Salty Matters

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Deepsea Hypersaline Anoxic Lakes & Basins (DHALs & DHABS)

John Warren - Friday, August 31, 2018



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.


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.


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.


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.


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.


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