the Cretaceous landsurface. Understanding the nature of the potential hydrological drainages and water source is a significant factor in controlling unexpected water during any salt mine expansion.


Lake Peigneur, Louisiana

Lake Peigneur is a natural water-filled solution doline that overlies the dissolving crest of the Jefferson Island Salt Dome Figure 13). The most recently risen part (salt spine) of the Jefferson Island stock crest, just west of the town of New Iberia, Louisiana, is now 250 m (800ft) higher than the adjacent flat-topped salt mass, which is also overlain by a cap rock. The boundary shear zone separating the spine from the less active portion of the crest contains a finer-grained “shale-rich” anomalous salt zone that had been penetrated in places by the former Jefferson Island mine workings. The known salt anomaly (BSZ) defined a limit to the extent of salt mining in the diapir, which was focused on extracting the purer salt within the Jefferson Island spine, in a mining scenario much like the fault shear anomaly, as mapped by Balk (1953), defined the extent of the workings at nearby Avery Island. The spine and its boundary “shear” zone are reflected in the topography of the Jefferson Island landscape, with a natural sub-circular solution lake, Lake Peigneur, created by the dissolving shallow crest of the most recently-active salt spine.

On November 20, 1980, one of the most spectacular sinkhole events associated with oilwell drilling occurred atop the Jefferson Island dome just west of New Iberia. Lake Peigneur disappeared as it drained into an underlying salt mine cavern and a collapse sinkhole, some 0.91 km2 in area, developed in the SE portion of the lake (Figure 13; Autin, 2002). In the 12 hours following the first intersection the underlying mine had flooded, and the lake was completely drained.

Drainage and collapse of the lake began when a Texaco oil rig, drilling from a pontoon in the lake, breached an unused section of the salt mine some 1000 feet (350 metres) below the lake floor (Figure 14a). Witnesses working below ground described how a wave of water instantly filled an old sump in the mine measuring some 200 ft across and 24 feet deep. The volume of floodwater engulfing the mine corridors couldn’t be drained by the available pumps. At the time of flooding the mine had four working levels and one projected future level. The shallowest was at 800 feet, it was the first mined level and had been exploited since 1922. The deepest part of the mine at the time of flooding was the approach rampways for a planned 1800 foot level. In 58 years of mine life, some 23-28 million m3 of salt had been extracted. Prompt reaction to the initial flood wave by mine staff allowed all 50 personnel, who were underground at the time, to escape without anything more than a few minor injuries.


The rapid flush of lake water into the mine, probably augmented by the drainage of natural solution cavities in the caprock below the lake floor, meant landslides and mudflows developed along the perimeter of the sinkhole, and that the lake was enlarged by 28 ha. The surface entry hole in the floor of Lake Peigneur quickly grew into a half-mile-wide crater. Eyewitnesses all agreed that the lake drained like a giant unplugged bathtub—taking with it trees, two oil rigs (worth more than $5 million), eleven barges, a tug boat and most of the Live Oak botanical gardens. It almost took local fisherman Leonce Viator Jr. as well. He was out fishing with his nephew Timmy on his fourteen-foot aluminium boat when the disaster struck. The water drained from the lake so quickly that the boat got stuck in the mud and they were able to walk away! The drained lake didn’t stay dry for long, within two days it was refilled to its normal level by Gulf of Mexico waters flowing backward into the lake depression through a connecting bayou (Delcambre Canal, aka Carline Bayou). But, since parts of the lake bottom had slumped into the sinkhole during the collapse, the final water level in some sections was higher than before relative to previous land features. It left one former lakefront house aslant under 12 feet of water.

Of course, an anthropogenically induced disaster like this attracted the lawyers like flies to a dead dingo. On 21 November 1980, the day after the disaster, Diamond Crystal Salt filed a suit against Texaco for an unspecified amount of damage. On 25 November, Texaco filed a countersuit against Diamond Crystal. The Live Oak Gardens sued both Diamond Crystal and Texaco. Months later, the State of Louisiana was brought into the suit since the incident occurred on state land. One woman sued Texaco and Wilson Brothers (the drillers) for $1.45 million for injuries (bruised ribs and an injured back) received while escaping from the salt mine. Less than a week before the scheduled trial, an out-of-court settlement was reached between the major players. Due to human error, related to a triangulation mistake when siting the drilling barge, Texaco and Wilson Brothers agreed to pay $32 million to Diamond Crystal and $12.1 million to the Live Oak Gardens.

An ongoing environmental catastrophe that was anticipated by environmental groups at the time of the accident never materialized. The lake quickly returned to its natural freshwater state, and with it the wildlife was largely un-affected. Nine of the barges eventually popped back up like corks (the drilling rigs and tug were never to be seen again). The torrent of water helped dredge Delcambre Canal so that it was two to four feet deeper. And of course, the former 1 metre deep Lake Peigneur was now 400 metres deep in the vicinity of the borehole!

Interestingly, the filling of the mine workings with water drastically slowed the rate of land subsidence atop the mine (Figure 14b). Measurements had been carried out between 1973 and 1983, some 7 years before the accident and continued for 2 years afterward (Thoms and Gerhle, 1994). Slowing reflects the post-accident reduction in the total pressure exerted on the roof of the mine to half its pre-accident levels. Prior to the accident, there was no hydrostatic pressure to alleviate some of the lithostatic pressure exerted by the weight of the overburden and so land subsidence above the mine workings was relatively rapid.

Although this incident is not directly related to any aspect of the salt mining operation and no human lives were lost (although three dogs perished), it clearly illustrates the speed of potential leakage following a breach in a cavern roof in any shallow storage facilities filled with low-density fluids. It also illustrates the usual cause of such disasters – human error in the form of a lack of due diligence, a lack of forward planning and a lack of communication between various private and government authorities. It also illustrates that filling a solution cavity with water slows the rate of subsidence atop a large salt cavity and that waters after the disturbance will return quietly to a state of density stratification.

The incident had wider resource implications as it detrimentally affected the profitability of other salt mines in the Five Islands region (Autin, 2002). Even as the legal and political battles at Lake Peigneur subsided, safe mining operations at the nearby Belle Isle salt mine came into contention with public perceptions questioning the structural integrity of the salt dome roof. Horizontal stress on the mineshaft near the level where the Louann Salt contacts the overlying Pleistocene Prairie Complex had caused some mine shaft deterioration. Broad ground subsidence over the mine area was well documented and monitored, as was near continuous ground water leakage into the mine workings. The Peigneur disaster meant an increased perception of continued difficulty with mine operations and an increased risk of catastrophic collapse was considered a distinct possibility. In 1985, a controlled flooding of the Belle Isle salt mine was completed as part of a safe closure plan.

Subsidence over the nearby Avery Island salt mine (operated by Cargill Salt) has been documented since 1986. This is oldest operating salt mine in the United States and has been in operation since the American Civil War, and after the Lake Peigneur disaster the mine underwent a major reconstruction and safety workover. Mine management and landowners did not publicly disclose the technical details of rates of subsidence, but field observations revealed the nature of the subsidence process. Subsidence along the mine edge coincided with a topographic saddle above an anomalous salt zone located inside the mined salt area, ground water had seeped into the mine, and there were a number of soil gas anomalies associated with the mine. Small bead-shaped sinkholes were initially noticed in the area in 1986, then over several years, a broad area of bowl-shaped subsidence and areas of gully erosion formed (Autin, 2002). Reconstruction has now stabilised this situation. Much of the subsidence on Avery Island was a natural process that occurs atop any shallow salt structure. Dating of middens and human artifacts around salt solution induced water-filled depressions atop the dome shows dissolution-induced subsidence is a natural process that extends back well beyond the 3,000 years of human occupation documented on the island.

Compared to the other salt domes of the Five Islands, Cote Blanche Island has benefited from a safe, stable salt mine operation throughout the mine life (Autin, 2002). Reasons for this success to date are possibly; (i) mining operations have not been conducted as long at Cote Blanche Island as other nearby domes, (ii) the Cote Blanche salt dome may have better natural structural integrity than other islands, thus allowing for greater mine stability (although it too has anomalous zones, a salt overhang, and other structural complexities), and (iii) the salt is surrounded by more clayey (impervious) sediments than the other Five Islands, perhaps allowing for lower rates of crossflow and greater hydrologic stability.

Haoud Berkaoui oilfield, Algeria

Located in the Sahara, some 32 km southwest of Ouargla City, the Haoud Berkaoui oilfield is an area of subsidence where numerous exploration and development wells were completed in the 1970s. Of these, the OKN32 and OKN32BIS wells have collapsed into an expanding collapse doline. It surfaced in October 1986 when a crater, some 200 metres across and 75 metres deep, formed (Morisseau, 2000). Today the solution cavity continues to expand and is now some 230 by 600 metres across. Its outward progression is continuing at a rate around 1 metre per year. The collapse is centred on two oil wells drilled in the late 1970s. The problem began in 1978, when the OKN32 oil exploration well was abandoned because of stability problems in Triassic salt at a depth around 650 metres. The target was an Ordovician sandstone at a depth of 2500m. Because of the technical problems associated with significant caverns at the level of the salt, the well was abandoned without casing being set in the salt, probably facilitating the escape of artesian waters (Morisseau, 2000; Bouraoui et al., 2012).


