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.

 

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.

 

Dissolving salt (1 of 5): Landforms

John Warren - Thursday, August 31, 2017

 

Introduction

This series of five articles discuss effects of nearsurface evaporite dissolution at various scales and times, in terms of; 1) landform expression, 2) evaporite-hosted caves and speleothems 3) natural evaporite karst geohazards 4)Anthropogenically enhanced karst features, 5) economic associations tied to evaporite paleokarst (hydrocarbon and metal).

As salt dissolves in the subsurface, it creates void spaces of various sizes and shapes into which overlying strata can drape or brecciate, so creating characteristic landforms at both local (m to km) and regional (km to tens of km) scales (Figure 1). Water sources producing these features can be shallow unconfined meteoric or marine, confined shallow or deeply circulating meteoric, or hypogene, basinal and hydrothermal. Solution-related evaporitic landforms can be either active karst or palaeokarst features. Paleokarst refers to ancient karstic features that are no longer active and tied to ancient basin flushing hydrologies and now buried land surfaces. Active karst is responding to the modern ambient hydrology and typifies landscape elements discussed in the following section.


Evaporite solution karst landforms overprint nearby carbonate or evaporite strata, with many descriptive features and terms are common to textures and geometries across both host lithologies. Evaporite-related karst includes many varieties of karren, sinkholes (dolines and caves), blind valleys, poljes and subsidence basins (Ford and Williams, 2007; Warren, 2016). Bedrock exposures of gypsum or salt at a finer scale are sculpted into irregular curved and grooved surfaces called rillenkarren or other varieties of gravity-oriented solution flutes (Macaluso and Sauro, 1996; Stenson and Ford, 1993).

Local-scale karst geometries that define evaporite-related subsidence range in lateral extent from the metre-scale cones of suffosion karst to collapse cones to kilometre-scale doline depressions (Figure 1). In regions of soil-mantled outcrops of bedded evaporites, these features tend to occur at higher densities, within or near broad subsidence valleys. In the context of the lateral extents of evaporite-related landforms, it is also useful to separate landforms associated with the dissolution of bedded evaporite versus dissolution landforms and namakier residues that occur above variably active salt diapir crests. Features related to dissolving diapirs and namakiers are discussed in an earlier article (Salty Matters, March 10, 2015).

The geometry of the karst set-up hydrology is distinct in sediments above diapiric versus bedded evaporite substrate, and this article focuses on landforms related to a dissolving substrate made up of shallow dipping salt beds. The next section outlines characteristics feature of evaporite karst at the metre to kilometre scale, as defined in Figure 1. Then we will compartmentalise these features using the classification of Gutiérrez et al., 2008, and finally look at three of the more distinctive regions responding to the dissolution of subcropping evaporite beds.

Evaporite solution dolines

Dolines are closed circular to elliptical hollows or depressions, often funnel shaped, with diameters ranging from a few metres to a few kilometres and depths from a metre or so, to hundreds of metres (Figure 1; Ford and Williams, 2007). They indicate subsidence and collapse in underlying salt or carbonate units. Valley sides in larger subsidence dolines can be steep and expose karstified gypsum, or can be gentler slopes covered by soil.

Larger dolines may enclose one or more smaller sinkholes and can be further subdivided into, suffusion, subsidence sag, and collapse dolines (Figure 1). A collapse sinkhole is initiated by collapse of a roof span into an underlying solution cavity. A subsidence sag doline indicates a more diffuse broader and gentler lowering of the ceiling above the dissolving bed. Suffosion dolines are small-width high-density karst features (1-2m diameter) covered and filled with soil and debris that has washed or fallen into closely-spaced fissures cutting into the evaporite bed. They typically indicate that a dissolving salt mass is very close to the landsurface.

Dissolution, collapse and suffosion processes are more active, more rapid, more frequent and more noticeable in regions of shallow evaporite units, compared to carbonate terranes at the same depths. For shallow weathering evaporites versus carbonates, at roughly the same depths in the Perm and Bashkir regions of Russia, Gorbunova (1979) reported doline densities of 32 and 10/km2 respectively.


Suffosion dolines

Suffosion in a karst terrane describes the downward migration of unconsolidated cover deposits through voids. High levels of dispersed impurities (mostly clays and muds, often dolomitic and other dissolution residues) in a rapidly dissolving evaporite mass, means a soil or carapace of insoluble residues quickly covers the subcropping top and edges of a dissolving salt mass (Figures 1a, 2). Downwash transport into growing fissures and voids in the shallow underlying evaporite is via downwashing of fine particles carried by percolating waters, cohesionless granular flows, viscous sediment gravity flows (non-Newtonian), freefall of particles, and sediment-laden water. The carapace is continually undermined by ongoing rapid solution, as residual debris is continually and rapidly washed into the doline crevices into the growing cavities in the evaporite. This creates a unique soil-covered dimpled landscape, which is typified by a high-density doline terrain where the dissolving evaporite is only a metre or so below the landsurface.


Densities of up to 1000 suffosion dolines/km2 occur in many evaporite-cored fold axes or at subcrop contacts of evaporites with other outcropping lithologies. Densities of 1100-1500/km2 have been documented in evaporite karst in the Italian Alps and in regions to the west of Sivas, Turkey (Figure 3; Belloni et al., 1972; Kacaroglu et al., 1997). The inherently high solubility of evaporite salts explains similar densely-packed schlotten depressions and large karren shafts seen in gypsum karst of Antigonish County, Nova Scotia, where extrapolated doline densities in the latter case range up to 10,000/km2 (Martinez and Boeher, 1997). Extrapolated, because such dense networks do not extend over any more than a square kilometre or two and are typically found near retreating escarpment edges underlain by shallow subcropping salts. Such high densities tend to occur in more humid rather than arid areas, with high hydraulic gradients so that the resulting suffosion dolines have diameters around 5 m or less.


Subsidence Dolines

Solution (subsidence) dolines, as described in much of the literature tend to be larger down-warped doline craters or bowl-like subsidence depressions (Figure 1c). In contrast to the numerous small steep-sided suffosion and collapse dolines formed atop shallow subcropping shallow evaporites, these larger, bowl-like dolines can have diameters of 100-500 m and depression depths of 10-20 m or more. Most subsidence sag dolines have a well-developed soil cover and a thick sediment fill, with dissolving salt units found depths measured in tens to hundreds of metres below the land surface (Figure 4). Compared to collapse and suffusion dolines, subsidence or sag dolines have lower angle to near flat slopes into the deeper parts of the doline hollow, the doline walls at the outcrop level are usually hosted in non-evaporites, with the dissolving salt bed lying some distance below the landsurface. Compared to suffusion karst, subsidence dolines occur in regions of more deeply buried evaporites (tens of metres) including; Italy (Belloni et at. 1972; Burri 1986; Ferrarase et al., 2002), Spain (Gutiérrez, 1996) and the Pecos Valley of West Texas and New Mexico (Gustavson et al., 1982; Davies, 1984a, b; Quinlan et al., 1986). At the larger end of the scale of subsidence doline development, a subsidence doline merges into a subsidence basin (Figure 1d). The latter is a large solution hollow that had created enough accommodation space to be considered a small sedimentary basin, generally filling with varying combinations of fluvial lacustrine and other continental sediments.


Collapse dolines

Collapse dolines are steep-sided sinkholes, often defined by cave entrances that contain large blocks of roof material (Figure 1c). They form when solution of an underlying evaporite bed creates a roof span that can no longer be supported by the overlying lithology. Collapse doline walls are frequently asymmetrical; one wall is steep, and the other one is gentle (Figure 5). Soil covered doline floors, when not water covered, tend to display either concave-up or flat geometries. Apart from blocks of the collapsed roof span, active doline floors can be veneered by thin collapse breccias in a matrix of insoluble residues. The high density of dolines in areas of evaporite subcrop and the ubiquity of the associated breccias indicates the inherently higher solubility of the evaporite salts compared with interbedded and overlying carbonates. Active collapse sinkholes atop shallow bedded evaporites typify terminations of dry arroyos in many deserts and may also line up along active or former river courses, as in Bottomless Lake State Park, New Mexico (Figure 6).


Sinkhole Classification

Following the definitions of Waltham et al. (2005), Gutiérrez et al. (2008) constructed a nongenetic classification of subsidence dolines or sinkholes (Figure 7). It is based on two observations that refer to the material affected by downward gravitational movements (cover, bedrock or caprock) and the primary type of process involved (collapse, suffosion or sagging). The classification also applies to both evaporite and carbonate karst. The term “cover” refers to allogenic unconsolidated deposits or residual soil material, bedrock to karst rocks and caprock to overlying non-karst rocks. Collapse indicates the brittle deformation of soil or rock material either by brecciation or the downward migration deposits through conduits and its progressive settling, while sagging is the ductile flexure (bending) of sediments caused by the lack of basal support. In practice, more than one material type and several processes can be involved in the generation of sinkholes. These more complex sinkholes can be described using combinations of the proposed terms with the dominant material and/or process followed by the secondary one (e.g. cover and bedrock collapse sinkhole, bedrock sagging and collapse sinkhole).


The cover material may be affected by any of three subsidence mechanisms. The progressive corrosional lowering of the rockhead may cause the gradual settling of the overlying deposits by sagging, producing cover sagging sinkholes. An important applied aspect is that the generation of these sinkholes does not require the formation of cavities These depressions arc commonly shallow, have poorly defined edges and may reach several hundred metres across.

Cover deposits may migrate downward into fissures and conduits developed in the rockhead by action of a wide range of processes collectively designated as suffusion: downwashing of particles by percolating waters, cohesion-less granular flows, viscous gravity flows (non-Newtonian), fall of particles, and sediment-laden water flows. The downward transport of the cover material through pipes and fissures may produce two main types of sinkholes depending on the rheological behaviour of the mantling deposits. Where the cover behaves as a ductile or loose granular material, it may settle gradually as undermining by suffosion progresses. This creates sags and slumps in the overburden materials. Where cover behaves in a more brittle manner, collapse breccias form.

Sinkholes that result from the combination of several subsidence processes and affect more than one type of material are described by combinations of the different terms with the dominant material or process followed by the secondary one (e.g. bedrock sagging and collapse sinkhole). The mechanism of collapse includes any brittle gravitational deformation of cover and bedrock material, such as upward stoping of cavities by roof failure, development of well-defined failure planes and rock brecciation. They define suffosion as the downward migration of cover deposits through dissolutional conduits accompanied by ductile settling. Sagging is the ductile flexure of sediments caused by differential corrosional lowering of the rockhead or interstratal karstification of the soluble bedrock. Sag plays a major role in the generation of sinkholes across broad areas underlain by shallow dissolving evaporites (not included in previous genetic classifications mostly based on carbonate karst). Likewise, collapse processes are more significant in extent and rate in areas underlain by evaporites than in carbonate karst, primarily due to the order of magnitude greater solubility of the evaporites and the lower mechanical strength and ductile rheology of gypsum and salt rocks (Warren, 2016).


Broader-scale evaporite landforms

There are some unique aspects to regional landform and doline density associations above dissolving evaporites because; 1) halite, and to a lesser extent anhydrite, units are many times more soluble than carbonate or siliciclastic strata, and 2) matrix in most salt units away from its retreating edges tends to remain impervious even as salt masses approach the landsurface (Figure 8; melting block of ice analogy, see Warren, 2016). Such features form at the basin-scale in regions where an ancient salt bed is moving from the mesogenetic to the telogenetic realm, especially in regions where mountains are growing adjacent to the uplifted evaporite basin. The broader-scale landscape features not seen in carbonate karst in similar regional settings are; 1) laterally migrating subsidence basins (Figure 8a) and 2) regions of breccia chimneys (Figure 8b).

Subsidence basins

A subsidence basin is a large (tens of kilometres width) solution hollow that creates enough accommodation space to be considered a small sedimentary basin with widths measured in tens of kilometres (Figure 8a). In a continental setting, the basin is filling with varying combinations of fluvial and lacustrine sediments.

