Source: http://www.saltworkconsultants.com/blog/tag/Red_Sea/
Timestamp: 2019-04-21 08:27:40+00:00

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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).
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.
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.
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.
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.
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.
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Most subsurface evaporites ultimately dissolve and, through their ongoing dissolution and alteration, can create conditions suitable for metal enrichment and entrapment in subsurface settings ranging from the burial diagenetic through to the metamorphic and igneous realms. This article looks at a few examples tied to halokinesis, a more comprehensive set of examples and more detailed discussion is given in Chapters 15 and 16 in Warren (2016). Because most, if not all, of any precursor salt mass that helped form these metalliferous deposits via dissolution, has gone, the resulting metal and other accumulations tend to be at or near the edges of salt basins, or in areas where most or all of the actual salts are long gone (typically via complete subsurface dissolution or metamorphic transformation, so that only breccias, weld and indicator mineral suites remain).
A lack of a direct co-occurrence with evaporite salts is perhaps why the metal-evaporite association is not recognised by some in the economic geology community. The significance of disappearing salt masses in focusing and enhancing metal precipitation, via the creation of chloride-rich and sulphate-rich brines, may not be evident without the conceptual tools needed to recognise the former presence of evaporites, post-salt halokinetic structural geometries, and meta-evaporite mineral associations.
The various ore tonnage-grade plots in Warren (2016), shows that many metal accumulations with an evaporite association tend to plot at the larger end of their respective deposit groupings.
Evaporite dissolution helps create "prepared ground."
I am not saying all large metal accumulations require evaporites or the highly-saline subsurface fluids that they can generate. Although, some recent papers do argue for a widespread role of evaporites in a Pb-Zn association (Fusswinkel et al., 2013; Wilkinson et al., 2009) and in sedimentary redbed copper deposits (Rose, 1976; Hitzman et al., 2010).
Typically, the conceptualisation of an evaporite in the economic geology literature is as a bedded evaporite and brine source (Figure 1). Likewise, this article and the relevant chapters in Warren (2016) detail a number of megagiant ore deposits where dissolving evaporite bodies have contributed in some way to a metal accumulation (Table 1). However this current article, like Warren 2016) focuses on the mechanisms and indicators tied to a halokinetic-ore association. Halokinesis is an aspect of evaporites that is not widely discussed in the field of ore deposit models.
Not all sediment-associated ore deposits are associated with evaporites. Only in those ore deposits classified as anorogenic and/or continental margin can subsurface evaporite masses can be involved in the same unusual concentration and alteration conditions that lead to the creation of metalliferous ore deposits (evaporite associations are indicated by E in Figure 2). At other times and locations hydrothermal mineral salts, especially anhydrite (CaSO4)which can supply sulphur as it dissolves, can be an integral part of the ore accumulation, but their occurrence may be unrelated to aridity. Hydrothermal anhydrite and other burial/magmatic hydrothermal salts tend to form in high salinity conditions inherent to the ore-forming environment and not necessarily to the presence of precursor evaporites; as in the formation of carbonatites (e.g. Afrikanda and Bayan Obo; Wu, 2008), or pegmatites and some IOCG deposits (hydrothermal anhydrite is indicated by HA in Figure 2). In some such hot subsurface settings the role of any nearby buried true” evaporite may be, via its dissolution or alteration, to aid in the creation of highly-saline high-temperature basinal brines (Chapter 16; Warren, 2016). According to whether the resulting brines are chloride or sulphate-rich, they can act as either enhanced metal carriers or fixers.
The role of evaporites creating metalliferous ores is two-fold; 1) In solution (halite-dominant precursor) they can act as chloride-rich metal carriers and 2) Locally, asCaSO4 beds or masses alter and disintegrate, their dissolution products, especially if trapped, can supply sulphur (mostly via bacteriogenic or thermogenic H2S). Dissolutional interfaces set up chemical interfaces that act as foci during brine mixing so manufacturing conditions suitable for precipitation of metal sulphides or native elements. As a consequence, most evaporite-associated ore systems tend to epigenetic, rather than syngenetic. Subsurface salt beds and masses are merely the solid part of a sizeable ionic recycling system, dissolved metals are another part, and zones of mixing between the two are typically sites where metal sulphides tend to gather.