When it reached the 600m level, the well had already passed through 50 metres of anhydrite (220-270m depth) along with interbedded anhydrite clay and dolomite 270m -450 m depth). These are evaporite sediments that, in their undisturbed state, can act as aquicludes or aquitards to any access by unconfined phreatic groundwaters, although at such shallow depths the evaporite beds are likely also to be variably overprinted by active-phreatic dissolution processes. Prior to drilling it was thought that the Senonian halite extended continuously to a depth of 600 m in the well and in turn was underlain by 50 m of anhydrite (600-650m depth). Below the halite-anhydrite is an artesian aquifer (Albian) with a natural hydraulic head that is larger than the surface aquifer head by 2.5 MPa (Morisseau, 2000)

In 1979 a second well, OKN32BIS, was drilled located some 80 metres from the previous well and it successfully obtained its Ordovician target. But in March 1981, the lining of this second well broke, probably because of cavity collapse at a level around 550 metres (once again the regional level of salt) and the well was lost. Five years later, on October 1 1986, a large surface crater formed, centred on these two wells. The initial at-surface diameter was 200 metres and it was 75 metres deep, today it is even larger (Figure 15). Cavern diameter below the stope breakout at that time was estimated to be 300m and water flows to be around 2000 m3/hour.

Since the initial stope breakout, leaching has become progressively less effective and expansion rates have slowed (Morisseau, 2000). This is because cavern growth and water outflow flow are thought to take place preferentially near the centre of the collapse, which is now far from the collapsing cavern walls Dissolving salt may be salinising the crossflowing groundwaters, leading to undocumented, but possible, ongoing degradation of freshwater oases in the region. Continuing expansion is evidenced by the development of fresh centripetal cracks about the expanding collapse margin. Using MT-InSAR analysis, Bouraoui et al. (2012) documented ongoing subsidence near the crater, with an average subsidence of 4 mm per year (between 1992 and 2002). The zone of current zone of subsidence is centred on the OKN32 location and is slowly migrating north east.

As in the USA (see Table 2 and Warren, 2016, Chapter 13 for examples), the loss of these wells, in this case during their active life, emphasizes the need for caution when planning well abandonment in a salt bed, especially when it is highly likely that the salt is acting as a seal, or at least an aquitard, to a regional artesian system. The fact that the first well (OKN32) was lost during drilling argues that a natural breach or cavity was already present in the salt bed and perhaps was already stoping its way to the surface. It is also possible that the inappropriate completion and cementation of casing levels, prior to the well’s abandonment, may have accelerated cavity expansion. In hindsight, the loss of the second well some 5 years later was highly likely as was catastrophic cavity collapse 5 years after that; the OKN32BIS wellhead was situated only 80 metres from OKN32 and was dealing with the same cavern-ridden salt geology.

Summary

Regarding anthropogenically-enhanced salt karst, it is important to note that a salt mass used for storage has never failed catastrophically. Weak points tend to occur wherever “the outside has access to the inside,” so problems tend to be mostly where mine expansion breaches a salt edge (Warren, 2017). Likewise, almost all the problems related to well and cavity failure are more a matter of human error, either by negligence, or a lack of understanding by on-the-ground personnel. There is the same general rule of thumb when it comes to salt cavities and salt mines, and that is, “keep it in the salt!” Most failures and breaches occur when mining or solution leaching operations allow the cavity to contact the edge of the salt. There undersaturated water crossflows can exaggerate any uncontrolled dissolution problems. Often the salt edge is irregular due to natural dissolution and assumptions of flat or gently curved shapes to a salt edge are oversimplifications.

References

Autin, W. J., 2002, Landscape evolution of the Five Islands of south Louisiana: scientific policy and salt dome utilization and management: Geomorphology, v. 47, p. 227-244.

Bouraoui, S., Z. Cakir, R. Bougdal, and M. Meghraoui, 2012, MT-InSAR monitoring of ground deformation around the Haoud Berkaoui sinkhole (SE Algeria): Geophysical Research Abstracts, EGU General Assembly 2012, held 22-27 April, 2012 in Vienna, Austria, v. 14, EGU2012-3344.

Buffet, A., 1998, The collapse of Compagnie des Salins SG4 and SG5 drilling: Proc. S.M.R.I. Fall Meeting, Rome,, p. 79-105.

Gendzwill, D., and N. Martin, 1996, Flooding and loss of the Patience Lake potash mine: CIM Bulletin, v. 89, p. 62-73.

Goodman, W. M., D. B. Plumeau, J. O. Voigt, and D. J. Gnage, 2009, The History of Room and Pillar Salt Mines in New York State,” in S. Zuoliang, ed., Proceedings, 9th International Symposium on Salt, Beijing, China, September 4–6, 2009, v. 2, Gold Wall Press, Beijing, China, p. 1239–1248.

Gowan, S. W., and S. M. Trader, 1999, Mine failure associated with a pressurized brine horizon: Retsof Salt Mine, western New York: Environmental & Engineering Geoscience, v. 6, p. 57-70.

Gowan, S. W., and S. M. Trader, 2003, Mechanism of sinkhole formation in glacial sediments above Retsof Salt Mine, Western New York, in K. S. N. Johnson, J. T. , ed., Evaporite karst and engineering/environmental problems in the United States: Norman, Oklahoma Geological Survey Circular 109, p. 321-336.

Morisseau, J. M., 2000, Uncontrolled leaching of salt layer in an oil field in Algeria: Proc. S.M.R.I. Fall Meeting Technical Session, San Antonio, p. 330-333.

Nieto, A., and R. A. Young, 1998, Retsof Salt Mine Collapse and Aquifer Dewatering, Genesee Valley , Livingston County , NY, in J. Borchers, ed., Poland Symposium Volume: Land Subsidence, Spec. Pub. 8, Assoc. Engineering Geologists, p. 309-325.

Payment, K. A., 2000, Loss of the Retsof salt mine: legal analysis of liability issues, in R. M. Geertmann, ed., Proc. 8th World Salt Symp., Salt 2000, The Hague, v. 1: Amsterdam, Elsevier, p. 399-404.

Tepper, D. H., W. H. Kappel, T. S. Miller, and J. H. WilliaMS, 1997, Hydrogeologic effects of flooding in the partially collapsed Retsof salt mine, Livingston County, New York: US Geol. Survey Open File Report, v. 97-47, p. 36-37.

Thoms, R. L., 2000, Subsidence and sinkhole development over salt caverns: An introduction to the technology of solution mining; Spring 2000 Technical Class, p. 127-141.

Thoms, R. L., and R. M. Gehle, 1994, Analysis of a Solidified Waste Disposal Cavern in Gulf Coast Salt Dome: SMRI Fall Mtg. (1994) 637.

Thoms, R. L., and R. M. Gehle, 2000b, Winnfield mine flooding and collapse event of 1965: Proc. S.M.R.I. Fall Meeting Technical Session, San Antonio, p. 262-274.

Van Sambeek, L. L., S. W. Gowan, and K. A. Payment, 2000, Loss of the Retsof Mine: Engineering Analysis: Proceedings, 8th World Salt Symposium, The Hague, The Netherlands, May 7–11, R. M. Geertman (ed.), Elsevier Science Publishers B.V., Amsterdam, The Netherlands, pp. 411–416.

Von Tryller, H., 2002, The Cavern Field No. 11 in Ocnele Mari - History, Present and Future: Solution Mining Research Institute Proceedings, Spring Meeting, 28 April 1 May, 2002, Banff, Canada, p. 10 pp.

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

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

Yager, R. M., 2013, Environmental Consequences of the Retsof Salt Mine Roof Collapse, US Geological survey Open File Report 2013–1174, 10 p.

Yager, R. M., T. S. Miller, and W. M. Kappel, 2001, Simulated effects of 1994 salt-mine collapse on ground-water flow and land subsidence in a glacial aquifer system, Livingston County, New York: US Geological Survey Professional Paper, p. 1-80.

Yager, R. M., P. E. Misut, C. D. Langevin, and D. L. Parkhurst, 2009, Brine Migration from a Flooded Salt Mine in the Genesee Valley, Livingston County, New York: Geochemical Modeling and Simulation of Variable-Density Flow, USGS Professional Paper 1767, 59 p.

Young, R. A., and G. S. Burr, 2006, Middle Wisconsin glaciation in the Genesee Valley, NY: A stratigraphic record contemporaneous with Heinrich Event, H4: Geomorphology, v. 75, p. 226-247.