Subsidence troughs are large-scale elongate depositional depressions created by interstratal solution along the dissolving edge of shallow dipping ancient uplifted salt bodies. The largest solution-induced depositional basins tend to occur along the margins of the great interstratal halite deposits, creating a solution form that may be represented by a shallow retreating salt slope at the surface, with a laterally migrating monoclinal drape of beds in the vicinity of a salt scarp, which is defined the dissolving edge of the underlying salt bed or beds. Subsidence basins are filled or partly filled by clastic sediments (Olive, 1957; Quinlan, 1978; Simpson, 1988). If the subsidence zones atop a retreating salt edge lack a significant volume of sediment fill, it called is a subsidence trough, as seen in the vicinity of the outcropping edge of the Jurassic Hith Anhydrite in Saudi Arabia (see part 2).

Salt-edge leaching of shallow dipping salt underscores a set of self-perpetuating processes. Fractures created by the collapse of overburden and intrasalt beds contribute to an expanding accommodation space generated by salt cavities and the lateral retreat of the salt scarp. The ongoing salt dissolution provides new fissures and sinks that act as additional conduits for further percolation of meteoric or upwelling of undersaturated basinal waters into the salt. This, in turn, instigates yet more salt solution in the vicinity of the collapse and typically, so creating an elongate corridor of subsidence at the surface, which can extend for many kilometres parallel to the dissolving salt edge

Hence, large solution-induced depositional basins and monoclines outline the edges of the great interstratal salt deposits of the world. Depression corridors some 5-500 km long, 5-250 km wide, with up to 100 to 500 m of subsidence induced relief and sediment fill define saline karst plains along the edges of bedded and dissolving saline giants in the Devonian evaporite subcrop of Canada (Figure 9b; De Mille et al., 1964; Tsui and Cruden 1984; Christiansen and Sauer, 2002; Tozer et al., 2014; Broughton et al., 2017), the Permian Basin of New Mexico, Oklahoma and Texas (Figure 9a; Anderson, 1981; Davies, 1984a, b; Bachman, 1984), the Perm region of Russia (Gorbunova, 1979) and the Jurassic Hith region of Saudi Arabia (Amin and Bankher, 1997a, b;). Associated regolith and sediment infill typically reduces the regional solution edge landscape to a few tens of metres of relief.

Salt Breccia Chimneys

Undersaturated waters not only influence the dissolving edges of a shallow-dipping salt bed, but can also cut up through the salt as a series of salt chimneys, located well out in the salt basin, with diameters between 20 and 250 m. Most are plumb vertical, with reported heights ranging from tens of metres to kilometers. Higher examples have usually stoped upward through one or more cover formations, which may include siliciclastics, coal measures, extrusive volcanics, etc. (Figure 9b). Breccia pipes in an evaporite mass may exhibit one of four possible states (Figure 8b): a) Active, and propagating upwards towards the surface, with fault and collapse bounded edges, but not yet expressed at the surface - a blind chimney; b) Active or inactive, expressed at the surface as a closed depression or a depression with a surface outflow channel; c) Inactive, and buried by later strata - paleokarst chimney; d) Inactive and standing up as a positive relief feature on the landsurface, because the breccia (generally cemented) is more resistant to weathering than the upper cover strata - so forming a residual pipe with positive relief. Some upstanding features are firmly cemented and resistant structure, such as the castiles above an erosional plain in the Delaware Basin of Texas (Hill, 1990)

Downward flexure in the uppermost strata atop a growing pipe tends to form a bowl of subsidence and a sag doline structure.

In some intrasalt pipes, the halite or gypsum is entirely removed, leaving only an accumulation of insolubles and collapsed breccia blocks. In others, portions of the less soluble salts can remain in the pipe at the level of the mother salt. Mature pipes are typically filled by a jumble of intrasalt and suprasalt breccia blocks. In some, a lower brecciated zone is succeeded by an upper zone in which the overlying strata failed as a coherent block, settling downwards via a cylindrical pattern of steep to vertical faults with downthrows of up to 200 m. Such variably filled and zoned salt breccia chimneys have created one type of barren zone in the potash ores of the Prairie Evaporite in Canada, and we shall discuss this in detail in part 5(economic associations).

Most breccia pipes originate via a point-source dissolution breakthrough of a halite bed (occasionally gypsum or anhydrite) and are located above a fracture junction, an anticlinal crest, or a buried reef that channels groundwater into a local area in the salt (Figure 8b). Hence, salt breccia chimneys and pipes form best where there is an artesian head to create subsalt pressures. In the Western Canada Sedimentary Basin the head comes from the adjacent uplifted Rocky Mountains, while the Delaware and Guadalupe Mountains play a similar role in the Delaware Basin. A pipe creates anomalous chemical interfaces with adjacent and overlying strata and so may be targets for the later precipitation of economic ores, including uranium and base metals (part 5).


Delaware Basin, West Texas

The Delaware Basin of West Texas and southeast New Mexico, in the southwest portion of the Permian Basin, contains bedded evaporites of the Late Permian Castile, Salado and Rustler Formations (Figure 9a). Outcrops and subcrops of these three formations constitute an area of widespread subsidence troughs, collapse sinks, dolines and breccia chimneys, all created by ongoing removal of underlying bedded Permian salts. The Delaware is an eastward-dipping basin, mainly surrounded by the Capitan Reef and its equivalents (Figure 9a). The original extent of the evaporites in the basin was much greater than today due to ongoing salt dissolution.

The area called the “Gypsum Plain” of Texas lies to the west of the Pecos River and comprises about 2,600 square kilometres of subcropping gypsum of the Castile Formation (Olive, 1957; Quinlan, 1978; Kirkland and Evans, 1980; Anderson, 1981; Stafford et al., 2008a,b; Holt and Powers, 2010). South of Carlsbad and Carlsbad Caverns, the plain exhibits many small examples of solution subsidence troughs, typically 0.7–15 km in length, 100–1500 m wide, but no more than 5–10 m deep (Figure 9a: Stafford et al., 2008a). The plain is also where the “Castile” landforms are found and indicate an overprint of bacterial and thermochemical sulphate reduction (Kirkland and Evans, 1976). Additional gypsum outcrops are present to the east in the Rustler Hills and to the north in Reeves County, Texas. Permian strata in region of the "Gypsum Plain" once contained significant volumes of halite that were dissolved well before the evaporite succession reached the landsurface.


In the centre of the Delaware Basin, the thick evaporites of the Castile and Salado Formations retain their halite beds as do the thinner evaporites of the overlying Rustler Formation. All are underlain by relatively permeable carbonates and siliciclastics, including some prolific hydrocarbon reservoirs in Permian backreef of the Central Basin Platform (Figure 10). Reef mounds, fractures and faults in these underlying sediments have provided focused conduits for upward stoping breccia chimneys through the buried evaporites as well as the subsurface formation of now-exhumed Castiles to the west. An eastward-flowing deeply-circulating regional artesian hydrology, in combination with centripetal escape of buoyant hydrocarbon-rich basinal waters, drives the formation of these chimneys, with their surface expressions occurring in areas such as the Wink Sink (chain of breached chimneys) and the Gypsum Plain (Castiles).

The upper sides of the shallow dipping salt beds are also affected by the hydrologies of the zone of active phreatic circulation. The Pecos River has migrated back and forth across the top of subcropping evaporites for much of the Tertiary. Its ancestral positions drove substantial salt dissolution, now evidenced by large sediment-filled subsidence basins and troughs in the centre of the Delaware Basin (Figures 9b, 10). The regional eastward dip of the Delaware basin sediments means first halite, and then gypsum has disappeared along the updip eastern edge of the basin. Relatively undisturbed salt remains along the more deeply buried western side of the basin that abuts and covers the Central Basin Platform. Bachman and Johnson (1973) estimate the horizontal migration rate of the dissolution front across the basin as high as 10-12 km per million years so that more than 50% of the original halite is gone. Multiple smaller examples of sediment-filled subsidence troughs occur at the edge of the gypsum plain of the Delaware Basin south of Carlsbad Caverns, where depositional troughs, 0.7 to 15 km in length, 100 to 1,500 m wide and no more than 5-10 m deep, are well documented, as are subsidence sinks within the subsidence swales, such as Bottomless Lake, which is a region where the watertable intersects a collapse chimney (Figure 10c; Quinlan et al. 1986).

The San Simon swale is a 25 km2 depression defining a residual karst feature atop the Capitan Reef on the northeastern margin of the Basin (Figures 9a, 10a, b). San Simon Sink sits atop a subsidence chimney within the San Simon swale; it is the lowest point in the depression and is some 30 m deep and 1 km2 in area. It, in turn, encloses a secondary collapse sink some a few hundred metres across and 10 metres deep (Figure 10b). During a storm in 1918, the San Simon sinkhole formed as a gaping hole about 25 metres across and 20 metres deep in the lower part of the sink. In one night, nearly 23,000 cubic metres of soil and bedrock disappeared into the collapse cavern. Annular rings that cut the surface around the San Simon sinkhole today suggest ongoing subsidence and readjustment of the sinkhole is still occurring in response to earlier collapses. The position of the San Simon sink over the Permian reef crest led Lambert (1983) and many others to suggest that the sinkhole originated as a groundwater cavity breakthrough, atop a series of stoping reef-focused breccia pipes or chimneys. Sinkhole breakouts, which can emerge in a matter of hours, continue to form across this dissolution basin, in some case aided by poorly-monitored brine extraction operations and improperly-cased water wells. But the majority of the sinkholes in the Delaware Basin are natural, not anthropogenically enhanced (Land, 2013).


Nash Draw is a southwesterly trending depression or swale, some 25 km long and 5-15 km wide, at the northern end of the Delaware Basin with its sump in a salt lake (Poker Lake) at the southern end of the draw (Figure 11). The underlying evaporitic Rustler Formation and parts of the Salado Formation have largely dissolved so that more than 100 caves, sinks, fractures, swallow-holes, and tunnels make up a complex local karst topography in the Draw, which is still active today (Figure 11b, c; Bachman, 1981, Powers et al., 2006; Goodbar and Goodbar, 2014).

An extensive drilling program conducted for the nearby WIPP site (now a low-level radioactive waste repository) showed that natural dissolution of halite in the Rustler and upper Salado formations is responsible for the subsidence and overall formation of Nash Draw (Lambert, 1983; Holt and Powers, 2010). To the immediate west of Nash Draw, the WIPP/DOE drilling program defined the formation of a solution trough in the Dewey Lake Redbed; it was created by preferential leaching of halite beds in the Rustler Formation, with interstratal anhydrite and breccia residuals (Figure 12). This dissolution occurred at a level some 400 metres above the salt-encased storage level of the WIPP waste isolation facility and so is not considered a significant risk factor in terms of longterm site stability (Holt and Powers, 2010). A heated scientific (and at times not so scientific!) debate of just how deep surface karst penetrates into the bedded halite of the Salado Formation in the vicinity of the WIPP site continues today.


Further north, near the subcropping western edge of the Northwest Permian Shelf, is Bottomless Lakes State Park, located some 20 kilometres southeast of Roswell. Encircling the lakes are the gypsum, halite and dolomitic redbeds of the Artesia Group and San Andres Formation. Away from the lakes are numerous other sinkholes and collapse dolines, most of which are circular, steep walled or vertical holes, 50-100 metres across and 30-60 metres deep, with the greatest density of features aligned along the eastern side of the Pecos River floodplain (Figure 6). Water in the various sinks that make up the Bottomless Lakes is crystal clear and brackish to saline (6,000-23,000 ppm), attesting to its passage through subsurface layers of gypsum and salt. Although the lakes are around 30 metres deep, dark-green moss and algae coat the bottoms giving the impression of great depth and hence the name of the park (Lea Lake; Figure 10d.). To the west, many of the playas in depressions near Amarillo in the High Plains of Texas have a similar genesis as solution depressions atop dissolving Permian salt, but most do not intersect the regional watertable and so do not hold permanent surface water (Paine, 1994). Further south, near Carlsbad Caverns, there are other subsidence chimneys that form lakes where they intersect the watertable and so are also locally known as “Bottomless” (Figure 9a).