At the world-scale, both evaporite and ore systems are driven by plate tectonics. Halite-dominated sequences, deposited in the drawdown basin centres, tend to dissolve in burial, and so supply chloride ions to the brine system. Salt beds that are thick enough tend to flow and thus focus the upward and centripetal passage of basinal and hydrothermal fluid flows. Dissolving gypsum or anhydrite beds, typically deposited higher on the basin platform or diagenetically accumulated along salt dissolution edges and salt welds (touchdowns) can supply sulphur, via bacterial or thermochemical sulphate reduction, while simultaneously focusing the subsalt metalliferous brine flows into the precipitation interface.
When the chemistries of the dissolving salt beds and the metal carriers interact so that redox fronts, salinity contrasts, and other precipitative interfaces are set up, an ore deposit can form. Thus, in base and precious metal exploration in evaporitic terranes, we are ultimately searching for those parts of a subsurface ionic cycling system where the salt dissolution, salt beds and metal systems have interacted to create economic levels of metalliferous precipitates.
Conceptually, this evaporite-related notion of regional fluid flow in a sedimentary/metasedimentary host is somewhat different to the internal process and local mineralised halo models that dominate our understanding of those world-class ore deposits related to the interior workings of igneous systems. The latter is known as an orthomagmatic system where internal igneous processes of fractional crystallisation and liquid immiscibility largely control ore formation. Ores are deposited in an evolving framework of world-scale tectonics and magmatism across time, from Archaean greenstones to those of present-day sialic plate tectonism. Examples, where buried evaporites have been assimilated into a magma chamber, are discussed chapter 16 in Warren, 2016. Then there are the various ore deposits that are external to (paramagmatic) or unrelated to the emplacement of igneous bodies (nonmagmatic). In both cases, the mineralisation is typically part of an ongoing long-term sedimentary burial history, tied to dissolving and flowing salt masses and associated hydrothermal circulation.
Evidence for hydrothermally-induced low-moderate temperature mineralisation is often best preserved in textures in the hydrothermally altered rock matrix, typically located outside the actual ore deposit (in its hydrothermal alteration halo). From the hydrothermal fluid perspective, one should see the role of evaporites and metal sulphides as each contributing its part to a larger scale “mineral systems” paradigm; much in the same way as, in a petroleum system, the integration of concepts of source, carrier, seal and trap are fundamental requirements to understand and predict economic oil and gas accumulations.
This holistic ore systems approach is not fully encompassed in some economic geology studies that use sequence stratigraphic sedimentological approaches for ore deposit prediction in greenschist terrains (Ruffel et al. 1998; Wilkinson and Dunster, 1996). In my opinion, this approach can shift the interpretation paradigm too far into the depositional realm. The problem with classic sequence stratigraphic criteria, when trying to understand ore genesis, is that sequence stratigraphy does not handle well the concept of a mobile ephemeral subsurface salt body that climbs the stratigraphy via autochthonous and allochthonous process sets (halokinesis). As the salt flows, it dissolves and so brings with it the associated epigenetic influences of brine-driven diagenesis and metasomatism.
Current sequence stratigraphic paradigms in the economic geology realm are dominated by the assumption that the geometry of the units in the depositional system, and associated fault characteristics, are relatively static within the buried sediment prism. Yet, in terms of most sediment-hosted hypersaline ore deposits, what is most important in understanding the metal-evaporite association is the understanding of; 1) Evaporite dissolution and halokinesis, 2) Migration of subsurface fluids, 3) Creation of shallower or lateral-flow redox fronts along with, 4) Opening and closing of fault/shear focused fluid conduits, typically tied to, the coming and going of bedded and halokinetic salts. These factors, rather than primary sediment wedge geometries, are the dominant controls as the mineralising system passes from the diagenetic into the metamorphic realm.