Zamfirescu, F., M. Mocuta, T. Constantinecu, E. Medves, and A. Danchiv, 2003, The main causes of a geomechanical accident of brine caverns at field II of Ocnele Mari - Romania: RMZ - Materials and Geoenvironment, v. 50, p. 431-434.

Zuber, A., J. Grabczak, and A. Garlicki, 2000, Catastrophic and dangerous inflows to salt mines in Poland as related to the origin of water determined by isotope methods: Environmental Geology, v. 39, p. 299-311.



 

 

Salt Dissolution (3 of 5): Natural Geohazards

John Warren - Tuesday, October 31, 2017


Introduction

Surface constructions and other anthropogenic activities atop or within evaporite karst terranes is more problematic than in subcopping carbonate terranes due to inherently higher rates of dissolution and stoping (Yilmaz et al., 2011; Cooper and Gutiérrez, 2013; Gutiérrez et al., 2014). Overburden collapse into nearsurface gypsum caves can create stoping chimneys, which break out at the surface as steep-sided dolines, often surrounded by broader subsidence hollows. Such swallow-holes, up to 20 m deep and 40 m wide, continue to appear suddenly and naturally in gypsum areas throughout the world.

Unlike the relatively slow formation of limestone karst, gypsum/halite karst develops on a human/engineering time-scale and can be enhanced by human activities (Warren, 2016, 2017). For example, in 2006, the Nanjing Gypsum mine in China broke into a phreatic cavity in a region of gypsum karst, driving complete flooding of the mine in some three days. Associated groundwater drainage caused a sharp drop in the local piezometric level of up to 90 m in a well in nearby Huashu village. Resultant ground subsidence severely damaged nearby roads and buildings (Wang et al., 2008). In Ukraine, dewatering of gypsum karst to facilitate sulphur mining substantially increased the rate of gypsum dissolution and favoured the expansion of sinkholes within an area affected by the associated cones of water-table depression (Sprynskyy et al., 2009). Natural evaporite karst enhanced by intrastructure focusing of drainage creates the various scales of problem across the Gypsum Plain of West Texas and New Mexico (Stafford et al., 2017).

Although halite is even more susceptible to rapid dissolution than gypsum, it typically is not a major urban engineering problem; large numbers of people simply do not like to live in a climate that allows halite to make it to the surface. However, in the Dead Sea region, the ongoing lowering of the water level encouraged karstic collapse in newly exposed mudflats and has damaged roads and other man-made structures (Frumkin et al. 2011; Shviro et al., 2017). Catastrophic doline collapse atop poorly managed halite/potash mines and solution brine fields is an additional anthropogenically-induced or enhanced geohazard in developed regions is discussed in detail in Warren, 2016 (Figure 1).


Gypsum karst is a documented natural hazard in many parts of Europe (Figure 2), and similar areas of shallow subcropping gypsum are common in much of the rest of the world (Table 1). For example, areas surrounding the city of Zaragoza in northern Spain are affected, as is the town of Calatayud (Gutiérrez and Cooper, 2002; Gutiérrez, 2014). Gypsum dissolution is responsible for subsidence and collapse in many urban areas around northern Paris, France (Toulemont, 1984), in urban areas in and around Stuttgart and other towns peripheral to the Harz Mountains in Germany (Garleff et al., 1997), in Pasvalys and Birzai in Lithuania (Paukstys et al., 1999), in the Muttenz-Pratteln area in northwestern Switzerland (Zechner et al., 2011), in the Perm area of Russia (Trzhtsinsky, 2002), in the Sivaz region of Turkey (Karacan and Yilmez, 1997), in the region centred on the city of Mosul in northern Iraq (Jassim et al., 1997) and in a number of areas of rapid urban development in eastern Saudi Arabia (Amin and Bankher, 1997a, b). Large subsidence depressions caused by gypsum dissolution in China have opened up in the Taiyuan and Yangquan regions of Shanxi Coalfield and the adjacent Hebei Coalfield.


Variation in the watertable level, induced by groundwater pumping or uncontrolled brine extraction, can be an anthropogenic trigger for dolines surfacing. As the watertable declines it causes a loss of buoyant support to the ground, it also increases the flow gradient and water velocity, which facilitates higher rates of crossflow and deeper aquifer recharge in subsequent floods and so reduces the geomechanical strength of the cover and washes away roof span support (Figures 1, 3). Dolines can also be associated with groundwater quality issues. Collapse dolines or sinkholes are frequently used as areas or sumps for uncontrolled dumping industrial and domestic waste. Because of the direct connection between them and the regional aquifer, uncontrolled dumping can cause rapid dispersion of chemical and bacterial pollutants in the groundwater. In the case of Riyadh region Saudi Arabia, a lake of near-raw sewage has appeared in Hit Dahl (cave) and is likely related to the increased utilisation of desalinated water for sanitation and agriculture (Warren, 2016). In the Birzai region of Lithuania numerous sinkholes developed in Devonian gypsum subcrop are in direct connection across the regional hydrology. Accordingly, the amount of agricultural fertilizer use is limited to help protect groundwater quality.

One of the problems associated with rapid surfacing of evaporite collapse features is that any assignment of sinkhole cause will typically lead to an assignment of blame, particulary when anthropogenic infrastructure has been damaged or destroyed by the collapse, or lives may have been lost. Areas of natural evaporite karst are typically areas of relatively shallow evaporites. Shallow evaporites make such regions suitable for extraction via conventional or solution mining. When a collapse does occur in a mined area, one group (generally the miners) has a vested interest in arguing for natural collapse, the others, generally the lawyers and their litigants, will argue for an anthropogenic cause. The reality is usually a combination of natural process enhanced to varying degrees by human endeavours. In the examples in this section, much of the driving process for the collapse is natural, while the cause of any unexpected karst-related disaster is typically geological ignorance combined with political/community intransigence. See Chapter 13 for a further discussion of karst and stope examples that include collapses and explosions where the anthropogenic drivers can dominate.

Problems in the Ripon area, Yorkshire, UK

The town of Ripon, North Yorkshire, and town’s surrounds experiences the worst ongoing gypsum-karst related subsidence in England (Figures 3, 4; Cooper and Waltham, 1999). Some 43 events of subsidence or collapse in the caprock over the Ripon gypsum have occurred over the last 160 years, within an area of 7 km2 (Figures 4). This gives a mean rate of one new sinkhole every 26 years in each square kilometre. Worldwide, the highest documented event rate occurs in Ukraine, in an area of thin and weak clay caprocks above interstratal gypsum karst, where new sinkholes appear at a rate of 0.01 to 3.0 per year per km2 (Waltham et al., 2005). In the Ripon area, numerous sags and small collapses also typify surrounding farmlands. Subsidence features are typically 10-30m in diameter, reach up to 20m in depth and can appear at the surface in a matter of hours to days (Figure 3). To the east of the town, one collapse sinkhole in the Sherwood Sandstone is 80 m in diameter and 30 m deep, perhaps reflecting the stronger roof beam capacity of the Sherwood Sandstone.

When a chimney breaks through, the associated surface collapse is very rapid (Figure 3 b-e). For example, one such subsidence crater, which opened up in front of a house on Ure Bank terrace on 23rd and 24th April, 1997, is documented by Cooper (1998.) as follows (Figure 3b).

“...The hole grew in size and migrated towards the house, to measure 10m in diameter and 5.5m deep by the end of Thursday. Four garages have been destroyed by the subsidence. This collapse was the largest of one of a series that have affected this site for more than 30 years; an earlier collapse had demolished two garages on the same site, and a 1856 Ordnance Survey map shows a pond on the same site. The hole is cylindrical but will ultimately fail to become a larger, but conical, depression. As it does so, it may cause collapse of the house, which is already damaged, and the adjacent road. The house and several nearby properties have been evacuated and the nearby road has been closed. The gas and other services, which run close to the hole, have also been disconnected in case of further collapse.”