Western Canadian karst

The distribution and timing of chimneys and subsidence troughs created by the subsurface dissolution of the Prairie evaporite are well known and tied to the distribution of oil sands and the quality of potash ores (part 5). Dissolution drape features are more pronounced nearer the retreating edges of the thick multilayered subsurface Devonian salt succession while breccia chimneys occur above the buried salt successions (Figure 9b). For example, the Rosetown Low and the Regina Hummingbird Trough accumulated more than 100 metres of depression trough and drape sediment during interstratal dissolution of the underlying Prairie salt (Devonian) in southern Saskatchewan, especially in Cretaceous time (Figure 13a,b; DeMille et al., 1964; Simpson, 1988). This evaporite-related Cretaceous subsidence is also the principal control on the distribution of the Athabasca oil sands in subsidence basins along the eastern side of the evaporite extent (Tozer et al., 2014). Uplift of the ancestral Rocky Mountains likely created the potentiometric head that drove much of the subsalt aquifer flow. As in the Delaware Basin, the positions of sub-salt reefs and pinnacles focused many of the upwells of deeply circulated meteoric water that ultimately created solution breccia layers and breccia chimneys (Figure 8b).

In various circum-salt subsidence troughs, now filled or partially filled with late Mesozoic and Tertiary sediment thicks, the concurrence of a supra-unconformity thick, adjacent to the sub-unconformity feather-edge of a bedded salt sequence, is at a scale that is easily recognised in seismic and constitutes one of the classic signatures of a salt collapse-induced hydrocarbon trap (Figure 8a). For now, we will focus on the influence of dissolving evaporites on the modern Canadian landscape in regions of active karst, but we will return to the topic of economic hydrocarbon associations with paleokarst in this basin in part 5.

Evaporite karst domes, and laterally extensive solution breccia units, tied to a waxing and waning cover of glacial ice and permafrost, were first documented in Canada in northern Alberta and the Northwest Territories (e.g. De Mille et al., 1964). Karst domes remain as surface features of positive relief once the surrounding evaporite mass has completely dissolved and are outlined by megabreccia with caverns. Unlike breccia chimneys, the cores of evaporitic karst domes can expose blocks from below, as well as above the original bedded and folded evaporite level. The evaporitic karst domes in western Canada are related to the stratigraphic level of former bedded salts; elsewhere in the world others are the remains of now dissolved salt thrusts, diapirs and allochthons (see diapiric breccias and rauhwacke discussion in Warren, 2016). Hence, domes and residual units are dramatic landscape features in the gypsiferous terrain of northern Alberta, Canada (Wigley et al., 1973; Tsui and Cruden, 1984). Similar features, tied to the dissolution of bedded evaporites, typify the Arkhangelsk gypsum-residue karst region inland of the Barents Sea coast of Russia (Korotkov, 1974).

Canadian karst domes range from 10 to 1000 m or more in length or diameter, and can rise to 25 m above the surrounding land surface. Many domes are highly fractured, with individual overburden blocks displaced by heaving and sliding, with the residual gypsum showing well-developed flow foliation.

At the extreme end of disturbance and dissolution range, the domes breccias are megabreccias; positive relief features made up of collapsed or even upthrust jumbles of large blocks, some the size of houses. The largest reported Canadian megabreccia example is in a steep-limbed anticline that extends along the shore of Slave Lake for a distance of 30 km. It is up to 175m in height with a brecciated crestal zone that is marked by a ‘chaotic structure and trench-like lineaments’ (Aitken and Cook 1969).

These brecciated landforms develop upon a distinctive geological association in the central Mackenzie Valley region, NWT, at a stratigraphic level equivalent to anhydrites and halites in the deeper subsurface (Hamilton and Ford, 2002 ). In widespread outcrops the breccia unit is formally named the Bear Rock Formation, when covered by consolidated strata it is termed the Fort Norman Fm (Meijer Drees, 1993; Law, 1971). It is centred in Late Silurian-Early Devonian strata and defines an outcrop and subcrop belt more than 50,000 km2 in extent. In core, the Fort Norman Formation is 250-350 m or more in thickness. It consists of a thin upper limestone (Landry Member), a central Brecciated Member and, in some cores, undisturbed lower sequences of inter­bedded dolostones and anhydrite remnants. It is conformably overlain by 90-150 m of limestones and calcareous shales (Hume Formation-Eifelian). This evaporite dissolution breccia is possibly derived from the dissolution of not one, but several subcropping Devonian salt layers. That is, although not much discussed in the literature, the mother level may not be tied to a single stratigraphic layer.

Typical till-covered collapse and subsidence karst of the Bear Lake Formation can develop through the Hume and higher formations as a consequence of interstratal dissolution of the salt layers in the Fort Norman and equivalents, where meteoric groundwater circulation and sulphate dissolution have been recorded at core depths as great as 900 m (Figure, 14, 15).


In the NWT the Bear Rock Formation is considered to host to much of this widespread solution breccia, which up to 250 m thick. If present, the Landry Member is a brecciated limestone no more than 20 m thick. The main breccia level forms a visually set of outcrops made up of chaotic, vuggy mixtures of limestone, dolomite, and dedolomite clasts, variably cemented by later calcites, with small residual clasts and secondary encrustations of gypsum. Pack breccias in the Bear Creek and its equivalents displaying rubble fabric (predominant), crackle or mosaic fabric, are common and cliff-forming (Stanton, 1966; James & Choquette, 1988). Float breccias are rarer and tend to be recessive. Meijer Drees (1985) classifies the Bear Rock Formation as a late diagenetic solution breccia created by meteoric waters. Hamilton (1995) shows that calcite, dolomite and sulphate dissolution, plus dedolomitisation with calcite precipitation, are continuing today. They are re-working older breccia fabrics, creating new ones, as well as forming a suite of surficial karstic depressions and subsidence troughs ranging from metres to several km in scale.

The Canadian example constitutes an important set of observations that also relate to many other regions with widespread, basin-scale evaporite dissolution breccias, namely; evaporite solution breccias are multistage and can encompass significant time intervals. Similar-appearance evaporite solution breccias can form at different times in a basin's burial history, from different superimposed undersaturated hydrologies. Individual breccias in any single sample are typically responses to multistage, multi-time diagenetic-fluid overprints. This is also why potash ores in the Prairie evaporite preserve evidence of multiple times of potash mobilisation and mineralogy (part 5; Warren, 2016).

Detailed evaporite karst landform studies have focused on Bear Rock Mountain (type area) and the Mackay Range, which are outlying highlands on the east and west banks respectively of the Mackenzie River, and terrain between Carcajou Canyon and Dodo Canyon in the Mackenzie Mountains (Hamilton, 1995). These sites were covered by the Laurentide Ice Sheet (Wisconsinan) but were close to its western, sluggish margin. They are at the boundary between widespread and continuous permafrost in the ground today and can display some year-round groundwater circulation via taliks. Thaw/freeze and solifluction processes compete with ongoing evaporite dissolution to mould the topography.


Regionally, the principal karst landforms hosted in and above the Bear Rock Formation are varieties of sinkholes, blind valleys, solution-subsidence troughs and fault-bounded depressions (e.g. Figures 14, 15). Sinkholes range from single colla­pse features a few metres in diameter to merged or compound dolines up to 1.5 km2 and 100 m deep. Smaller individual sinkholes and collapse dolines may retain seasonal meltwater ponds, and there are permanent lakes draining to marginal ponors in some of the larger subsidence troughs. Blind valleys have developed where modern surface streams flow for several km into the evaporite karst zone from adjoining rocks. There are many relict, wholly dry valleys that may have been created by glacial meltwaters. Subsidence troughs have developed at the surface along contacts between the Bear Rock breccias and underlying, massive dolostones that are typically gently dipping. In the centre of the Mackay Range and at Bear Rock itself are solution depressions formed where the breccias make-up hanging-wall strata on steeply inclined fault planes on the collapse or pipe edge. The Mackay example is 3.2 km long, 1.0 km wide and-160 m deep (Ford, 1998).

Where patches of the Landry Member survive above the main breccia level(s), they are often broken into large, separate slabs that tilt into adjoining depressions in sharply differing directions, creating a very distinctive topography of dissolution draping and block rotation (see also Dahl Hit in Warren, 2016). Some slabs are rotated through 80-90°. Ridges (inter-sinkhole divides) that are wholly within the main breccias often display stronger cementation, represented by pinnacles as much as 30 m high The many sinking streams pass through the permafrost via taliks and emerge as sub-permafrost springs at stratigraphic contacts or topographic low points.

According to Hamilton (1995), the variety and intensity of karst landform assemblages on the Bear Rock Formation are like no other in Canada, and he notes he had not seen or read of very similar intensive karst topographies elsewhere. He attributes their distinctiveness to repeated evaporite dissolution and brecciation, with dedolomitisation and local case hardening, throughout the Tertiary and Quaternary, with these processes occurring multiple times in mountainous terrains subject to episodes of glaciation, permafrost formation and decay, and to vigorous periglacial action.

The spectacular karst domes that typify dolines and glacially associated surfaces atop an evaporite subcrop and are particularly obvious in regions of anticlinal salt-cored structures within regions of widespread permafrost. The accumulation of ground ice in initial fractures in the evaporite layer probably contributes to the heaving, folding and other displacement of breccia fragments in the dome. Tsui and Cruden (1984) attributed the examples that they studied in the salt plains of Wood Buffalo National Park Canada (Lat. 59-60°N) to hydration processes operating on subcropping bedded gypsum during the postglacial period. Ford and Williams (2007) argued such features indicate local injection of gypsum residuals during times of rapidly changing glacial ice loading. Whether they are created by glacial unloading/reworking or are a type of gravitationally-displaced dissolution breccia in a permafrost region is debatable; that in Canada they are a widespread type of evaporite karst residue is not. They characterise those parts of the permafrost-influenced region defined by dissolution of shallow subcropping folded Devonian salts in Northern Alberta, where the evaporites are typically exposed in anticlinal crests are today still retain fractionated gypsum residuals.

And so, in addition to widespread breccias hoisting karst domes, there are numerous natural collapse dolines atop dissolving shallow salt beds. A classic example is the water filled doline some km NE of Norman Wells (Figure 15).


Holbrook Anticline, Arizona

Subsidence driven by natural salt solution at depth generates regional-scale drape or monoclinal folding in strata atop the retreating salt edge. This is the corollary of the formation of a subsidence trough (Figure 16a). The 70-km long dissolution front in the Permian Supai salt of Arizona is defined by more than 500 sinkholes, fissures, chimneys and subsiding depressions some 40-50m across and 20-30m deep (Neal, 1995; Johnson, 2005). Away from the main dissolution zone the northeasterly-dipping Schnebly Hill Formation (aka Supai Salt) is composed of up to 150 metres of bedded halite, with local areas of sylvite along its northern extent (Figure 16a). Atop the solution front, there are a number of topographic depressions with playa lakes that in total cover some 300 km2. Salt-dissolution induced features to include areas known as; The Sinks, Dry Lake Valley, and the McCauley Sinks (Neal et al. 1998; Martinez et al. 1998). The Sinks region includes more than 20 steep-sided caprock sinkholes, with some that are more than 100 metres across and up to 30 metres deep (Figure 16b). The McCauley Sinks are the likely surface expression of compound breccia pipes and chimneys (Neal and Johnson, 2002).

Regional expanses of Supai Salt removal produce a regional gravity depression, largely coincident with a surface topographic depression, where there is as much as 100 metres of collapse and topographic displacement. The solution front is essentially coincident with the updip end of the Holbrook Anticline, a flexure defined by dip reversal in an otherwise northeasterly dipping succession (Figure 16a).

Rather than orogenically driven, the Holbrook Anticline is a subsidence-induced monoclinal flexure created by the northeasterly migrating dissolution front (Figure 16a). It may be the largest single solution-collapse fold structure in the world (Neal, 1995). The reverse dip of the flexure directly overlies the salt-dissolution front and marks the location of two major collapse depressions known as The Sinks and Dry Lake Valley, both occur where salt is within 300 m of the surface (Peirce, 1981). Although it has periodically held surface water, reports of several hundred acre-feet of flood water in the Dry Lake Valley playa draining overnight supports the notion of active fissure and cavern formation related to salt removal at depth. Major surface drainage events took place in 1963, 1979, 1984 and 1995, with more than 50 new sinkholes forming in the valley during that period. The continuing rapid appearance of new sinkholes testifies to the ongoing nature of dissolution in the underlying evaporites. According to Neal (1995), dissolution front features began forming in the landscape in the Pliocene and continue to form today.