It is interesting that in a benchmark paper, discussing and classifying the world’s ore deposits in a plate-tectonic-time framework, Groves et al., (2007) list almost all the major ore categories shown in Figure 2b as belonging to the group of “...sediment-hosted deposits of non-diagnostic or variable geodynamic setting.” Into this category, they place all stratiform to stratabound sediment-hosted deposits with variable proportions of Pb, Zn, Cu (including Zambian Copper belt, Kupferschiefer and SedEx deposits). They go on to note (p. 26) that, although there is general agreement that the majority of these various deposits formed during active crustal extension, either in intracratonic rift basins or passive margin sediment hosts, there is considerable controversy concerning their broader scale tectonic setting at the time of mineralisation and the driving force for hydrothermal fluid flow at the time of their mineralisation.
Perhaps this lack of model specificity in the varied interpretations of sediment-hosted deposits reflects the fact that one piece of significant information is missing from many ore genesis models. Namely, that the greater majority of these poorly classified sediment-hosted deposits sat atop, or adjacent to, or beneath, what were once thick evaporite sequences (Table 1). In many cases, the salt mass is long gone. It was the dissolution of these salt masses, either bedded or halokinetic-allochthonous, that focused much of the ore-fluid flow in the sedimentary-diagenetic realm. The loss of salt as the basin sediments passed from the low temperature diagenetic into the metamorphic realm, and as the metalliferous fluid flow was focused into permeable conduits about, below or above the dissolving and retreating, or flowing salt edges, is how salt-related ore deposits form.
This is why the majority of these salt-aided deposits tend to occur outside salt basins that retain substantial salt masses still in the diagenetic realm. The deposits are a response to the dissolution and flow of evaporites, or the residual seawater bitterns created in underlying and subjacent settings as the salt beds were deposited, not to the presence of actual undissolved primary evaporite masses. As we see in Proterozoic and Archaean meta-evaporites and most Precambrian evaporite associations, the original salt mass is long gone from the hosting succession, via varying combinations of halokinesis, dissolution and metasomatism (Warren, 2016, Chapter 13; Salt Matters blog, August 28, 2016).
Ore deposits of Precambrian tend to be linked to evaporite alteration products and residues and rarely preserve actual sedimentary salts (other than local remains of minor hydrothermal anhydrite). In younger Phanerozoic deposits, such as Kupferschiefer, the Atlantis II deep and Dzhezkazgan, portions of actual salt (brine source) can remain in the more deeply buried parts of the basin.
Metal sulphide precipitates are not rare or unique in the subsurface diagenetic fluid milieu, what is essential in the prediction of ore-grade levels of metal sulphide buildups is understanding where and why the metal precipitation system is focused into particular structurally-controlled positions and encompass time frame/fluid volumes sufficient to build an ore deposit.
That is, evaporite-associated ore deposits are no more than ancient subsurface hydrology-specific associations where the precipitation system was stable enough, for long enough, to allow higher, ore-grade levels of metals sulphides to accumulate from carrier brines at particularly favourable and stable chemical and temperature interfaces. As such, metal precipitation sites are part of an ore-forming process set, spread across the epigenetic and syngenetic realms (Table 1). They are part of the regional evolution of the fluid plumbing from the time of deposition, into burial, and on into the realm of metamorphic transformation. This means to understand the ore system tied an evaporite-entraining system holistically; one must integrate local ore paragenesis with various aspects of the basin-scale geology, sedimentology, sequence stratigraphy, diagenetic-metamorphic-igneous facies, fluid flow conduits and structural evolution of the evaporitic basin.