Cooper (1998) found the sites of most severe subsidence in the Ripon area (including the house at Ure Terrace and in the vicinity of Magdelen's Road) are located at the sides of the buried Ure Valley, an area where the significant volumes of water seeps from the gypsum karst levels into the river gravels (Figure 4). In 1999 the Ure Terrace sinkhole was filled using a long conveyor belt that was cantilevered over the hole so that no trucked needed to back up close to the sinkhole opening. The hole was surcharged to a height of 0.5m. The hole remains unstable, but the collapse of the fill is monitored to document fill performance and the fill is periodically topped up. After the sinkhole was filled, the road adjacent to the sinkhole was re-opened and the site of the sinkhole fenced. The severely damaged Field View house remains standing next to the sinkhole. The nearby Victorian Ure lodge was not directly damaged by the 1997 sinkhole, but its western corner fell within the council-designated damage zone, and was left unoccupied. It fell into disrepair and was subsequently demolished (Figure 3b). A similar fate befell houses damaged by the surfacing of collapse sinkholes in and around Magdelen's Road, which is located a few hundred metres from Ure Terrace (Figure 3c-e). Shallow subcropping Zechstein gypsum (rehydrated anhydrite) occurs in two subcropping bedded units in this area, one is in the Permian Edlington and the other is in the Roxby Formation (Figure 4b). Together they form a subcrop belt about a kilometre wide, bound to the west by the base of the lowest gypsum unit (at the bottom of the Edlington Formation) and to the east by a downdip transition from gypsum to anhydrite in the upper gypsum-bearing unit of the Roxby Formation. The spatial distribution of subsidence features within this belt relates to joint azimuths in the Permian bedrock, with gypsum maze caves and subsidence patterns following the joint trends (Cooper, 1986). Most of the subcropping gypsum is alabastrine in the area around Ripon, while farther to the east, where the unit is thicker and deeper, the calcium sulphate phase is still anhydrite.

Fluctuations in the watertable level tied to heavy rain or long drought are thought to be the most common triggering mechanism for subsidence transitioning to sinkhole collapse. Many of the more catastrophic collapses occur after river flooding and periods of prolonged rain, which tend to wash away cavern roof span support. Subsidence is also aggravated by groundwater pumping; first, it lowers the watertable and second, it induces considerable crossflow of water in enlarged joints in the gypsum. When recharged by a later flood, the replacement water is undersaturated with respect to gypsum.


Thomson et al. (1996) recognised four hydrogeological flow units driving karst collapse in the Ripon area (Figure 4):

1) Quaternary gravels in the buried valley of the proto-River Ure

2) Sherwood Sandstone Group

3) Magnesian Limestone of the Brotherton Fm. and the overlying/adjacent gypsum of the Roxby Fm.

4) Magnesian limestone of the Cadeby Fm. plus the overlying/adjacent gypsum of the Edlington Fm.

Local hydrological base level within this stratigraphy is controlled by the River Ure, especially where the buried Pleistocene valley (proto-Ure) is filled by permeable sands and gravels, as these unconsolidated sediments, when located atop a breached roof beam, are susceptible to catastrophic stoping to base level (Figure 4). In the area around Ripon the palaeovalley cuts down more than 30 m, reaching levels well into the Cadeby Formation, so providing the seepage connections or pathways between waters in all four units wherever they intersect the palaeovalley. There is considerable groundwater outflow along this route with artesian sulphate-rich springs issuing from Permian strata in contact with Quaternary gravels of the buried valley (Cooper, 1986, 1995, 1998).

The potentiometric head comes from precipitation falling on the high ground of the Cadeby formation to the west and the Sherwood Sandstone to the east. Groundwater becomes largely confined beneath glacial till as it seeps toward the Ure Valley depression, but ultimately finds an exit into the modern river via the deeply incised sand and gravel-filled palaeovalley of the proto-Ure. Waters recharging the Ure depression pass through and enlarge joints and caverns in the gypsum units of the Edlington and Roxby Formations, so the highest density of subsidence features are found atop the sides of the palaeovalley. This region has the greatest volume of artesian discharge from aquifers immediately beneath the dissolving gypsum bed. Although created as an active karst valley, the apparent density of subsidence hollows is lower on the present Ure River floodplain than the surrounding lands as floodplain depressions are constantly filled by overbank sediments (Figure 4b).

Cooper (1998) defined 16 sinkhole variations in the gypsum subsidence belt at Ripon, all are types of entrenched, subjacent and mantled karst. Changes in karst style are caused by; the type of gypsum, the nature and thickness of the overlying deposits, presence or absence of consolidated layers overlying the gypsum and the size of voids/caverns within the gypsum.

To the west of Ripon, the gypsum of the Edlington Formation lies directly beneath glacial drift. These unconsolidated drift deposits and the loose residual marl atop the dissolving gypsum gradually subside into a pinnacle or suffusion (mantled) karst. But between Ripon town and the River Ure, the limestone of the Brotherton Formation overlies the Edlington Formation. There the karst develops as large open caverns beneath strong roof spans (entrenched karst). Ultimate collapse of the roof span creates rapid upward-stoping caverns in loosely consolidated sediment. Stopes break though to the surface as steep-sided collapse dolines or chimneys with sometimes catastrophic results. A similar entrenched situation is found east of the Ure River but there karstified gypsum units of both the Edlington and the Roxby formations are involved.


There are also thick beds of gypsum in the Permian Zechstein sequence that forms the bedrock in the Darlington area. In this area, subsidence features attributed to gypsum dissolution are typically broad shallow depressions up to 100 m in diameter, and the ponds, known as Hell Kettles, are the only recognized examples of steep-sided subsidence hollows around Darlington (Figure 5). Historical records suggest that one of the ponds formed in dramatic fashion in AD 1179 (Cooper 1995). The southern pond appears to be the most likely one to have formed at that time because it is many metres deep and is fed from below by calcareous spring water that is rich in both carbonate and sulphate. The 2D profiles have revealed evidence of foundering in the limestone of the Seaham Formation at depths of c. 50 m (Figure 5; Sargent and Goulty, 2009). The foundering is interpreted to have resulted from dissolution of gypsum in the Hartlepool Anhydrite Formation at ≈ 70 m depth. The reflection images of the gypsum itself are discontinuous, suggesting that its top surface has karstic topography. The 3D survey also acquired and interpreted by Sargent and Goulty (2009) reveals subcircular hollows in the Seaham Formation up to 20 m across, which are again attributed to foundering caused by gypsum dissolution.


Problems with Miocene gypsum, Spain

Karstification has led to problems in areas of subcropping Miocene gypsum in the Ebro and Calatayud basins, northern Spain (Figure 6). Cliff sections and road cuts indicate the widespread nature of karstification in the gypsum outcrops and subcrops in Spain (Figure 7b) Areas affected are defined by subsidence or collapse in Quaternary alluvial overburden and include; urban areas, communication routes, roads, railways, irrigation channels and agricultural fields (Figure 7a; Soriano and Simon, 1995; Elorza and Santolalla, 1998; Guerrero et al., 2013; Gutiérrez et al., 2014). In the region there can be a reciprocal interaction between anthropic activities and sinkhole generation, whereby the ground disturbance engendered by human activity accelerates, enlarges and triggers the creation of new sinkholes. Subsidence is particularly harmful to linear constructions and buildings and numerous roads, motorways and railways have been damaged (Figure 7a, b). Catastrophic collapse and rapid karst chimneying into roads and buildings can have potentially fatal consequences. For example, several buildings have been damaged around the towns of Casetas and Utebo. In the Portazgo industrial estate some factories had to be pulled down due to collapse-induced instability (Castañeda et al., 2009). A nearby gas explosion was attributed to the breakage of a gas pipe caused by subsidence. The local water supply is also disrupted by subsidence and pipe breakage so that 20,000 inhabitants periodically lose their water supply. The most striking example of subsidence affecting development comes from the village of Puilatos, in the Gallego Valley. In the 1970's this town was severely damaged by subsidence and abandoned before it could be occupied (Cooper 1996).


Collapse affects irrigation channels in the countryside with substantial economic losses (Elorza and Santolalla, 1998). In 1996 a doline collapse surfaced and cut the important Canal Imperial at Gallur village. New dolines often form near unlined irrigation canals. The ongoing supply of fresh irrigation waters to field crops can also encourage sinkhole generation in the fields. Though not directly visible, natural sinkholes also form in the submerged beds of river channels cutting regions of subcropping gypsum.

On December 19th, 1971, a bus fell from a bridge into the Ebro River at Zaragoza, near where the ‘San Lazaro well’ (a submerged gypsum sinkhole) is located (Figure 8a). Ten people lost their lives in this accident , while the remainder of the passengers were rescued, after being stranded on the bus roof in the flowing river for some hours (Figure 8b). After survivors were rescued, river waters washed the bus from the foot of the bridge supports into the nearby 'San Lazaro well (collapse sinkhole) in the water-covered floor of the river. Nine of the ten bodies in the bus were never found, although the bus was later recovered from the sinkhole. Locals suggested that bodies were carried deeper into the various interconnect phreatic sinkhole caverns fed by this losing stream.


Karstification in the Zaragoza region is characterised by the preferential intrastatal dissolution of glauberite bed, which are more soluble than the gypsum interbeds, this leads to collapse and rotation of gypsum blocks and river capture (Guerrero et al., 2013).