The caprock collapse sinkholes centre in the Coconino Sandstone bed that overlie the salt, which is located 200-300 metres below the at-surface features. These caprock sinkholes define regions of focused breccia pipe development and discreet upward cavity migration (chimney stoping). The underlying salt is bedded and there are no indications of halokinesis anywhere in the basin. Sinkhole regions lie just ahead of a dissolution front that is migrating downdip, driven by the widespread dissolution of halite.

Other karst features attributed to evaporite dissolution in the Holbrook basin are; pull-apart fissures, graben sinks (downward-dropped blocks), breccia pipes and plugs, and numerous small depressions with and without sinkholes (Neal et al., 1998, 2001). Interestingly, many of the “karst” features occur in sandstones, not limestones; such caprock collapse sinkholes can form wherever pervasive dissolution has removed the underlying salt, independent of caprock lithology. The presence of the more than 500 karst features in the Holbrook basin, some of which formed in days, evokes practical karst hazard and infrastructure concerns, even in such a sparsely populated region (Martinez et al., 1998).

Implications

Evaporite-related karst landforms are in many ways similar carbonate karst features. But, the much higher solubilities of halite and other evaporite salts compared to limestone and dolomite means there are additional features unique to regions of salt dissolution.

In the subsurface, a bedded evaporite can be composed of thick impervious halite with intrasalt beds composed of anhydrite dolomite and calcite. The edges of this halite bed can dissolve even in the deep mesogenetic (burial) realm, wherever the edge of the salt is in contact with undersaturated waters. Insoluble residues bands start to form and can take the form of a basal anhydrite or a fractionated caprock. If the rise of undersaturated water is focused, a breccia cavern can form, and once it breaches the salt, the cavity will contain collapse blocks of the less soluble intrasalt beds. The cavity can then stope to the surface, forming a breccia pipe. A transtratal breccia pipe can rise through kilometres of overburden before attaining the surface. There is no equivalent process-response in the carbonate realm.Subsidence basins, troughs and megabreccia plains are also features that owe their origins to the rapid dissolution rates inherent to salt bed-fresher water contacts

In a similar fashion, the rapid dissolution of halite in the shallower parts of a salt basin means the evaporite karst features will transition from mesogenetic or bathyphreatic cavern formation to meteoric phreatic to vadose effects. This is the emphasis of the second article in this series, which will discuss processes forming vadose and phreatic caves in evaporites.

References

Aitken, J. D., and D. G. Cook, 1969, Geology, Lake Belot, District of Mackenzie: Geological Survey of Canada Map 6.

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

Anderson, E. J., 1981, Deep-seated salt dissolution in the Delaware Basin, Texas and New Mexico: in Wells, S.G., Lamber, W., and Callender, J.F., eds., Environmental geology and hydrology in New Mexico: Special Publication 10, New Mexico Geological Society, p. 133-145.

Bachman, G., and R. B. Johnson, 1973, Stability of Salt in the Permian Salt Basin of Kansas, Oklahoma, Texas and New Mexico: U. S. Geol. Surv. Open-File Rept.. v. 4339-4.

Bachman, G. O., 1981, Geology of Nash Draw, Eddy County, New Mexico.: U. S. Geol. Surv. Open-File Rept. 81-3.

Bachman, G. O., 1984, Regional geology of Ochoan evaporites, northern part of Delaware Basin: New Mexico Bureau of Mines and Mineral Resources, Circular, v. 184, p. 22.

Belloni, S., B. Martins, and G. Orombelli, 1972, Karst of Italy, in M. Herak, and V. T. Springfield, eds., Karst: Important karst regions of the Northern Hemisphere: Amsterdam, Elsevier, p. 85-128.

Broughton, P. L., and D. Cotterill, 2017, Breccia pipe and sinkhole linked fluidized beds and debris flows in the Athabasca Oil Sands: dynamics of evaporite karst collapse-induced fault block collisions: Bulletin of Canadian Petroleum Geology, v. 65, p. 200.

Burri, E., 1986, Various aspects of karstic phenomena in the urbanised area of Gissi and neighbouring areas (southern Abruzzo, Italy): Le Grotte d'Italia, v. 4, p. 143-161.

Christiansen, E. A., and E. K. Sauer, 2001, Stratigraphy and structure of a Late Wisconsinan salt collapse in the Saskatoon Low, south of Saskatoon, Saskatchewan, Canada: an update: Canadian Journal of Earth Sciences, v. 38, p. 1601-1613.

Davies, P. B., 1984a, DeepSeated Dissolution and Subsidence in Bedded Salt Deposits: doctoral thesis, Stanford University.

Davies, P. B., 1984b, Structural analysis of a deep seated salt dissolution collapse chimney; implications for nuclear waste disposal: Neues Jahrbuch fuer Mineralogie, Abhandlungen, v. 149, p. 163-175.

De Mille, G., J. R. Shouldice, and H. W. Nelson, 1964, Collapse structures related to evaporites of the Prairie Formation, Saskatchewan: Geological Society America Bulletin, v. 75, p. 307-316.

Ferrarase, F., T. Macaluso, G. Madonia, P. A., and U. Sauro, 2002, Solution and recrystallisation processes and associated landforms in gypsum outcrops of Sicily: Geomorphology, v. 49, p. 25-453.

Ford, D., and P. D. Williams, 2007, Karst hydrology and Geomorphology: New York, John Wiley and Sons, Ltd, 562 p.

Ford, D. C., 1998, Principal features of evaporite karst in Canada: Suppl. Geogr. Fis. Dinam. Quat. III, Proceedings of fourth International Conference on Geomorphology, Italy 1997:Karst Geomorphology, v. 12, p. 11-19.

Goodbar, J., and A. Goodbar, 2014, A Method to Investigate Karst Groundwater Flow in Nash Draw Eddy County, New Mexico To Delineate Potential Impacts of Potash Industry Discharge and Runoff U.S. Geological Survey Karst Interest Group; Conference Paper April 2014 held at Carlsbad New Mexico.

Gorbunova, K. A., 1979, Gidratatsiya angidrita i soputstvuyushchiye yey yavleniya: Koval'chuk, A. I. Karst i gidrogeologiya Predural'ya. Akad. Nauk Sssr, Ural. Nauchn. Tsentr, Inst. Geol. Geokhim., Tr.

Gustavson, T. C., W. W. Simpkins, A. Alhades, and A. D. Hoadley, 1982, Evaporite dissolution and development of karst features on the Rolling Plains of the Texas Panhandle: Earth Surface Processes and Landforms, v. 7, p. 545-563.

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., J. Guerrero, and P. Lucha, 2008, A genetic classification of sinkholes illustrated from evaporite paleokarst exposures in Spain: Environmental Geology, v. 53, p. 993-1006.

Hamilton, J., and D. Ford, 2002, Karst geomorphology and hydrogeology of the Bear Rock Formation — a remarkable dolostone and gypsum megabreccia in the continuous permafrost zone of northwest Territories, Canada: Carbonates and Evaporites, v. 17, p. 114-115.

Hamilton, J. P., 1995, Karst geomorphology and hydrogeology of the northeastern Mackenzie Mountains, District of Mackenzie, N.W.T., Canada: Doctoral thesis, McMaster University, 532 p.

Hill, C. A., 1990, Sulphuric acid speleogenesis of Carlsbad Caverns and its relationship to hydrocarbons, Delaware Basin, New Mexico and Texas: American Association of Petroleum Geologists Bulletin, v. 74, p. 1685-1694.

Holt, R. M., and D. W. Powers, 2010, Evaluation of halite dissolution at a radioactive waste disposal site, Andrews County, Texas: Geological Society of America Bulletin, v. 122, p. 1989-2004.

James, N. P., and P. W. Choquette, 1988, Paleokarst: New York, Springer, 416 p.

Johnson, K. S., 2005, Subsidence hazards due to evaporite dissolution in the United States: Environmental Geology, v. 48, p. 395-409.

Kirkland, D. W., and R. Evans, 1976, Origin of limestone buttes, Gypsum Plain, Culberson County, Texas: American Association of Petroleum Geologists, Bulletin, v. 60, p. 2005-2018.

Kirkland, D. W., and R. Evans, 1980, Origin of castiles on the Gypsum Plain of Texas and New Mexico: Guidebook New Mexico Geological Society, v. 31, p. 173-178.

Korotkov, A. N., 1974, Caves of the Pinego-Severodvinskaja karst (in Russain): Geog. Soc. USSR, Leningrad.

Lambert, S. J., 1983, Dissolution of Evaporites in and Around the Delaware Basin, Southeastern New Mexico and West Texas: Sandia Nat’l. Labs. Report SAND82-0461, Albuquerque, NM, 96 pp.

Land, L., 2013, Evaporite Karst in the Permian Basin Region of West Texas and southeastern New Mexico: The Human Impact, in L. Land, D. H. Doctor, and J. B. Stephenson, eds., Sinkholes and the engineering and environmental impacts of karst; Symposium 2; Proceedings of the 13th Multidisciplinary Conference, May 6 through 10, 2013, Carlsbad, New Mexico, p. 113-121.

Law, J., 1971, Regional Devonian Geology and Oil and Gas Possibilities, Upper Mackenzie River Area: Bulletin of Canadian Petroleum Geology, v. 19, p. 437-484.

Macaluso, T., and U. Sauro, 1996, Weathering and Karren on exposed gypsum surfaces: International Journal of Speleology, v. 25, p. 115-126.

Martinez, J. D., K. S. Johnson, and J. T. Neal, 1998, Sinkholes in Evaporite Rocks: American Scientist, v. 86, p. 38.

Meijer Drees, N. C., 1985, Evaporitic deposits of western Canada: Geological Survey of Canada, v. Paper 85-20, p. 118 pp.

Meijer Drees, N. C., 1993, The Devonian succession in the subsurface of the Great Slave and Great Bear Plains, Northwest Territories: Geological Survey of Canada Bulletin, No. 393, 222 p.

Neal, J. T., 1995, Supai salt karst features: Holbrook Basin, Arizona, in B. F. Beck, ed., Karst geohazards: engineering and environmental problems in karst terrane. Proc. 5th conference, Gatlinburg 1995, Balkema, p. 53-59.

Neal, J. T., R. Colpitts, and K. S. Johnson, 1998, Evaporite Karst in the Holbrook Basin, Arizona, in J. Borchers, ed., Joseph F. Poland Symposium on Land Subsidence: Sudbury, MA, Assoc. Eng. Geologists Spec. Pub. 8, p. 373-384.

Neal, J. T., R. Colpitts, and K. S. Johnson, 2001, Evaporite Karst in the Holbrook Basin, Arizona: SMRI (Solution MIning Research Institute) Report presented at the Fall 2001 Meeting 7 - 10 October 2001 Albuquerque, New Mexico, USA.

Olive, W. W., 1957, Solution-subsidence troughs, Castile Formation of Gypsum Plain, Texas and New Mexico: Geological Society of America, v. 68, p. 351-358.

Paine, J. G., 1994, Subsidence beneath a playa basin on the Southern High Plains, U. S.A.; evidence from shallow seismic data: Geological Society of America Bulletin, v. 106, p. 233-242.

Powers, D. W., R. L. Beauheim, R. M. Holt, and D. L. Hughes, 2006, Evaporite karst features and processes at Nash Draw, Eddy County, New Mexico, in L. Land, V. Lueth, B. Raatz, P. Boston, and D. Love, eds., Caves and Karst of Southeastern New Mexico: New Mexico Geological Society, Guidebook 57, p. 253-266.

Quinlan, J. F., 1978, Types of karst, with emphasis on cover beds in their classification and development: Doctoral thesis, University of Texas at Austin, 323 p.