Stratabound deposits of copper (±Ag), hosted by variably-dipping continental clastic sedimentary rocks, occur in Central Andean intermontane basins and are known to postdate compressive deformation/uplift events in the region (Flint, 1986, 1989). The deposits are relatively small with variable host-rock depositional ages and include; Negra Huanusha, central Peru (Permo-Triassic); Caleta Coloso, northern Chile (Lower Cretaceous); Corocoro, northwestern Bolivia (Oligo-Miocene); San Bartolo, northern Chile (Oligo-Miocene); and Yasyamayo, northwestern Argentina (Miocene-Pliocene).
The Corocoro area has produced the largest amount of copper in these Andean examples, something like 7.8 million tonnes of copper at a grade of 7.1% (Cox et al., 2007). The location of mineralisation is controlled by structurally-focused redox fronts in bedded sediment hosts, which abut a steeply-dipping translatent thrust fault (Figure 3). Deposits are irregular, usually elongate lenses of native metal, sulphides, and their oxidation products. Typically, deposits are hosted in alluvial fan and playa sandstones or conglomerate facies that also contain abundant gypsum and lesser halite. The undersides of some copper sheets at Corocoro even preserve mudcrack polygons and bed-parallel burrow traces (Savrda et al., 2006). Ore mimicry of mudcracks is not a feature controlled by on-for-one-replacement of organic material deposited in a sandstone; rather it is following pre-existing permeability/redox contrasts.
Corocoro deposits have been mined sporadically since they were first exploited by the local Indians, prior to the Spanish invasion in the 16th century and were largely exhausted in half a century of more intense mining operations that began in 1873 (Figure 3). Sandstone and conglomerate matrices show evidence of bleaching and leaching of the original redbed host with numerous red-greybed redox interfaces visible in the mined sequences. Ore minerals (dominantly native copper) are secondary fills within secondary intergranular pores created by the dissolution of earlier carbonate and sulphate masses and intergranular cement. Twelve grey sandstone beds, which were host to the long worked-out native copper ores, occur within a stratigraphic thickness of 60 m, in a unit known as the Ramos Member that still hosts abundant CaSO4 as gypsum (Figure 3).
Ores are stratabound, but not necessarily stratiform, and the larger masses of native copper are typically shallower and present as vein fills. Sometimes the copper pseudomorphs large orthogonal-ended aragonite prisms, which can be several centimetres across. There are two main styles of mineralization; 1) Ore minerals as a matrix to stratiform detrital silicates, typically low dipping and commonly highlighting primary sedimentary structures, such as cross stratification, 2) Ores in stacked channelized sand bodies, that show steep dips in structurally complex and folded zones with local brecciation (Figure 3). Native copper commonly fills thin laterally extensive sheets in tectonic fractures in the limbs of tight folds. Ljunggren and Meyer (1964) interpreted these folded diagenetic sheets of copper as a remobilization products precipitated during deformation of earlier matrix-pore filling copper.
The ore-hosting clastic horizons are consistently located in the highly gypsiferous Vetas Member of the Ramos Formation, which was deposited as redbeds in braidplains or fluviodeltaic playa margins centripetal to the edges of saline evaporitic lakes that were accumulating gypsum and halite (Figure 4; Flint, 1989). Abundant gypsum is still present in the Ramos Member as nodules and satinspar vein fills. Both are secondary evaporite textures likely implying the dissolution of previously more voluminous CaSO4 and NaCl beds and masses. Gypsum along with celestite are the most common gangue minerals associated with native copper veins in all the Corocoro deposits (Singewald and Berry, 1922). In the geological analysis of the first two decades of last century, the copper-bearing beds of the westerly-dipping series were called "vetas" and those of the easterly-dipping beds "ramos" and, as a matter of convenience, the names became attached to the rocks themselves. The term "veta" is Spanish for vein and "ramo" the Spanish for branch (native copper). The 1922 paper by Singewald and Berry noted that the veta horizons were traceable continuously for over 5 km in outcrop, but they found no apparent primary trends related to ramos outcrops (Figure 3).