Sometimes even well-intentioned attempts to remediate culturally significant buildings under threat of evaporite karst collapse can exacerbate collapse problems. Gutiérrez and Cooper (2002) cite examples from the city of Calatayud, Spain. Subsidence-induced differential loading across doline edges drives the tilting of the 25-metre high tower (mudéjar) of the San Pedro de Los Francos church, which leans towards and overhangs the street by about 1.5 metres. (Figure 9) In places, the brickwork of the church indents the pre-existing tower fabric, which probably dates from the 11th Century or the beginning of the 12th Century. This indentation and the non-alignment of the church and the tower walls indicates that most of the tower tilting occurred prior to the construction of the church. In 1840, the upper 5m of the tower was removed and the lower part buttressed for the safety of the Royal family, who visited the town and stayed in the palace opposite. On 3rd June 1931, San Pedro de Los Francos church was declared a “Monument of Historical and Artistic value.” Due to its ruinous condition, the church was closed to worship in 1979. Micropiling to improve the foundation was started in 1994, but this corrective measure was interrupted when only half of the building was underpinned. Very rapid differential settlement of the building took place in the following year, causing extensive damage and aggravating the subsidence problem.


Colegiata de Santa María la Mayor was constructed between the 13th and 18th centuries, it has an outstanding Mudéjar (a 72 m high tower) and numerous Renaissance features; it is considered the foremost monument in the city of Cataluyud. As with the San Pedro de los Francos Church, recent micropiling work, applied to only one part of the cloister, has been followed by alarming differential movements that have drastically accelerated the deterioration of the building. Large blocks have fallen from the vault of the “Capitular Hall” and cracks up to 150 mm wide have opened in the brickwork of the back (NW) elevation, which has now been shored up for safety. The dated plaster tell-tales placed in these cracks to monitor the displacement demonstrate the high speed of the deformation produced by subsidence in recent years. On the afternoon of 10 September 1996, the fracture of a water supply pipe flooded the cloisters and the church with 100 mm of muddy water. Ten years earlier a similar breakage and flood had occurred. These breaks in the water pipes are most likely related to karst-induced subsidence. Once they occur, the massive input of water to the subsurface may trigger further destruction via enhanced dissolution, piping and hydrocollapse (Gutiérrez and Cooper, 2002).


Gypsum karst in Mosul, Iraq

A similar quandary of multiple areas of structural damage from gypsum-induced subsidence affects large parts or the historic section of the city of Mosul in northern Iraq (Jassim et al., 1997). The main part of its old quarter is over a century old and some buildings are a few hundred years old. Mosul lies on the northeastern flank of the Abu Saif anticline and near to its northern plunge (Figure 10a). It was built on the western bank of the Tigris River on a dip slope of Middle Miocene Fatha limestone that is directly underlain by bedded gypsum and green marl (equivalent to Lower Fars Formation). Houses in the old city were built on what seemed to be at the time a very sound rock foundation.

Water distribution in the city was done on mule back in the early part of last century and the estimated water consumption did not exceed 10 litres per person per day (Jassim et al., 1997). Discharge from households was partly to surface drainage and partly to shallow and small septic tanks. The modern piped system of water distribution did not start until the 1940s, resulting in a sudden increase in water consumption (presently around 200 litres per person per day) and it was not associated with a complementary sewer system. Increased water consumption meant larger and deeper septic tanks were dug at the perimeter of buildings (which never seemed to fill) resulting in a dramatic increase in water percolating downwards, water that was also more corrosive than previously due to the increased use of detergents and chlorination. This water passes through the permeable and fractured limestone to the underlying gypsum. On its way through the limestone it enlarges and creates new dissolution cavities, but eventually finds its way into the older gypsum karst maze, which is then further widened as water drains back into the Tigris (Figure 10b). Caverns in the gypsum enlarge until the roof span collapses. Since the 1970s more and more buildings in the old city have fractured and many are subject to sudden collapse. The problem is further intensified due to the expansion of the city in the up-dip direction (west and southwest) including the construction of industrial, water-dependent centres with integrated drainage. Water seeping/draining from these newly developed up-dip areas eventually passes under the old city before discharging in the Tigris river. The process was slightly arrested in the 1980s by the completion of a drainage system for the city, but the degradation of the old city continues.

Coping: man-made structures atop salts

The towns of Ripon in the UK and Pasvales and Birzai in Lithuania house some 45,000 people, who currently live under the ongoing threat of catastrophic subsidence, caused by natural gypsum dissolution (Paukstys et al., 1999). Special measures for construction of houses, roads, bridges and railways are needed in these areas and should include: incorporating several layers of high tensile heavy duty reinforced plastic mesh geotextile into road embankments and car parks; using sacrificial supports on bridges so that the loss of support of any one upright will not cause the deck to collapse; extending the foundations of bridge piers laterally to an amount that could span the normal size of collapses; and using ground monitoring systems to predict areas of imminent collapse (Cooper 1995, 1998).


Dams to store urban water supplies are costly structures and failure can lead to disaster, large scale mortality and financial liability (for example, Cooper and Gutiérrez, 2013). For example, at two and a half minutes before midnight on March 12, 1928, the St. Francis Dam (California) failed catastrophically and the resulting flood killed more than 400 people (Figure 11). The collapse of the St. Francis Dam is considered to be one of the worst American civil engineering disasters of the 20th century and remains the second-greatest loss of life in California’s history, after the 1906 San Francisco earthquake and fire. The collapse was partly attributed to dissolution of gypsum veins beneath the dam foundations. The Quail Creek Dam, Utah, constructed in 1984 failed in 1989, the underlying cause being an unappreciated existence of, and consequent enlargement of, cavities in the gypsum strata beneath its foundations.

Unexpected water leakage from reservoirs, via ponors, sinkholes and karst conduits, leads to costly inefficiency, or even project abandonment. Unnaturally high hydraulic gradients, induced by newly impounded water, may flush out of the sediment that previously blocked karst conduits. It can also produce rapid dissolutional enlargement of discontinuities, which can quickly reach break-through dimensions with turbulent flow. These processes may significantly increase the hydraulic permeability in the region of the dam foundation, on an engineering time scale.

Accordingly, numerous dams in regions of the USA underlain by shallow evaporites either have gypsum karst problems, or have encountered gypsum-related difficulties during construction (Johnson, 2008). Examples include; the San Fernando, Dry Canyon, Buena Vista, Olive Hills and Castaic dams in California; the Hondo, Macmillan and Avalon dams in New Mexico; Sandford Dam in Texas; Red Rock Dam in Iowa; Fontanelle Dam in Oklahoma; Horsetooth Dam and Carter Dam in Colorado and the Moses Saunders Tower Dam in New York State. Up to 13,000 tonnes of mainly gypsum and anhydrite were dissolved from beneath a dam in Iraq in only six months causing concerns about the dam stability (Figure 13). In China, leaking dams and reservoirs on gypsum include the Huoshipo Dam and others in the same area. The Bratsk Dam in eastern Siberia is leaking, and in Tajikistan the dam for the Nizhne-Kafirnigansk hydroelectric scheme was designed to cope with active gypsum dissolution occurring below the grout curtain. Gypsum karst in the foundation trenches of the Casa de Piedra Dam, Argentina and El Isiro Dam in Venezuela, caused difficult construction conditions and required design modifications.


Another illustration of the problems associated with water retaining structures and the ineptitude, or lack of oversight, by some city planners comes from the town of Spearfish, South Dakota (Davis and Rahn, 1997 ). As discussed earlier in this chapter, the Triassic Spearfish Formation contains numerous gypsum beds in which evaporite-focused karst landforms are widely documented across its extent in the Black Hills of South Dakota (Figure 12). The evaporite karst in the Spearfish Fm. has caused severe engineering problems for foundations and water retention facilities, including wastewater stabilization sites. One dramatic example of problems in water retention atop gypsum karst comes from the construction in the 1970s of now-abandoned sewage lagoons for the City of Spearfish.

Despite warnings from local ranchers, the Spearfish sewage lagoons were built in 1972 by city authorities on alluvium atop thick gypsum layers of Spearfish Formation. Ironically, at one point during lagoon construction, a scraper became stuck in a sinkhole and required four bulldozers to pull it out. Once filled with sewage, within a year the lagoons started leaking badly; the southern lagoon was abandoned after four years because of ongoing uncontrollable leaks, and the northern lagoon did not completely drain, but could not provide adequate retention time for effective sewage treatment. Attempts at repairs, including a bentonite liner, were ineffective, and poorly treated sewage discharged beneath the lagoon’s berm into a nearby surface drainage. The lagoons were abandoned completely in 1980. This was after a US $27-million lawsuit was filled in 1979 by ranchers whose land and homes were affected by leaking wastewater. A mechanical wastewater treatment plant was constructed nearby on an outcrop of the non-evaporitic Sundance Formation. The engineering firm that designed the facility without completing a knowledgeable geological site survey was reorganised following the lawsuit.