Quinlan, J. F., R. O. Ewers, J. A. Raty, and K. S. Johnson, 1986, Gypsum karst and salt karst of the United States of America: Le Grotte d'Italia, v. 4, p. 73-92.

Simpson, F., 1988, Solution-generated collapse (SGC) structures associated with bedded evaporites; significance to base-metal and hydrocarbon localization: Geoscience Canada, v. 15, p. 89-93.

Stafford, K. W., R. Nance, L. Rosales-Lagarde, and P. J. Boston, 2008, Epigene and Hypogene Gypsum karst manifestations of the Castile formation: Eddy County, new Mexico and Culbesron County, Texas, USA: International Journal of Speleology, v. 37, p. 83-98.

Stafford, K. W., L. Rosales-Lagarde, and P. J. Boston, 2008, Castile evaporite karst potential map of the Gypsum Plain, Eddy County, New Mexico and Culberson County, Texas: A GIS methodological comparison., v. 70, no. 1, p. 35–46.: Journal of Cave and Karst Studies, v. 70, p. 35-46.

Stanton, R. J., Jr, 1966, The solution brecciation process: Geological Society of America Bulletin, v. 77, p. 843-847.

Stenson, R. E., and D. C. Ford, 1993, Rillenkarren on gypsum in Nova Scotia: Geographie Physique et Quaternaire, v. 47, p. 239-243.

Tozer, R. S. J., A. P. Choi, J. T. Pietras, and D. J. Tanasichuk, 2014, Athabasca oil sands: mega-trap restoration and charge timing: Bulletin American Association Petroleum Geologists, v. 98, p. 429-447.

Tsui, P. C., and D. M. Cruden, 1984, Deformation associated with gypsum karst in the Salt River Escarpment, northeastern Alberta: Canadian Journal of Earth Sciences, v. 21, p. 949-959.

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.

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

Wigley, T. M. L., J. J. Drake, J. F. Quinlan, and D. C. Ford, 1973, Geomorphology and geochemistry of a gypsum karst near Canal Flats, British Columbia: Canadian Journal of Earth Sciences, v. 10, p. 113-129.


 

 

Salt as a Fluid Seal: Article 2 of 4: Internal fluid source

John Warren - Wednesday, January 20, 2016

 

Black Salt: as an indicator of overpressure

The previous article in this series on salt leakage focused on black and dark salt created by ingress or interaction of undersaturated waters with relatively shallow halokinetic salt masses, with entry zones often tied to intervals of salt shear. The resulting black or dark salt textures are one style of “anomalous” salt. This article looks at fluid entry into salt in subsurface intervals of high pore pressure, exemplified by the “black salt” in the Ara salt seals of Oman. Such intervals are often tied to burial-pressure and temperature-related changes to the dihedral angle of salt (halite).


Dihedral angle changes and the permeability of salt

Permeability in intercrystalline pore networks in re-equilibrating and crystallising subsurface salt is tied to the dihedral angle  at solid-solid-liquid triple junctions (Figure 1; Lewis and Holness, 1996, Holness and Lewis, 1997). When the halite dihedral angle is higher than 60° under static laboratory conditions, this contact angle equates to the maintenance of closure of polyhedral grain boundaries by halite precipitation, and so at these lower temperatures both bedded and halokinetic recrystallized salt is impermeable (Schenk and Urai, 2004; Holness and Lewis, 1996). In this temperature range, the small amount of brine present in the salt is distributed in micrometer-sized isolated fluid inclusions at termini of salt crystal polygon apices. In contrast, when the solid-solid-liquid interfaces of increasingly heated and pressurised polyhedral halite attain dihedral angles that are less than 60° then the fluid-inclusion filled intercrystal cavities link up and the salt mass becomes permeable.


At burial temperatures >100°-150°C and pressures of 70 MPa or more, the dihedral angle has decreased to values <60°, driving a redistribution of the fluid into a thermodynamically stable network of connected, fluid-filled channels or fused fluid strings at grain boundary triple junctions. This transition may be related to the observation by Peach and Spiers (1996) that, during natural deformation of rocksalt at great depths, salt undergoes natural hydraulic fracturing or dilatancy. The dihedral angle is, therefore, a thermodynamic property that changes with pressure P and temperature T. Holness and Lewis’s experiments demonstrated that buried salt masses, subject to high pressures and elevated temperatures, can acquire intercrystalline or polyhedral permeability comparable to associated with intergranular permeability in sand.

This typically occurs at higher temperatures and pressures where intercrystal water positions link within flowing or static, but texturally re-equilibrated, salt and so creates continuous fluid strings along evolving intercrystalline junctions in the burial-recrystallised salt. The newly attained intercrystal configuration allows penetration and throughflow of hot, dense brines or hydrocarbons into and through the altered mass of salt polyhedrons. In Oman has created characteristic haloes of black salt about pressurised salt-encased carbonate slivers (next section).

At the same time as a recrytallising salt mass passes into the earlier stages of the greenschist facies, the salt is dissolving and altering to sodic scapolite (Warren, 2016; Chapter 13). Thus, through the later stages of diagenesis and into early to medium grades of metamorphism, the salt and its daughter products may act as sources and conduits for flow of chloride-rich metalliferous brines and salt slurries. This occurs as bedded and halokinetic salt evolves from a dense impermeable salt mass into permeable salt with higher dihedral angles and so explains salt’s significant role in the creation of many of massive base metal and IOCG deposits (Warren, 2016; Chapters 15, 16).


Black salt and overpressure in Oman

The transition in dihedral angle with increasing pressure and temperature explains the occurrence of black (bitumen-charged) haloes in salt encasing some carbonate-sliver reservoirs in the Ara Salt of Oman (Figure 2; Kukla et al., 2011a, b). Once this recrystallization occurs, the previous lower P&T mosaic halite loses its ability to act as an aquitard or aquiclude (seal) and can instead serve as a permeable conduit for escaping highly-pressurised and hydrocarbon-rich formation waters. According to Lewis and Holness, the depth at which the recrystallization occurs may begin as shallow as a few kilometres (Figure 1). But, their pressure bomb laboratory-based static-salt experiments did not completely encompass the ability of natural salt to pressure creep and self-seal by longer-term diffusion-controlled pressure solution (Warren 2016, Chapter 6). Even if the changing dihedral angles alter and open up permeability at such shallow depths, there is no guarantee that subsequent flowage associated with pressure solution will not re-anneal these new pores. The ability of salt to continue to act as a highly efficient hydrocarbon seal to depths of 6-10 km means, in my opinion, that bedded salt does may become a relative aquifer until attaining depths of 6-10 km or more. This occurs certainly at temperatures and pressures where the sequence is entering the greenschist realm. In extremely overpressured situations the transition of dihedral angles is much shallower, as in the 40-50m thick black salt rims that typify the salt-encased hydrocarbon-charged carbonate stringers in the Ara Salt of Oman (Kukla et al., 2011b). Once it does transform into polyhedral halite, a former aquiclude becomes an aquifer flushed by chloride-rich brines, likely carrying other volatiles.

A release of entrained inclusion (±intercrystalline) water at temperatures > 300-400°C (early greenschist) influences the textures of deeply buried halite. Most of the inclusions in chevron halite and other inclusion-rich cloudy primary salts are due to entrained brine inclusions and not mineral matter. Figure 3 plots the weight loss of various types of halite during heating. It clearly shows cloudy (inclusion-rich) halite releases up to 5 times more brine (0.2-0.5 wt%) than clear coarsely crystalline halite. An analysis of all fluids released during heating shows carbon dioxide and hydrogen contents are much lower than the water volumes: CO2/H2O < 0.01 and H2/H2O < 0.005. Organic compounds, with CH4, are always present (<0.05% H2O), and are twice as abundant in cloudy halite. There are also traces of nitrogen and, in some samples, hydrogen sulphide and sulphur dioxide (Zimmermann and Moretto, 1996).


The influence of overpressure driving changes in the dihedral angle of pressurised salt is most clearly seen in black-salt encased Late Neoproterozoic to early Cambrian intra-salt Ara (stringer) reservoirs of the South Oman Salt Basin (Figures 2, 4, 5; Kukla et al., 2011b). These carbonate bodies are isolated in salt and frequently contain low-permeability dolomites and are characterised by high initial hydrocarbon production rates due to overpressure. But not all stringers are overpressured, and a temporal relationship is observed defined by increasingly overpressured reservoirs within stratigraphically younger units. There are two separate pressure trends in the stringers; one is hydrostatic to slightly-above hydrostatic, and the other is overpressured from 17 to 22 kPa.m−1, almost at lithostatic pressures (Figure 4).


The black staining of the halite is caused by intragranular microcracks and grain boundaries filled with solid bitumen formed by the alteration of oil (Figures 2, 5). The same samples show evidence for crystal plastic deformation and dynamic recrystallization. Subgrain-size piezometry indicates a maximum differential paleostress of less than 2 MPa. Under such low shear stress, laboratory-calibrated dilatancy criteria suggest that oil can only enter the rock salt at near-zero effective stresses, where fluid pressures are very close to lithostatic. In Schoenherr et al.’s (2007b) model, the oil pressure in the carbonate stringer reservoirs reservoir increases until it is equal to the fluid pressure in the low, but interconnected, porosity of the Ara Salt, plus the capillary entry pressure (Figure 5). When this condition is met, oil is expelled into the rock salt, which dilates and increases its permeability by many orders of magnitude. Sealing capacity is lost, and fluid flow will continue until the fluid pressure drops below te minimal principal stress, at which point rock salt will reseal to maintain the fluid pressure at lithostatic values. Inclusion studies in the halite indicate ambient temperatures at the time of entry were more than 90°C, implying hydrocarbons could move into interconnected polyhedral tubes in the halite. These conduits were created in response to changes in the polyhedral angle in the halite in response to elevated temperatures (Lewis and Holness, 1996).


Hydrocarbon-stained “black salt” can extend up to 100 metres from the pressurised supplying stringer into the Ara salt of Oman (Figure 2, 5). It indicates a burial-mesogenetic pressure regime and is not the same process set as seen in the telogenetic “black salt” regions of the onshore Gulf of Mexico. The latter is created by dissolution, meteoric water entry, and clastic contamination, as in the crests of nearsurface diapirs such as Weeks Island (Warren 2015). An Ara stringer enclosed by oil-stained salt but now below the lithostatic gradient likely indicates a later deflation event that caused either complete (C) or partial (E) loss of overpressures. Alternatively, stringers showing overpressure, but below the lithostatic gradient (E), might be explained by regional cooling or some other hitherto unexplained mechanism (Figure 4a; Kukla et al., 2011a, b).

Structural, petrophysical and seismic data analysis suggests that overpressure generation in the Ara is driven initially by rapid burial of the stringers in salt, with a subsequent significant contribution to the overpressure from thermal fluid effects and kerogen conversion of organic-rich laminites with the stringer bodies. If the overpressured stringers come in contact with a siliciclastic minibasin, they will deflate and return to hydrostatic pressures (A) in Figure 4. When the connection between the minibasin and the stringers is lost, they can regain overpressures because of further oil generation and burial (A’). If hydrocarbon production in undeflated stringers stops relatively early, the fluid pressures do not reach lithostatic pressures (B). If hydrocarbon generation continues, the fluid pressures exceed the lithostatic pressure (red star), leading to dilation and oil expulsion into the rock salt to what is locally known as “black salt” (D and E).

As well as these examples of overpressure associated with older evaporites, overpressure readily develops in salt-sculpted Tertiary basins. For example, overpressure occurs in salt shear (gumbo) transitions beneath some, but not all, shallow salt allochthons in Green and Mahogany Canyon regions in the Gulf of Mexico (Beckman, 1999: Shaker 2008). Where salt allochthons are climbing the stratigraphy, subsalt sealing and associated overpressure can occur beneath the salt mass at shallower levels than is observed in overpressured shale basins.