Six mineralised layers of each kind were in exploited in mining during the first two decades of last century, the thicknesses of which varied from a few centimetres to 7 meters (Figure 3). Sheets and masses of native copper, called charque, were up to 600 pounds in weight, but more significant volumes of copper were extracted from vetas sandstones where copper was found as diffuse minute grains, pellets, or granular masses of the native metal. Associated with the enriched copper zones were more oxidised minerals as malachite, chrysocolla, azurite, domeykite, and chalcocite. Singewald and Berry (1922) noted gypsum and salt were the principal gangue minerals, while silver minerals were rare. The vetas sediment hosts tended to be coarser grained, often conglomeratic; whereas the ramos sediment hosts were finer-grained with copper present as smaller particles and masses.
The currently accepted interpretation of the Corocoro copper is that it formed during early diagenesis within a saline playa depositional environment, and in combination with dissolution of the adjacent bedded lacustrine evaporites (Figure 4). This bedded combination is thought to have controlled the formation, transport and precipitation of the copper ore (Flint, 1989). Playa sandstones, sealed between impervious evaporitic mudstone layers, created the plumbing for focused metalliferous fluid migration toward the basin margin. It is argued that the carbonaceous material at Corocoro was likely concentrated in the sandstones and conglomerates and not in the shalier members of the sedimentary sequence (Eugster, 1989).
The organics were considered strata-entrained as primary plant matter (e.g. spores) preferentially in the sandstones, along with later possible catagenic/hydrothermally cracked products migrating as hydrocarbons out of the basin. This created locally reducing pore environments in the aquifers wherever these reduced fluids met with somewhat more oxidising updip pore waters. This updip migration of saline reducing waters, in combination with sulphur supplied as H2S from the adjacent dissolving calcium sulphate beds and nodules, as well as from dissolving intergranular sulphate cement, precipitated copper in the newly created secondary porosity. The pore water chemistry and flow hydrology of this sandstone-hosted Cu system is thought to show many affinities with diagenetic uranium-redox precipitating systems, as defined by Shockey and Renfro (1974).
However, there is, in my mind, a possible anomaly in this model, which assumes organics were deposited in fluvial sandstones at the time of deposition. It is highly unusual to have higher plant material accumulating in large volumes in sandstone in a setting that is sufficiently arid and oxidising to precipitate ongoing interbeds of halite and gypsum. Such settings are typically too dry to allow abundant higher plant growth. Also, groundwaters that are flowing basinward through bajada sandstones in Neogene sediments of the Andes are ephemeral or too oxidising to facilitate the long-term reducing conditions needed to preserve significant volumes of high plant remains in the sandstone aquifers.
What is also interesting in this sedimentological/diagenetic model of Tertiary age cupriferous redbeds deposits in the Andes, centred on Corocoro, but not considered in any detail in the published literature base, is the question..., What controlled the folding, and the associated brecciation and perhaps even subsurface brine interfaces responsible for the Cu precipitation? All the stratabound Bolivian Cu deposits accumulated in sediment hosts that were deposited in fault-bound intermontane groundwater sumps. All are located in hydrologic lows in the crustal shortening tectonic scenario that typifies the Tertiary history of the Andes.
The variable ages of the host sediments and the predominance of evaporite indicators including gypsum in outcrop (often as diagenetic residues, not primary, features in the fluvial hosts) and all intimately tied to the Corocoro ore forces the question...., “was the fluid focusing driving the Cu precipitation a response to compression-driven halokinesis in an evolving salt-lubricated thrust belt?” Did this on-ground scenario occur in a halokinetic hydrology, that was possibly related to a combination of thrust-driven telogenesis, redox setup, evaporite dissolution and aquifer focusing of brines with dissolution aiding local slumping? This, along with associated strike-slip prisms, could better explain the stability of redox interfaces in sandstone aquifers across timeframes needed to accumulate significant native copper volumes. After all, most of the ore textures are passive precipitates, mainly in pre-existing porosity. If so, perhaps these deposits are not a variation on a roll-front uranium theme, which is predicated on dispersed primary organic material in the host sandstones (Shockey and Renfro, 1974).