Likewise, the development of Chamshir Dam atop Gascharan Formation outcrop and subcrop in Iran is likely to create ongoing infrastructure cost and water storage problems (Torabi-Kaveh et al., 2012). The site is located in southwest of Iran, on Zuhreh River, 20 km southeast of Gachsaran city. The area is partially covered by evaporite formations of the Fars Group, especially the Gachsaran Formation. The dam axis is located on limestone beds of Mishan Formation, but nearly two-thirds of the dam reservoir is in direct contact with the evaporitic Gachsaran Formation. Strata in the vicinity of the reservoir and dam site have been brecciated and intersected by several faults, such as the Dezh Soleyman thrust and the Chamshir fault zone, which all act in concert to create karst entryways, including local zones of suffusion karst. A wide variety of karstic features typify the region surrounding the dam site and include; karrens, dissolution dolines, karstic springs and cavities. These karst features will compromise the ability of Chamshir Dam to store water, and possibly even cause breaching of the dam, via solution channels and cavities which could allow significant water flow downstream of the dam reservoir. As possible and likely partial short term solutions, Torabi-Kaveh et al. (2012) recommend the construction of a cutoff wall and/or a clay blanket floor to the reservoir

Difficulties in building hydraulic structures on soluble rocks are many, and dealing with them greatly increases project and maintenance costs. Gypsum dissolution at the Hessigheim Dam on the River Neckar in Germany has caused settlement problems in sinkholes nearby. Site investigation showed cavities up to several meters high and remedial grouting from 1986 to 1994 used 10,600 tonnes of cement. The expected life of the dam is only 30-40 years, with continuing grouting required to keep it serviceable.

Grouting costs in zones of evaporite karst can be very high and may approach 15 or 20% of the dam cost, currently reaching US$ 100 million in some cases. In karstified limestones grouting is difficult, yet in gypsum it is even more difficult due to the rapid dissolution rate of the gypsum. Karst expansion in limestone occurs on the scale of hundreds of years, in gypsum it can be on the order of a decade or less. Grouting may also alter the underground flow routes, so translating and focusing the problems to other nearby areas. In the Perm area of Russia, gypsum karst beneath the Karm hydroelectric power station dam has perhaps been successfully grouted, a least in the short term, using an oxaloaluminosilicate gel that hardens the grout, but also coats the gypsum, so slowing its dissolution. The Mont Cenis Dam, in the French Alps, is not itself affected by the dissolution of gypsum. However, the reservoir storage zone is leaking and photogrammetric study of the reservoir slopes showed ongoing doline activity over gypsum and subsidence in the adjacent land.


Probably the worst example tied to and evaporite karst hazard is the significant dam disaster waiting to happen that is the Mosul Dam in Iraq (Figure 13; Kelley et al., 2007; Sissakian and Knutsson, 2014; Milillo et al., 2016). It is ranked as the fourth largest dam in the Middle East, as measured by reserve capacity, capturing snowmelt from Turkey, some 70 miles (110 km) north. Built under the despotic regime of Saddam Hussein, completed in 1984 the Mosul Dam (formerly known as Saddam Dam) is located on the Tigris river, some 50 km NW of Mosul.

The design of the dam was done by a consortium of European consultants (Sissakian and Knutsson, 2014), namely, Swiss Consultants group, comprising: Motor Columbus; Electrowatt; Suiselectra; Societe Generale pour l’Industrie. The construction was carried out by a German-Italian consortium of international contractors, GIMOD joint venture, comprising: Hochtief; Impregilo; Zublin; Tropp; Italstrade; Cogefar. The consultants for project design and construction supervision comprised a joint venture of the above listed Swiss Consultants Group and Energo-Projekt of Yugoslavia, known as MODACON.

As originally constructed the dam is 113 m in height, 3.4 km in length, 10 m wide in its crest and has a storage capacity of 11.1 billion cubic meters (Figure 13b). It is an earth fill dam, constructed on evaporitic bedrock atop a karstified high created by an evaporite cored anticline in the Fat’ha Formation, which consists of gypsum beds alternating with marl and limestone (Figure 13a, 14). To the south, this is same formation with the same evaporite cored anticlinal association that created all the stability problems in the city of Mosul (Figures 10). The inappropriate nature of the Fat’ha Formation as a foundation for any significant engineering structure had been known for more than a half a century. Then again, absolute rulers do not need to heed scientific advice or knowledge. Or perhaps he didn’t get it from a well-paid group of Swiss-based engineering consultants. As Kelley et al. (2007) put it so succinctly....“The site was chosen for reasons other than geologic or engineering merit.”

The likely catastrophic failure of Mosul Dam will drive the following scenario (Sissakian and Knutsson, 2014); “... (dam) failure would produce a flood wave crest about 20 m deep in the City of Mosul. It is estimated that the leading edge of the failure flood wave would arrive in Mosul about 3 hours after failure of the dam, and the crest of the flood wave would arrive in Mosul about 9 hours after failure of the dam. The total population of the City of Mosul is about 3 million, and it is estimated that about 2 million people are in locations within the city that would be inundated by a 20 m deep flood wave. The City of Baghdad is located about 350 km downstream of Mosul Dam, and the dam failure flood wave will arrive after 72 hours in Baghdad and (by then) would be about 4 m deep.”



The heavily karsted Fat’ha Formation is up to 352 m thick at the dam and has an upper and lower member. The lower member is dominated by carbonate in its lower part (locally called “chalky series”) and is in turn underlain by an anhydrite bed known as the GBo. Gypsum beds typify its upper part,and the evaporite interval is capped by a limestone marker bed. The upper member, crops out as green and red claystone with gypsum relicts, around the Butmah Anticline. Thickness of individual gypsum beds below the dam foundations can attain 18 m; these upper member units are intensely karstified, even in foundation rocks, with cavities meters across documented during construction of the dam (Figure 14). Gypsum breccia layers are widespread within the Fatha Formation and have proven to be the most problematic rocks in the dam’s foundation zone. The main breccia body contains fragments or clasts of limestone, dolomite, or larger pieces of insoluble rocks of collapsed material. The upper portion of the accumulation grades upward from rubble to crackle mosaic breccia and then a virtually unaffected competent overburden. Breccia also may form without the intermediate step of an open cavity, by partial dissolution and direct formation of rubble. As groundwater moves through the rubble, soluble minerals are carried away, leaving insoluble residues of chert fragments, quartz grains, silt, and clay in a mineral matrix. These processes result in geologic layers with lateral and vertical heterogeneity on scales of micro-meters to meters.

High permeability zones in actively karsting gypsum regions can form rapidly, days to weeks, and quickly become transtratal. So predicting or controlling breakout zones via grouting and infill can be problematic (Kelley et al., 2007; Sissakian and Knutsson, 2014). For example, four sinkholes formed between 1992 and 1998 approximately 800 m downstream in the maintenance area of the dam (Figure 13a). The sinkholes appeared in a linear arrangement, approximately parallel to the dam axis. Another large sinkhole developed in February 2003, east of the emergency spillway when the pool elevation was at 325 m. The Mosul Dam staff filled the sinkhole the next day, with 1200 m3 of soil. Another sinkhole developed in July 2005 to the east of the saddle dam. Six borings were completed around the sinkhole and indicated that the sinkhole developed beneath overburden deposits and within layers of the Upper Marl Series. Another cause for concern at Mosul Dam in recent years is a potential slide area reported upstream of the dam on the west bank. The slide is most likely related to the movement of beds of the Chalky Series over the underlying GBo (anhydritic) layer.

To “cope” with ongoing active karst growth beneath and around the Mosul Dam, a continuous grouting programme was planned, even during dam construction, and continues today, on a six days per week basis. It pumps tens of thousands of tons of concrete into expanding karst features each year (Sissakian and Knutsson, 2014; Milillo et al., 2016). The dam was completed in June 1984, with a postulated operational life of 80 years. Due to insufficient grouting and sealing in and below the dam foundation, numerous karst features, as noted above, continue to enlarge in size and quantity, so causing serious problems for the ongoing stability of the dam. The increase in hydraulic gradient created by a wall of water behind the dam has accelerated the rate of karstification in the past 40 years.

Since the late 1980s, the status of the dam and its projected collapse sometime within the next few decades has created ongoing nervousness for the people of Mosul city and near surroundings. All reports on the dam since the mid 1980s have underlined the need for ongoing grouting and monitoring and effective planning of the broadcasting of a situation where collapse is imminent. For “Saddam’s dam” the question is not if, but when, the dam will collapse. To alleviate the effects of the dam collapse, Iraqi authorities have started to build another “Badush Dam” south of Mosul Dam so that it can stop or reduce the effects of the first flood wave. However this new dam has a projected cost in excess of US$ ten billion and so lies beyond the financial reach of the current Iraqi government. Problems related to the dam increased with the takeover of the region by the forces of ISIL.