In terms of extension and compression regimes within a single allochthon tongue, Shaker (2008) noted that in extensional regions in halokinetic basins the magnitude and direction of the principal stresses are controlled by sediment load, salt thickness, and salt emplacement-displacement history. Therefore, the maximum principal stress is not necessarily represented by the sheer weight of the overburden, as is usually assumed in quiescent terranes. Salt buoyancy often acts upward and has the tendency to accelerate and decelerate the principal stress above and below the salt, respectively. A distinctive shift of the pore pressure envelopes and normal compaction trends takes place across the salt body in several wells drilled trough salt below minibasins in the Mississippi Canyon, Green Canyon, and Garden Banks areas of the Gulf of Mexico. A lower pore pressure gradient has been observed below the salt and a higher gradient above the salt barrier. On the salt-rooted minibasin scale, a high-gradient was also observed in areas where the salt was emplaced and a lower gradient where the salt withdrew (Shaker and Smith, 2002). On the other hand, in the compressional portion of a salt allochthon system, lateral stress generated by the salt movement piling up salt at the foot of the slope acts as the maximum principal stress, whereas the load of sediment represents the minimum stress.


Extreme overpressuring is commonplace in subsalt settings in the Gulf of Mexico at depths of 3000-4000 m and its variability creates drilling problems, as evidenced by the BP Horizon spill and explosion on April 20, 2010. Gas generated at greater depths in these regions can be trapped under the salt seal at pressures approaching lithostatic. It means drilling under the allochthonous salt on the Gulf Coast slope can intersect undercompacted sediments that are moderately to extremely overpressured and friable (Hunt et al., 1998). The influence of highly effective Jurassic salt seals on pressure gradients in the Neogene stratigraphy of the Gulf of Mexico is seen in the increased mud weights typically required for safe drilling, once an evaporite allochthon is breached by the drill (Table 1). Many wells intersecting salt allochthons in the deepwater realm of the Gulf of Mexico and the circum-Atlantic Salt basins are overpressured at some depth below the base of salt with mud weights controlling pressures ranging from 14 to 17.5 ppg.

Implications

This and the previous article (Warren, 2015) demonstrate that black salt is a form of anomalous salt that indicates salt has leaked, however, the locations and conditions where leakage has occurred are distinct. The black salt encountered in the salt mines of the US Gulf Coast are indicative of meteoric water entry in relatively shallow conditions in regions where the salt is in contact with the surrounding shales of muds that enclose the diapir salt core. In other words, fluid entry is from the outside of the salt mass and fluids move into the salt from its edges and likely enhance  the porosity in the intercrystalline salt. In contrast, the black salt occurrences in the Ara Salt of Oman are indicative of overpressure haloes, generated internally via hydrocarbon and fluid expulsion in carbonate slivers, which are are fully encased in salt. This creates naturally hydrofractured envelopes in the salt mass in zones where pressure and temperature induced changes in the dihedral angle has generated intercrystalline fluid strings within the recrystallised polyhedral halite. The two settings of black salt formation are distinct.

There is not a single mechanism that creates black salt in a halokinetic salt mass. We shall discuss the implications of this in the next article which will include a look at leakage models in halokinetic salt systems both in terms of their seal integrity and the implications for short and  long term storage of hydrocarbons and nuclear waste. 

References

Beckman, J., 1999. Study reveals overpressure sources in deep-lying formations. Oil and Gas Journal, September: 137.

Holness, M.B. and Lewis, S., 1997. The structure of the halite-brine interface inferred from pressure and temperature variations of equilibrium dihedral angles in the halite-H2O-CO2 system. Geochimica et Cosmochimica Acta, 61(4): 795-804.

Hunt, J.M., Whelan, J.K., Eglinton, L.B. and Cathles III, L.M., 1998. Relation of shale porosities, gas generation, and compaction to deep overpressures in the US Gulf Coast. In: B.E. Law, G.F. Ulmishek and V.I. Slavin (Editors), Abnormal pressures in hydrocarbon environments. American Association Petroleum Geologists Memoir 70, Tulsa, OK, pp. 87-104.

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

Kukla, P.A., Reuning, L., Becker, S., Urai, J.L. and Schoenherr, J., 2011b. Distribution and mechanisms of overpressure generation and deflation in the late Neoproterozoic to early Cambrian South Oman Salt Basin. Geofluids, 11(4): 349-361.

Lewis, S. and Holness, M., 1996. Equilibrium halite-H2O dihedral angles: High rock salt permeability in the shallow crust. Geology, 24(5): 431-434.

O'Brien, J. and Lerche, I., 1994. Understanding subsalt overpressure may reduce drilling risks. Oil and Gas Journal, 92(4): 28-29,32-34.

Peach, C. and Spiers, C.J., 1996. Influence of crystal plastic deformation on dilatancy and permeability development in synthetic salt rock. Tectonophysics, 256: 101-128.

Schenk, O. and Urai, J.L., 2004. Microstructural evolution and grain boundary structure during static recrystallization in synthetic polycrystals of sodium chloride containing saturated brine. Contributions to Mineralogy and Petrology, 146: 671-682.

Schoenherr, J., Littke, R., Urai, J.L., Kukla, P.A. and Rawahi, Z., 2007a. Polyphase thermal evolution in the Infra-Cambrian Ara Group (South Oman Salt Basin) as deduced by maturity of solid reservoir bitumen. Organic Geochemistry, 38(8): 1293-1318.

Schoenherr, J., Urai, J.L., Kukla, P.A., Littke, R., Schleder, Z., Larroque, J.-M., Newall, M.J., Al-Abry, N., Al-Siyabi, H.A. and Rawahi, Z., 2007b. Limits to the sealing capacity of rock salt: A case study of the infra-Cambrian Ara Salt from the South Oman salt basin. Bulletin American Association Petroleum Geologists, 91(11): 1541-1557.

Shaker, S., 2008. The double edged sword: The impact of the interaction between salt and sediment on subsalt exploration risk in deep water. Gulf Coast Association of Geological Societies Transactions, 58: 759-769.

Warren, J.K., 2015. Salt as a fluid seal: Article 1,  Salty Matters blog; First published on Dec 19, 2015; www.saltworkconsultants.com.

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

Zimmermann, J.L. and Moretto, R., 1996. Release of water and gases from halite crystals. European Journal of Mineralogy, 8(2): 413-422.


 

 

Salt as a Fluid Seal: Article 1 of 4: External fluid source

John Warren - Saturday, December 19, 2015

 

Introduction

In the next few articles, I plan to discuss salt’s ability to act as a fluid seal in a variety of halokinetic settings, as well as looking at the nature of the sealing salt. Of particular interest are formative mechanisms driving textural and permeability variations in zones that typify the salt side of the sealing boundaries in sub-vertical salt stems versus the lower contact transitions in sub-horizontal allochthons. In the first few articles, we shall focus on local-scale scenarios and salt seal textures, including situations where salt has leaked, and the intercrystalline or tetrahedral/polyhedral pores contain fluid or mineralogical evidence of leakage and crystal boundary dissolution. Within the salt mass, this is usually indicated by occurrences of “black” or “dark” salt in anomalous salt zones, some of which are intersected by workings in a number of salt mines. In contrast, in most oil exploration scenarios we only have wireline signatures to interpret the deeper and typically offshore seal horizons. Following on from that discussion, we shall look at more regional examples of cross-formational leakage. Finally, we will discuss implications of possible leakage in terms of understanding and predicting outcomes with respect to both waste storage and hydrocarbon sealing

“Black” or “dark” salt in anomalous salt zones

The geological term “black salt” covers a variety of salt textures and associated mechanisms of formation. The term “black” salt also has a non-technical culinary association (kala namak[1]) but, other than in the footnote, I will not discuss it further in this series of articles. The geological descriptor “black” or “dark” salt is widely used in the US salt mining industry as a pointer to possible zones of current or past natural fluid entry into the salt mass. Colouring fluids can be brine, oil or gas, often with solid impurities dominated by shale, anhydrite or calcite-dolomite. These intrasalt “black” or “dark” salt zones in a mine were also referred to as “shear” zones and considered pointers to what are often unstable regions, liable to fluid entry, gassy outbursts and roof or wall collapse. “Shear”, “black” and “dark” salt zones are better described under the broader term “anomalous salt” zones, many of which were or are  in fluid contact with the enclosing non-salt sediment mass (Kupfer, 1990).

In a somewhat related fashion, the term “black salt” is used by the oil industry in Oman and Europe to indicate subsurface zones of overpressured salt, where natural hydrofracturing has occurred, and hydrocarbons have penetrated up to 100 m into the sealing salt mass. Hence its dark color (naturally hydrofractured salt and its textures are the focus of the second article in this series on salt leakage). Fluid entry in this type of “black’ salt is ascribed to temperature-related changes in the dihedral angle of the halite crystals in “black” salt zones. In a similar fashion, the term “black” salt was used in a recent paper in Science by Ghanbarzadeh et al. (2015) and the dihedral angle changes are ascribed to temperature increases in halokinetic salt intervals in the offshore Gulf of Mexico. There the authors argue temperature increases have changed the intercrystalline dihedral angle in a salt mass, and so facilitated the entry of fluids from adjacent strata into the salt body.

So, the term “black” salt is used in the geological community without reference to geological criteria that can separate what I consider are at least two distinct styles of “black” or “dark” salt formation and leakage. One type of salt leakage occurs when the salt is relatively shallow and subject to dissolution driven by the entry of meteoric and other near-surface undersaturated waters into folded and refolded shear (anomalous) salt zones in and about salt stems and décollements. This typically occurs when the flowing salt crest is relatively shallow and tends to occur in regions where the leaking “black” salt zone is in contact with the nonsalt boundary edges of the halokinetic salt mass. This process set ultimately leads to an accumulation of insoluble residues (clays, anhydrite, gypsum, calcite, etc.) that define a unit called a caprock. The term “cap” is somewhat of a misnomer as “caprock” units also form on the sides and undersides of a salt mass, wherever the salt unit is in contact with undersaturated cross-flowing formational waters (Warren, 2016). The other type of “black salt, exemplified by the Ara salt in Oman is related to deeper salt burial, salt flow and an association with intrasalt pressurized fluids (a focus of next article in this series on salt leakage). Accordingly, if we are not to confuse styles of “black” salt genesis (meteoric or undersaturated fluid entry versus intrasalt overpressures) then a better non genetic term should be used to describe zones of "black" or "dark" salt. Although less euphonious, the better term is “anomalous” salt. This describes all zones within halite-dominant intervals with features that are not typical of the bulk of the main diapiric salt mass (Kupfer, 1990).

In this first article we look at various types of anomalous salt in salt mines, largely related to the entry of, or interaction with, undersaturated relatively shallow formation waters. The next article focuses on salt leakage and "black" salt related to overpressure. Then, as we shall see in the third article on salt leakage, there are significant implications of occurrences of anomalous salt with respect to practicalities of safe intrasalt storage and fluid contamination with respect to separating the two types of black salt. This is especially so when working in the subsurface without the luxury of core or mine wall exposures. Ignoring the origins of “black” or “dark” salt, and the associated implications for wireline interpretation, means any conclusions in terms of waste storage outcomes and/or hydrocarbon seal potential, by generalizing lab-based experimental results on leaking salt to all “black” salt occurrences in halokinetic settings, will be somewhat confused (e. g. Ghanbarzadeh et al., 2015).