When one plots the position of Corocoro and other redbed copper across the region, the 1000-lb gorilla that has been standing in the corner of the room for the past century becomes obvious. The Corocoro redbed copper deposit is located on a salt-cored fault system linked across less than a kilometre to an outcropping gypsum-capped remnant of a salt diapir which crosscuts the anticlinal axis of a saline redbed/greybed Corocoro sequence and ties to the saline decollement of the Corocoro Fault (Figure 5). The same tie to salt-cored decollement and diapir proximity is true of other nearby redbed copper deposits to the south-southeast, such as Veta Verde and Callapa. It is highly likely that the saline fluid interfaces forming the redbed Cu deposits of Corocoro, Veta Verde and Callapa were halokinetically focused. A similar-salt lubricated set of thrusts and strike-slip faults typifies halokinetic anticline outcrops in Central Iran.
It is highly likely that much of the structuration that is controlling Corocoro ore positioning is a response to salt flow related uplift, brine conduits and fracture creation. Metal precipitation occurred at redox interfaces induced and controlled by regional salt-lubricated compressional tectonism, and the associated salt-structuration has driven the brine-interface redox hydrology.
Work by Rutland (1966) did make an observation that the Corocoro ore deposits are related to an unconformity between the Ramos and Vetas Formations. Previously, the unconformity was interpreted as directly due to the outcrop of the Corocoro Fault. He noted that the fault and the unconformity were one and the same. In the 1960s there was no notion of a salt weld but it was nonetheless a highly astute observation by Roy Rutland. He went on to note a similar unconformity is tied to the growth of the Chuquichambi salt diapir, some 100 km southeast of Corocoro. Unfortunately, the halokinetic implications of Rutland's work were not considered 20 years later in Flint's key 1989, paper inferring a mostly clastic sedimentological origin for the Corocoro and other similar SSC deposits.
A possible halokinetic/weld association also leads to the question... Were the salt lakes, that are considered an integral part of the depositional and saline ore-precipitation systems at Corocoro by Flint, also a response to dissolution of the same nearby diapiric structures, when they were active in the mid to late Tertiary? This tie, between diapir/weld brines sourced in the drainage hinterland and bedded evaporite - lacustrine mud interbeds accumulating in the groundwater outflow sumps, is the case with groundwater inflow for the Salar de Atacama infill, as it is in other Quaternary salt lakes in the region. The are many diapir remnants across the Andes region. It seems that the Corocoro style of Cu mineralisation is perhaps another example of suprasalt redox focusing in a halokinetic setting.
Whether the halokinetic scenario, or the currently accepted non-halokinetic bedded arid-lacustrine evaporite scenario, explains the Cu mineralisation Corocoro is yet to be tested. But in terms of future copper exploration for similar deposits, it probably requires an answer. A halokinetic association offers an exploration targeting mechanism, utilising satellite imagery and aerial/gravimetric data, prior to the acquisition of on-ground land positions and geochemical surveys.
This material on the HYC deposit will be expanded upon in an upcoming paper by Lees and Warren (in prep.). Before mining, the McArthur River (or HYC) Pb-Zn-Ag deposit, contained 227 million tonnes of 9.2% Zn, 4.1% Pb, 0.2% Cu and 41 ppm Ag (Logan et al., 1990; Pirajno, and Bagas, 2008). The deposit is hosted in the HYC Pyritic Shale member and lies adjacent to the Emu Fault in the McArthur Basin and adjacent to what are currently sub-economic base metal deposits in the Emu Fault zone known as the Cooley II and the Ridge II deposits (Figure 6a). Across all these deposits, major ore sulphides are pyrite, sphalerite and galena, with lesser chalcopyrite, arsenopyrite and marcasite. The mineralised region has an area of two km2 and averages 55 m in thickness (Figure 6b). It is elongated parallel to the major Emu growth Fault, which lies 1.5 km to the east, but is separated from the main ore mass by carbonate breccias of the Cooley Dolostone Member (Figure 6a-d).