Today, the Mosul dam is subsiding at a linear rate of ~15 mm/year compared to 12.5 mm/year subsidence rate in 2004–2010 (Milillo et al., 2016). Increased subsidence restarted at the end of 2013 after re-grouting operations slowed and at times stopped. The causes of the observed linear subsidence process of the dam wall can be found in the human activities that have promoted the evaporite–subsidence development, primarily in gypsum deposits and may enable, in case of continuous regrouting stop, unsaturated water to flow through or against evaporites deposits, allowing the development of small to large dissolution cavities.

Large vertical movements that typified the dam wall have resulted from the dissolution of extensive gypsum strata previously mapped beneath the Mosul dam. Increased subsidence rate over the past five years has been due to periods when there was little or no regrouting underlying the dam basement. Dam subsidence currently seems to follow a linear behavior but on can not exclude a future acceleration due to increased gypsum dissolution speed and associated catastrophic collapse of the dam (Milillo et al., 2016).

Given the existing geologic knowledge base in the 1980s, in my opinion, one must question the seeming lack of understanding in a group of well-paid consultant engineering firms as to the outcome of building such a major structure, atop what was known to be an active karstifying gypsum succession, sited in a location where failure will threaten multimillion populations in the downstream cities. The same formation that constituted the base to the Mosul dam was known at the time to be associated with ground stability problems atop similar gypsum-cored anticlines in the city of Mosul to the south. Even more concerning to the project rationale should have been the large karst cavities in highly soluble gypsum that were encountered a number of times during feasibility and construction of the dam foundations (Figure 14). Or, perhaps, as Lao Tzu observed many centuries ago, “ ...So the unwanting soul sees what’s hidden, and the ever-wanting soul sees only what it wants.”

Canals, like dams, that leak in gypsum karst areas can trigger subsidence, which can be severe enough to cause retainment failure. In Spain, the Imperial Canal in the Ebro valley, and several canals in the Cinca and Noguera Ribagorzana valleys, which irrigate parts of the Ebro basin, have on numerous occasions failed in this way. Similarly, canals in Syria have suffered from gypsum dissolution and collapse of soils into karstic cavities. Canals excavated in such ground may also alter the local groundwater flow (equivalent to losing streams) and so accelerate internal erosion, or the dissolution processes and associated collapse of cover materials. In the Lesina Lagoon, Italy, a canal was excavated to improve the water exchange between the sea and the lagoon. It was cut through loose sandy deposits and highly cavernous gypsum bedrock, but this created a new base level, so distorting the local groundwater flow. The canal has caused the rapid downward migration of the cover material into pre-existing groundwater conduits, producing a large number of sinkholes that now threaten an adjacent residential area.

Pipelines constructed across karst areas are potential pollution sources and some may pose possible explosion hazards. The utilization of geomorphological maps depicting the karst and subsidence features allied with GIS and karst databases help with the grouting and management of these structures. In some circumstances below-ground leakage {Zechner, 2011 #26} from water supply pipelines can trigger severe karstic collapse events. Where such hazards are identified, such as where a major oil and gas pipeline crosses the Sivas gypsum karst in Turkey, the maximum size of an anticipated collapse can be determined and the pipeline strength increased to cope with the possible problems.


Solving the problem?

Throughout the world, be it in the US, Canada, the UK, Spain, eastern Europe, or the Middle East, it is a fact that weathering of shallow gypsum forms rapidly expanding and stoping caverns, especially in areas of high water crossflow, unsupported roof beams, and unconsolidated overburden and in areas of artificially confined fresh water. Rapid karst formative processes and mechanism will always be commonplace and widespread (Table 2). Resultant karst-associated problems can be both natural and anthropogenically induced or enhanced. It is fact that natural solution in regions of subcropping evaporites is always rapid, and even more so in areas where it is encouraged by human activities, especially increased cycling of water via damming, groundwater pumping, burst pipes, septic systems, agricultural enhancement and uncontrolled storm and waste water runoffs to aquifers.

Typically, the best way to deal with a region of an evaporite karst hazard is to map the regional extent of the shallow evaporite solution front and avoid it (Table 3). In established areas with a karst problem the engineering solutions will need to be designed around hazards that will typically be characterised by short-term onsets, often tied to rapid ground stoping/subsidence events and quickly followed by ground collapse. If man-made buildings of historical significance are to be restored and stabilized in such settings, perhaps it is better to wait until funds are sufficient to complete the job rather than attempt partial stabilization of the worst-affected portions of the feature. Significant infrastructure (including roads, canals and dams) should be designed to avoid such areas when possible or engineered to cope with and/or survive episodes of ground collapse.

A piecemeal approach to dealing with evaporite karst can intensify and focus water crossflows rather than alleviate them. In the words of Nobel prizewinner, Shimon Peres; “If a problem has no solution, it may not be a problem, but a fact - not to be solved, but to be coped with over time.”


References

Alberto, W., M. Giardino, G. Martinotti, and D. Tiranti, 2008, Geomorphological hazards related to deep dissolution phenomena in the Western Italian Alps: Distribution, assessment and interaction with human activities: Engineering Geology, v. 99, p. 147-159.

Amin, A., and K. Bankher, 1997b, Causes of land subsidence in the Kingdom of Saudi Arabia: Natural Hazards, v. 16, p. 57-63.

Amin, A. A., and K. A. Bankher, 1997a, Karst hazard assessment of eastern Saudi Arabia: Natural Hazards, v. 15, p. 21-30.

Biddle, P. G., 1983, Patterns of drying and moisture deficit in the vicinity of trees on clay soils: Geotechnique, v. 33, p. 107-126.

Castañeda, C., F. Gutiérrez, M. Manunta, and J. P. Galve, 2009, DInSAR measurements of ground deformation by sinkholes, mining subsidence, and landslides, Ebro River, Spain: Earth Surface Processes and Landforms, v. 34, p. 1562-1574.

Cooper, A. H., 1986, Subsidence and foundering of strata caused by the dissolution of Permian gypsum in the Ripon and Bedale areas, North Yorkshire: Harwood, Gill M., Smith, Denys B. The English Zechstein and related topics. Univ. Newcastle upon Tyne, Newcastle upon Tyne, United Kingdom. Geological Society Special Publications, v. 22, p. 127-139.

Cooper, A. H., 1995, Subsidence hazards due to the dissolution of Permian gypsum in England: Investigation and remediation, in B. F. Beck, ed., Karst Geohazards - Engineering and Environmental Problems in Karst Terrane. Proceedings of the fifth multidisciplinary conference on sinkholes and the environmental impacts of karst, Gatlinburg, Tennessee: Rotterdam, A.A. Balkema, p. 23-29.

Cooper, A. H., 1998, Subsidence hazards caused by the dissolution of Permian gypsum in England: geology, investigation and remediation, in J. G. Maund, and M. Eddleston, eds., Geohazards in Engineering Geology, v. 15: London, Geological Society, London, p. 265-275.

Cooper, A. H., and F. Gutiérrez, 2013, Dealing with gypsum karst problems: hazards, environmental issues, and planning, in J. F. Shroder, ed., Treatise on geomorphology, Elsevier, p. 451-462.

Cooper, A. H., and J. M. Saunders, 2002, Road and bridge construction across gypsum karst in England: Engineering Geology, v. 65, p. 217-233.

Cooper, A. H., and A. C. Waltham, 1999, Subsidence caused by gypsum dissolution at Ripon, North Yorkshire: Quarterly Journal of Engineering Geology, v. 32, p. 305-310.

Dahm, T., S. Heimann, and W. Bialowons, 2011, A seismological study of shallow weak micro-earthquakes in the urban area of Hamburg city, Germany, and its possible relation to salt dissolution: Natural Hazards, v. 58, p. 1111-1134.

Davis, A., and P. Rahn, 1997, Karstic gypsum problems at wastewater stabilization sites in the Black Hills of South Dakota: Carbonates and Evaporites, v. 12, p. 73-80.

Driscoll, R., 1983, The influence of vegetation on the swelling and shrinking of clay soils in Britain: Geotechnique, v. 33, p. 93-105.

Elorza, M. G., and F. G. Santolalla, 1998, Geomorphology of the Tertiary gypsum formations in the Ebro Depression (Spain): Geoderma, v. 87, p. 1-29.

Ford, D. C., 1997, Principal features of evaporite karst in Canada: Carbonates and Evaporites, v. 12, p. 15-23.

Frumkin, A., M. Ezersky, A. Al-Zoubi, E. Akkawi, and A.-R. Abueladas, 2011, The Dead Sea sinkhole hazard: Geophysical assessment of salt dissolution and collapse: Geomorphology, v. 134, p. 102-117.