Black salt (dark salt) in anomalous salt in response to undersaturated fluid entry

Intervals of “black” or dark salt are described in US Gulf Coast salt mines in publications by Balk (1953), Kupfer (1976, 1990) and Looff et al., (2010), the following observations are largely based on their work. Nearsurface (<1-2 km) portions of mined or cored diapirs with “black” salt zones in the Gulf Coast USA and Germany are segmented into a number of intradiapir zones showing differential movement between adjacent salt spines or flowing masses. The more homogenous intervals of consistently mineable salt ore are separated by anomalous zones, formerly called “shear” zones. This association of homogenous spines separated by narrower shear or anomalous zones was first mapped in mine walls in the Jefferson Island salt dome by Balk (1953). His work was one of a series of classic papers mapping the internal structural complexities and shears in various mined salt diapirs in the US Gulf Coast and the Zechstein of Germany. Subsequent work by Kupfer (1976) on the same US Gulf Coast Five-Island salt mines (Jefferson Island, Avery Island, Weeks Island, Cote Blanche and Belle Isle) further refined notions of internal shear and occurrences of “black” or “dark” salt in diapirs. Today, only the Cote Blanche and Avery Island salt mines are still in operation along the Five Island Salt Dome Trend (Figure 1)

A shear zone in a diapiric structure forms where adjacent parts of a salt structure are moving (rising or falling) at different rates. Such zones tend to dominate the perimeter of a salt structure across which salt mass is rising or falling with respect to the adjacent sediment and so grade outward from the salt spine into a boundary “shale sheath”. Older shear zones and shale sheaths also are commonplace in re-folded intervals within a salt stem. Mapping of these zones by Balk (1953), Muehlberger and Claubaugh (1968) and Kupfer (1976) across many salt mines showed salt in a diapir must flow at different rates at different times. Otherwise, the complex and highly variable internal refolded drape and napkin folds seen in diapirs in all the world’s salt mines could not form. Figure 2 illustrates some internal complexities the diapir scale typifying various diapiric salt structures across the world and the dominantlyvertical flow fabrics in diapir stems and subhorizonatal flow textures in overhangs and salt tongues. Figure 3 shows that same vertical dominance (biaxial elongation) of salt crystals from cores collected in diapir stems cored various salt mines, while Figure 4 shows the typical vertical banding fold style that typifies diapir stems.



Walden and Jacoby (1963) were the first to call attention to a Gulf Coast anomalous salt zone. They documented a fault zone in the Avery Island salt mine that separated the region of salt being mined, across an anomalous zone, from the domal core. To call attention to the zonal ductile, not brittle, nature of intradiapir salt flow, Kupfer, 1974 changed the description of such anomalous zones from "fault” zones to "shear” zones and concluded most intradiapir shear zones were not faulted zones (defined by brittle fracture offset). In a later paper, he suggested abandoning of the genetic and misleading term "shear zone" and proposed replacement with the broader nongenetic term "anomalous salt zone" (Kupfer, 1990).


The term “anomalous salt,” as defined by Kupfer (1990), is based on his then more than twenty years experience in various salt mines in the US Gulf Coast. An anomalous salt zone is defined broadly as a zone of anomalous features in salt of whatever origin. He noted that typical anomalous salt zone features are different to the majority of features in the adjacent salt and involve varying combinations of anomalous features that include:

Textures--Coarse-grained, piokiloblastic, friable

Inclusions--Sediments, hydrocarbons, brine, gases

Structures--Sheared salt, gas outbursts, brine leaks, undue roof and wall slabbing, jointing, voids, and slight porosity development

Compositions--Potash/magnesium, high anhydrite content, very black salt (made up of disseminated fluid and solid impurities.

The terms “anomalous salt” and “anomalous zones” as defined by Kupfer (1990), are based on observations across the various Five Island salt mines of South Louisiana (Figure 1). As later refined in Kupfer et al. (1998), anomalous salt is a rocksalt zone that deviates from what are considered typical domal salt. Typical Gulf Coast rocksalt according to Kupfer is reasonably pure halite (97%+/- 2%), with minor amounts of disseminated anhydrite (CaSO4) being the primary non-halite impurity. Grainsize is considered to be uniform with grain diameters of 3 – 10 mm (0.12 – 0.39 inches). With continued mapping of Five Island mines, Kupfer et al. (1998) and Looff et al. (2010) documented an even wider variety of anomalous salt zone characteristics and concluded that the creation of anomalous zones need not be related to faulting or shearing, but also can be related to fluid entry and salt dissolution (Figure 5). Anomalous salt can be defined by impurity content, structure, colour, or other features. Anomalous features may not have sharp contacts or uniform thickness, and most are not continuous over long distances. Individual anomalous features commonly disappear for tens of metres (hundreds of feet) only to appear over some horizontal distance. The internal salt fabric of a salt dome is always composed of both typical (volumetrically dominant) and anomalous salt. Kupfer (op. cit.) noted that other salt deposits, including horizontally bedded nonhalokinetic salt deposits in the Permian of West Texas and the Devonian of Western Canada, all have anomalous zones of various origins.


Further work in both the salt mines and salt cavern storage industry has increasingly invoked the concept of anomalous features, anomalous zones and boundary shear zones although there is still a significant confusion over the appropriate use of the terminology (Looff et al., 2010). Because of the flow experienced by diapiric salt, most anomalous salt features parallel the near vertical internal banding of the salt. Many anomalous salt features may create zones of differing creep, strength, or dissolution characteristics that can impact the solution mining and operation of a salt storage cavern and some are tied to zones of problematic fluid entry in a mine. An anomalous zone is any zone in a salt diapir that contains 3 or more dissimilar anomalous features (Kupfer, 1990). The term “anomalous” implies nothing regarding the genesis of the zone. While many anomalous zones may extend laterally over hundreds of metres in length, they are variable in nature, near vertical, and parallel to layering (Figure 5). Typical widths are poorly known but are commonly in the order of 30-50m; however individual structures or anomalous features within the anomalous zone may be as thin as millimetres.

Boundary Shear Zones (BSZ) and Edge Zones (EZ) are the two types of anomalous zones that have a genetic interpretation (Looff et al., 2010). Boundary shear zones are those zones that bound an active salt spine where the salt experiences increased shear stress due to differential salt movement. An edge zone is similar to a boundary shear zone except, instead of being internal within the dome, it is confined to the periphery of the salt stock. Anomalous salt is not restricted to shear zones, however within and about as diapir edge one can expect most anomalous salt to be associated with shear zones (Kupfer, 1990; Looff, 2000).

Anomalous zones within a salt spine are in many cases the remnants of relict BSZ’s from older spines incorporated into younger active salt spines and this especially obvious with those boundary zones associated with clastic impurities (Figure 6). Boundary shear zones and edge zones around the dome tend to be more problematic for storage caverns as they are likely to contain greater occurrences of anomalous salt, higher impurity content (including gas and brine) and structural features that may degrade salt quality and enable leakage. Thus salt caverns can be constructed within boundary salt zones, but if possible, they should be avoided as they can result in non-optimal operating conditions, long-term operational difficulties and in the most severe cases contribute to the loss of cavern integrity (Looff et al., 2010). In the case of edge zones, additional distance to the edge of the salt dome needs to be maintained not only to cover any uncertainty regarding the placement of the edge of salt with respect of mine workings but also to account for the potential for degraded salt quality and to provide a sufficient pillar of good quality salt between the mine or cavern wall and the edge of dome.


A top-of-salt boundary between aggradational and dissolutional components atop diapirs in the Five Islands salt landscape typically coincides with underlying anomalous zones of differential shear within the underlying diapir typically indicated by “black” or “dark” salt zones in the various diapirs (Kupfer, 1976; Lock, 2000). Where such interior anomalous “black” salt zones have intersected the edge of the salt mass, they tend to create intervals with a greater propensity for water entry or gas outbursts and unstable roof zone liable to slabbing and collapse (Figure 6). Such anomalous zones can leak water into a mine, and over the longer term create stability problems, as illustrated by problems in; the now abandoned Weeks Island oil storage facility, the Avery Island Salt Mine, and the likely association of a subvertical zone with anomalous salt, and the enhanced fluid entry that occurred during the Lake Peigneur collapse, which was tied to 1980 flooding of the former Jefferson Island Salt Mine. Today, only two of the former mines in the Five Island Salt Dome trend remain unflooded. For a more detailed discussion of these and other salt leakage scenarios tied to undersaturated fluid entry into salt mines and caverns, see case histories in Chapter 13, Warren 2016.

“Black” or “Dark” Salt zones and leakage into the former Weeks Island storage facility

In the walls of the now-flooded Weeks Island salt mine, Kupfer (1976) noted that wide black beds of the internal “shear” zone are unusual and not found over most of the rest of the mine where salt was extracted. In places, the anomalous zone beds contain black clay (Room J-21), orange sandstone (S-20), and other fragments of clastic material (Paine et al., 1965). These clastic remnants typically occur as balls or roundish blebs ranging in size up to tens of cm in diameter. Petroleum leaked out of seams in this black salt zone and seams in the surrounding salt; the escaping fluid ranged from yellow grease and heavy, blue oil to very light, straw-yellow distillate. Methane and carbon dioxide were also common. The width (surmised) and structural complexity of the anomalous zone suggest that internal salt movement continued after a clastic boundary sheath-zone was incorporated into the salt stock (Figure 7).


The cause of the drainage and abandonment of the Weeks Island oil storage facility was an active subsidence sinkhole some 10 metres across and 10 metres deep, first noted near the edge of the SPR facility in May 1992, and perhaps reaching the surface about a year earlier. The growing doline depression was located on the south-central portion of the island, directly over a subsurface trough, which was obvious in the top-of-salt contours based on former mine records before conversion to a hydrocarbon storage facility (Figure7; Neal and Myers, 1995). Earlier shallow exploratory drilling around the Department of Energy service and production shafts in 1986 had identified the presence of irregularities and brine-filled voids along the top of salt mass across this region. A second, much smaller sinkhole was noticed in early February 1995, but it did not constitute a serious threat as it lay outside the area of cavern storage.

The first sinkhole occurs in a position of sharp change in landform slope (transition from high island to gully fill) and lies atop the projected alignment of what is known as Shear Zone E (a dark salt zone) in the underlying salt (Autin and McCulloh, 1995). Neal (1994) pointed out that Kupfer’s 1976 map of that part of the Weeks Island salt mine, located beneath the first sinkhole, was defined by black salt (also shown as Figure 8 which is based on the more recent Kupfer et al. (1998) map). Miners always avoided such “black” salt or “dark” salt zones in the various subsurface workings and the lateral extent of workings in many of the Five Island mines extended only as far as intersections with significant “black” or “dark” salt regions (Figure 6 & 7).


The volume of the first Weeks Island sinkhole (estimated as 650 m3 when first noted), its occurrence over a trough in the top of salt, and its position directly above the oil-filled mine caverns, meant it was of urgent concern to the SPR authorities, especially in terms of the stability of the roof of the storage cavern. This feature did not form overnight; it lies atop a shear zone that formed during the diapiric rise of the salt and capped by a rockhead valley containing Pleistocene sediment fill. Salt extraction during mine operations probably created tension across the shear zone, thereby favouring fracture enlargement in the anomalous salt zone, as early perhaps as 1970 (Figure 6; Waltham et al., 2005). Eventually, an incursion of undersaturated groundwater traversed the fracture zone across some 107 m, from a level equivalent to the rockhead down to the mine where it emerged. Over time, ongoing dissolution enlarged a void at the top of the anomalous salt zone, creating the collapse environment for the sinkhole first noted at the land surface in 1991. Investigations were undertaken in 1994 and 1995 into the cause of active at-surface sinkholes verified that water from the aquifer above the Weeks Island salt dome was seeping into the underground oil storage chamber at the first sinkhole site (Figures 7& 8; Neal and Myers, 1995; Neal et al., 1995, 1997). Drainage and decommissioning of the Weeks Island facility followed.

Beginning in 1994, and continuing until the abandonment of the facility, saturated brine was injected directly into the throat of the first sinkhole, which lay some 75 metres beneath the surface. This essentially arrested further dissolution and bought time for DOE (Department of Energy) to prepare for the safe and orderly transfer of crude oil to another storage facility. To provide added insurance during the oil transfer stage, a “freeze curtain” was constructed in 1995. It consisted of a 54 well installation around the principal sinkhole, which froze the overburden and uppermost salt to a depth of 67 metres (Figure 9; Martinez et al., 1998). Until the mine was filled with brine and its hydrocarbons removed, this freeze wall prevented groundwater flow into the mine via the region of black salt around the sinkhole. Dealing with this sinkhole was costly. Mitigation and the removal and transfer of oil, including the dismantling of infrastructure (pipelines, pumps, etc.), cost a total of nearly US$100 million; the freeze curtain itself cost nearly $10 million.


Following oil fill in 1980-1982, the Weeks Island facility had stored some 72.5 million BBL of crude oil in abandoned mine chambers. Then in November 1995, the Department of Energy (DOE) initiated oil drawdown procedures, along with brine refill and oil skimming, plus numerous plugging and sealing activities. In 1999, at the end of this recovery operation, about 98% of the crude oil had been recovered and transferred to other SPR facilities in Louisiana and Texas; approximately 1.47 MMBL remains in the now plugged and abandoned mine workings. In hindsight, based on an earlier leak into the mine, while it was an operational mine, and the noted presence of black salt in a shear zone in the mined salt, one might fault the initial DOE decision to select this mine for oil storage. In 1978 groundwater had already leaked into a part of the mine adjacent to the sinkhole and this was forewarning of events to come (Martinez et al., 1998). Injection of cement grout into the flow path controlled the leak into the operation mine at that time, but it could just as easily have become uncontrollable and formed a sinkhole then.


Black salt zones in the now-flooded Jefferson Island Salt Mine and the 1980 Lake Peigneur collapse

The most recently risen part of the Jefferson Island stock crest is now 250 m (800 ft) higher than the adjacent flat-topped salt mass, which is also overlain by a cap rock (Figure 10). The boundary separating the spine from the less active portion of the crest is a finer-grained and a “shale-rich” anomalous zone, penetrated by the former Jefferson Island mine workings. It 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. The spine and its boundary “shear” are reflected in the topography of the Jefferson Island landscape, with a solution lake, called Lake Peigneur, defining the zone of shallower salt created by the active spine. There 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 11; Autin, 2002; Warren 2016). In the 12 hours following the first intersection the underlying mine had flooded and the lake was completely drained. The lake is about 2.4 km in diameter, has a bean-shaped configuration, with a topographic promontory along the southeast shore of the lake rising to more than 23 m above sea level and the surrounding delta plain (Figure 10).

Drainage and collapse of the lake began when a Texaco oilrig, 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 11). 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. This old sump was in contact with a zone of anomalous “black” (shear zone) salt. 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. Some 23-28 million m3 of salt had been extracted in the preceding 58 years of mine life. The rapid flush of lake water into the mine, probably augmented by the drainage of natural solution cavities in the anomalous salt zone and associated collapse grabens below the lake floor, meant landslides and mudflows developed along the perimeter of the Peigneur sinkhole, so that post flooding the lake was enlarged by 28 ha.


With water filling the mine workings, 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 tugboat and a sizeable part of the Live Oak Botanical Garden. 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 backwards into the lake depression through a connecting bayou (Delcambre Canal, aka Carline Bayou) former what was a waterfall with the highest drop in the Stat of Louisiana. Since parts of the lake bottom had slumped into the sinkhole during the collapse, the final water level in some sections of the lake was higher than before relative to previous land features. This ground movement and subsidence left one former lakefront house aslant under 12 feet of water.

Implications for other salt mines with anomalous salt zone intersections.

The Peigneur disaster 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. During ongoing operations, horizontal stress on the mineshaft near the level where the Louann Salt contacts the overlying Pleistocene Prairie Complex across a zone of anomalous salt 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 m\Mine was completed as part of a safe closure plan.

Subsidence over the nearby Avery Island salt mine (operated by Cargill Salt) has been monitored since 1986 when small bead-shaped sinkholes were initially noticed in the above mine region. Subsidence monitoring post-1986 defined a broad area of bowl-shaped subsidence, within associated areas of gully erosion (Autin, 2002). Avery mine is today the oldest operating salt mine in the United States and has been in continual operation since the American Civil War. The mine underwent a major reconstruction and a improved safety workover after the Lake Peigneur disaster. Subsidence is still occurring today along the active mine edge, which coincides with a topographic saddle above an anomalous salt zone, which is located inside the mined salt area. At times, ground water has seeped into the mine, and there are a number of known soil gas anomalies and solution dolines on the island. These are natural features that predate mining. Much of the subsidence on Avery Island is a natural process as differential subsidence occurs atop any shallow salt structure with the associated creation of zones of anomalous salt (Warren, 2016). Dating of middens and human artifacts around salt-solution induced, water-filled depressions atop the dome, shows dissolution-induced subsidence is a natural process, as are short episodes of lake floor collapse, slumping and the creation of water-filled suprasalt dolines (circular lakes). Such landscape events and their sedimentary signatures have histories that extend back well beyond the 3,000 years of human occupation documented on Avery Island (Autin, 2002).

Compared to the other salt domes of the Five Islands region of Louisiana, the Cote Blanche Island salt mine 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 salt zones, a salt overhang, and other structural complexities), and (iii) the Cote Blanche Salt Mine is surrounded by more clayey (impervious) sediments than the other Five Islands diapirs, all with sandier surrounds, perhaps allowing for lower rates of undersaturated fluid crossflow and greater hydrologic stability.

Significance

And so, today, we know that anomalous salt zones near diapirs crests are often tied to subvertical fault or shear zones, and that many are also associated with the presence of past crossflows of undersaturated waters. Across the various US Gulf Coast mines (present and past) the anomalous (“shear”) salt zones within diapirs are known to be potential problematic leakage zones and so are avoided, if possible, during mining operations. This style of black salt distribution and the potential for intrasalt leakage must be taken into account when near-crestal and shallower portions of domes are to be utilised for any fluid or waste storage. Without an understanding of the significance of such “black” salt or anomalous salt layers, there are potential undefined leakage problems within some salt structures (Looff et al., 2010; Warren 2016).

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.

Autin, W. J., and R. P. McCulloh, 1995, Quaternary geology of the Weeks and Cote Blanche islands salt domes: Gulf Coast Association of Geological Societies Transactions, v. 45, p. 39-46.

Balk, R., 1953, Salt Structure of Jefferson Island Salt Dome, Iberia and Vermilion Parishes, Louisiana: Bulletin American Association Petroleum Geologists, v. 37, p. 2455-2474.

Ghanbarzadeh, S., M. A. Hesse, M. Prodanović, and J. E. Gardner, 2015, Deformation-assisted fluid percolation in rock salt: Science, v. 350, p. 1069-1072.

Kupfer, D., 1976, Shear zones inside Gulf Coast salt stocks help to delineate spines of movement: Bulletin American Association of Petroleum Geologists, v. 60, p. 1434-1447.

Kupfer, D., 1980, Problems associated with anomalous zones in Louisiana salt stocks, USA, in A. H. Coogan, and H. Lukas, eds., Fifth Symposium on Salt (Hamburg, Germany, June 1978), v. 1: Cleveland OH, Northern Ohio Geological Society, p. 119-134.

Kupfer, D. H., 1974, Boundary shear zones in salt stocks: in Fourth Symposium on Salt. Northern Ohio Geological survey, v. 1, p. 215-225.

Kupfer, D. H., 1990, Anomalous features in the Five Islands salt stocks, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 40, p. 425-437.

Kupfer, D. H., B. E. Lock, and P. R. Schank, 1998, Anomalous Zones Within the Salt at Weeks Island, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 48, p. 181-191.

Lock, B. E., 2000, Geologic Mapping of Salt Mines in Salt Diapirs: Approaches and Examples from South Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 50, p. 567-582.

Looff, K. M., 2000, Geologic and Microstructural Evidence of Differential Salt Movement at Weeks Island Salt Dome, Iberia Parish, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 50, p. 543-555.

Looff, K. M., K. M. Looff, and C. Rautman, 2010, Salt spines, boundary shear zones and anomalous salts: Their characteristics, detection and influence on salt dome storage caverns: Paper presented at Solution Mining Research Institute Spring 2010 Technical Conference, Grand Junction, Colorado, USA, 26-27 April 2010, 23 p.

Martinez, J. D., K. S. Johnson, and J. T. Neal, 1998, Sinkholes in Evaporite Rocks: American Scientist, v. 86, p. 38.

Muehlberger, W. R., and P. S. Clabaugh, 1968, Internal Structure and Petrofabrics of Gulf Coast Salt Domes: AAPG Memoir, v. 8, p. 90-98.

Neal, J. T., 1994, Surface features indicative of subsurface evaporite dissolution: Implications for storage and mining: Solution Mining Research Institute, Meeting paper, 1994 Spring meeting, Houston Texas.

Neal, J. T., S. Ballard, S. J. Bauer, B. L. Ehgartner, T. E. Hinkebein, E. L. Hoffman, J. K. Linn, M. A. Molecke, and A. R. Sattler, 1997, Mine-Induced Sinkholes Over the U.S. Strategic Petroleum Reserve (SPR) Storage Facility at Weeks Island, Louisiana: Geologic Mitigation Prior to and During Decommissioning, SAND96-2387A.: Presented at 6th Multidisciplinary Conference on Sinkholes and the Engineering & Environmental Impacts of Karst, Springfield, Missouri, April 6-9, 1997. Sandia National Laboratories, Albuquerque, NM.

Neal, J. T., S. J. Bauer, and B. L. Ehgartne, 1995, Sinkhole Progression at the Weeks Island, Louisiana, Strategic Petroleum Reserve (SPR) Site: Solution Mining Research Institute, Fall Meeting, San Antonio, Texas, October 1995. Sandia National Laboratories, Albuquerque, NM.

Neal, J. T., and R. E. Myers, 1995, Origin, Diagnostics, and Mitigation of a Salt Dissolution Sink-hole at the U,S. Strategic Petroleum Reserve Storage Site, Weeks Island Louisiana,: Sandia National Laboratories, Albuquerque, NM. Report Sandia SAND95-0222C Paper presented at the Fifth International Symposium on Land Subsidence, The Hague, October 1995. Proceedings of the Fifth International Symposium on Land Subsidence, IAHS Publ. No. 234.

Paine, W. R., M. W. Mitchell, R. R. Copeland Jr., and L. d. A. Gimbrede, 1965, Frio and Anahuac Sediment Inclusions, Belle Isle Salt Dome, St. Mary Parish, Louisiana: American Association Petroleum Geologists - Bulletin, v. 49, p. 616-620.

Walden, W., and C. H. Jacoby, 1963, Exploration by horizon­tal drilling at Avery Island, Louisiana, in A. C. Bersticker, ed., Symposium on Salt (First): Cleveland, OH, Northern Ohio Geo­logical Society, p. 367-376.

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.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Jan-Feb. 2015: Berlin, Springer, 1854 p.

 

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[1] Pungent-smelling condiment Kala Namak (black salt) is widely used in South Asia, it consists primarily of sodium chloride with trace impurities of sodium sulphate, sodium bisulphate, sodium bisulphite, sodium sulphide. Kala Namak is also known as Himalayan Black Salt, Sulemani Namak, Bit Lobon , Kala Noon or as Bire Noon in Nepal. Its characteristic smell and taste is mainly due to its elevated sulfur content, which to a western nose is reminiscent of rotten eggs, largely due to the presence of greigite. The various iron impurities impart a brownish pink to dark violet colour to the coarse translucent crystals and, when ground into a powder, transform into a light purple to pink color.

Traditionally, mined salt was transformed from the raw natural form of salt into commercially-sold kala namak through a reductive chemical process. This heating transforms some of the naturally occurring iron oxidew and sodium sulfates in the raw salt into pungent hydrogen sulfide and sodium sulfide daughter products (along with greigite.[ The various sulphate salt impurities in the halite typify the partially recrystallised meteoric overprints that typify textures and structures in nearsurface salt residues in the Himalayan thrust belt (see Richards et al., 2015 for documentation of the geological and structural characteristics of this salt – this article can be downloaded from the publications page on this website).

Historically, the transformation of Himalayan thrust belt salt into kala namak involved firing the raw salt in a furnace for 24 hours, while sealed in a ceramic jar containing charcoal along with small quantities of harad seeds, amla, bahera, babul bark, or natron. The fired salt was then cooled, stored, and aged prior to sale. Kala namak is still prepared in this manner in northern India with production concentrated in Hisar district, Haryana. Although the kala namak can still be produced from natural salts with the required compounds, it is now common to now manufacture it synthetically using halite from non-Himalayan sources. This is done through combining sodium chloride with smaller quantities of sodium sulfate, sodium bisulfate and ferric sulfate, which are then chemically reduced with charcoal in a furnace. Reportedly, it is also possible to create similar products through reductive heat treatment of sodium chloride, 5–10% of sodium carbonate, sodium sulfate, and some sugar.


 

 

 


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