Galve, J. P., F. Gutierrez, P. Lucha, J. Bonachea, J. Remondo, A. Cendrero, M. Gutierrez, M. J. Gimeno, G. Pardo, and J. A. Sanchez, 2009, Sinkholes in the salt-bearing evaporite karst of the Ebro River valley upstream of Zaragoza city (NE Spain) Geomorphological mapping and analysis as a basis for risk management: Geomorphology, v. 108, p. 145-158.

Garleff, K., H. Kugler, A. V. Poschinger, H. Sterr, H. Strunk, and G. Villwock, 1997, Germany, in C. Embleton, and C. Embleton, eds., Geomorphological hazards of Europe, Vol. 5. Developments in Earth Surface Processes, v. 5, p. 147-177.

Guerrero, J., F. Gutiérrez, and J. P. Galve, 2013, Large depressions, thickened terraces, and gravitational deformation in the Ebro River valley (Zaragoza area, NE Spain): Evidence of glauberite and halite interstratal karstification: Geomorphology, v. 196, p. 162-176.

Gutierrez, F., 2010, Hazards associated to karst (Chapter 13), in I. Alcántara-Ayala, and A. S. Goudie, eds., Geomorphological Hazards and Disaster Prevention, Cambridge University Press, p. 161-176.

Gutiérrez, F., 1996, Gypsum karstification induced subsidence - effects on alluvial systems and derived geohazards (Calatayud Graben, Iberian Range, Spain): Geomorphology, v. 16, p. 277-293.

Gutiérrez, F., 2014, Evaporite Karst in Calatayud, Iberian Chain, in F. Gutiérrez, and M. Gutiérrez, eds., Landscapes and Landforms of Spain: World Geomorphological Landscapes, Springer Netherlands, p. 111-125.

Gutiérrez, F., A. Cooper, and K. Johnson, 2008, Identification, prediction, and mitigation of sinkhole hazards in evaporite karst areas: Environmental Geology, v. 53, p. 1007-1022.

Gutiérrez, F., and A. H. Cooper, 2002, Evaporite dissolution subsidence in the historical city of Calatayud, Spain: Damage appraisal and prevention: Natural Hazards, v. 25, p. 259-288.

Gutiérrez, F., M. Parise, J. De Waele, and H. Jourde, 2014, A review on natural and human-induced geohazards and impacts in karst: Earth-Science Reviews, v. 138, p. 61-88.

Jassim, S. Z., A. S. Jibril, and N. M. S. Numan, 1997, Gypsum karstification in the Middle Miocene Fatha Formation, Mosul area, Northern Iraq: Geomorphology, v. 18, p. 137-149.

Johnson, K., 2008, Gypsum-karst problems in constructing dams in the USA: Environmental Geology, v. 53, p. 945-950.

Jones, C. J. F. P., and A. H. Cooper, 2005, Road construction over voids caused by active gypsum dissolution, with an example from Ripon, North Yorkshire, England: Environmental Geology, v. 48, p. 384-394.

Karacan, E., and I. Yilmaz, 1997, Collapse dolines in Miocene gypsum - An example from SW Sivas (Turkey): Environmental Geology, v. 29, p. 263-266.

Kelley, J. R., L. D. Wakeley, S. W. Broadfoot, M. L. Pearson, C. J. McGrath, T. E. McGill, J. D. Jorgeson, and C. A. Talbot, 2007, Geologic Setting of Mosul Dam and Its Engineering Implications: US Army Corps of Engineers; Engineer Research and Development Center Report ERDC TR-07-10.

Martinez, J. D., and R. Boehner, 1997, Sinkholes in glacial drift underlain by gypsum in Nova Scotia, Canada: Carbonates and Evaporites, v. 12, p. 84-90.

Milillo, P., R. Bürgmann, P. Lundgren, J. Salzer, D. Perissin, E. Fielding, F. Biondi, and G. Milillo, 2016, Space geodetic monitoring of engineered structures: The ongoing destabilization of the Mosul dam, Iraq: Nature Open Reports, v. 6, p. 37408.

Paukstys, B., A. H. Cooper, and J. Arustiene, 1999, Planning for gypsum geohazards in Lithuania and England: Engineering Geology, v. 52, p. 93-103.

Sargent, C., and N. R. Goulty, 2009, Seismic reflection survey for investigation of gypsum dissolution and subsidence at Hell Kettles, Darlington, UK: Quarterly Journal of Engineering Geology and Hydrogeology, v. 42, p. 31-38.

Shviro, M., I. Haviv, and G. Baer, 2017, High-resolution InSAR constraints on flood-related subsidence and evaporite dissolution along the Dead Sea shores: Interplay between hydrology and rheology: Geomorphology, v. 293, p. 53-68.

Sissakian, V., N. Al-Ansari, and S. Knutsson, 2014, Karstification Effect on the Stability of Mosul Dam and Its Assessment, North Iraq: Engineering and Mining Journal, v. 6, p. 84-92.

Sissakian, V. K., V. K. Al-Ansari, and S. Knutsson, 2015, Karst Forms in Iraq Journal of Earth Sciences and Geotechnical Engineering, v. 5, p. 1-26.

Soriano, M. A., and J. Simon, 1995, Alluvial dolines in the central Ebro basin, Spain: a spatial and developmental hazard analysis: Geomorphology, v. 11, p. 295-309.

Sprynskyy, M., M. Lebedynets, and A. Sadurski, 2009, Gypsum karst intensification as a consequence of sulphur mining activity (Jaziv field, Western Ukraine): Environmental Geology, v. 57, p. 173-181.

Stafford, K. W., W. A. Brown, T. Ehrhart. Jon, A. F. Majzoub, and J. D. Woodard, 2017, Evaporite karst geohazards in the Delaware Basin, Texas: review of traditional karst studies coupled with geophysical and remote sensing characterization: International Journal of Speleology, v. 46, p. 169-180.

Thierry, P., A. Prunier-Leparmentier, C. Lembezat, E. Vanoudheusden, and J. Vernoux, 2009, 3D geological modelling at urban scale and mapping of ground movement susceptibility from gypsum dissolution: The Paris example (France): Engineering Geology, v. 105, p. 51-64.

Tolmachev, V., A. Ilyin, B. Gantov, M. Leonenko, V. Khomenko, and I. A. Savarensky, 2003, The main results of engineering karstology research conducted in Dzerzhinsk, Russia (1952-2002), in B. Beck, ed., Sinkholes and the engineering and environmental impacts of karst: proceedings of the ninth multidisciplinary conference, September 6-10, 2003, Huntsville, Alabama, American Society of Civil Engineers, p. 502-516.

Tolmachev, V., and M. Leonenko, 2011, Experience in Collapse Risk Assessment of Building on Covered Karst Landscapes in Russia, in P. E. van Beynen, ed., Karst Management, Springer Netherlands, p. 75-102.

Torabi-Kaveh, M., M. Heidari, and M. Miri, 2012, Karstic features in gypsum of Gachsaran Formation (case study; Chamshir Dam reservoir, Iran): Carbonates and Evaporites, v. 27, p. 291-297.

Toulemont, M., 1984, Le karst gypseux du Lutetien superieur de la region parisienne; caracteristiques et impact sur le milieu urbain: Revue de Geologie Dynamique et de Geographie Physique, v. 25, p. 213-228.

Trzhtsinsky, Y., 2002, Human-induced activation of gypsum karst in the southern Priangaria (East Siberia, Russia): Carbonates and Evaporites, v. 17, p. 154-158.

Waltham, T., F. Bell, and M. Culshaw, 2005, Sinkholes and Subsidence: Karst and Cavernous Rocks in Engineering and Construction: Berlin Heidelberg, Springer Praxis Books, 382 p.

Wang, G., G. You, and Y. Xu, 2008, Investigation on the Nanjing Gypsum Mine Flooding, in H. Liu, A. Deng, and J. Chu, eds., Geotechnical Engineering for Disaster Mitigation and Rehabilitation: Proceedings of the 2nd International Conference GEDMAR08, Nanjing, China 30 May – 2 June, 2008: Berlin, Heidelberg, Springer Berlin Heidelberg, p. 920-930.

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

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

Yaoru, L., and A. H. Cooper, 1997, Gypsum karst geohazards in China, in B. F. Beck, and J. B. Stephenson, eds., Engineering Geology and hydrogeology of Karst Terrains: Proceedings of the Sixth Multidisciplinary Conference on Sinkholes and the Engineering and Environmental Impacts of Karst Springfield, Missouri, 6-9 April 1997, Balkema, Rotterdam, p. 117-126.

Yilmaz, I., M. Marschalko, and M. Bednarik, 2011, Gypsum collapse hazards and importance of hazard mapping: Carbonates and Evaporites, v. 26, p. 193-209.

Zechner, E., M. Konz, A. Younes, and P. Huggenberger, 2011, Effects of tectonic structures, salt solution mining, and density-driven groundwater hydraulics on evaporite dissolution (Switzerland): Hydrogeology Journal, v. 19, p. 1323-1334.

 


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