The sequence at McArthur River comprises dolomites of the Emmerugga Dolostone (with the Mara Dolostone and Mitchell Yard members), overlain by the Teena Dolostone with abundant aragonite splays indicative of a normal-marine tropical Proterozoic carbonate. Overlying the Teena Dolostone in the vicinity of the HYC deposit is the somewhat deeper water Barney Creek Formation and its equivalents, containing the W-Fold Shale member, while the ore is hosted in carbonaceous shales, with multiple lenses of fine-grained galena-sphalerite-pyrite, separated by inter-ore sedimentary breccias (Large et al., 1998). This unit contains numerous sedimentary features indicative of a deeper-water anoxic setting. For example, comparison with d13C values from isolated kerogen in the HYC laminites confirms that n-alkanes in Bitumen II are indigenous to HYC, indicating that the deposit formed under euxinic conditions. This supports a generally-held model for Sedex deposits the region, whereby lead and zinc reacted in a stratified water column with sulphide produced by bacterial sulphate reduction (Holman et al., 2014).
The ore-hosting organic-rich 1,643-Ma HYC Pyritic Shale Member of the Barney Creek Formation is much thicker in the HYC sub-basin than elsewhere in the Batten Trough Fault Zone (e.g., Glyde River Basin) and consists mainly of dolomitic carbonaceous siltstones (Figure 7; Davidson and Dashlooty, 1993; Bull 1998). I would argue this thickening reflects a combination of long-term local basinfloor subsidence, related to salt withdrawal, and brine stratification due to ongoing salt dissolution and focused outflow. Indicators of former salt allochthon tiers are widespread in the vicinity of the HYC deposit, but are absent in the Glyde River Basin.
In the HYC mine area, the ore interval is overlain by the HYC pyritic shale member and made up of pyritic bituminous and dolomitic shales and polymict breccias (Figure 7). Importantly, when contacts are walked out in outcrop, the polymict breccias are significantly transgressive to bedding, while drilled intersections in the vicinity of the HYC deposit and in the mine itself show the breccias are stratabound. Another interesting feature of these breccias is that they can contain mineralised clasts. More broadly, a variety of sedimentary breccias occur throughout the Barney Creek Formation stratigraphy, especially along the eastern margin of the HYC half graben and tend to pass updip into the breccias of the Cooley Dolostone (Figure 6a).
Williams 1976, defined three breccia types (I, II and III) in the HYC area. Type I breccia beds occur in the lower half of the HYC Pyritic Shale Member and contain clasts characteristic of lithologies in formations of the McArthur Group below the Barney Creek Formation (Table 2). In the northern end of the sub-basin, the breccias are of a chaotic nature with no sorting and minor grading of clasts (Figure 6b). The underlying shale beds are frequently contorted and squeezed between the breccia fragments, which reach a maximum size of approximately 10 m. Toward the south, the thickness and maximum clast size of individual breccia beds decrease (Figure 6b). All breccia units are thickest adjacent to the Emu Fault Zone and likely record sediment sinks controlled by rapid fault-controlled basin subsidence during Barney Creek time. Inter-ore breccias amalgamate and thicken to the north-north-east of HYC, and occupy a position toward the foot of what is interpreted as a more substantial breccia lens, dominated by sediment gravity ﬂow deposits (Figure 6d; Logan et al., 2001).
In a subsequent study, Ireland et al. (2004a) identified four distinct sedimentary breccia styles within Type I breccias: framework-supported polymictic boulder breccia; matrix-supported pebble breccia; and gravel-rich and sand-rich graded turbidite beds (Table 2). The boulder breccias can be weakly reverse-graded and show rapid lateral transition into the other facies, all of which are interpreted as more distal manifestations of the same sedimentary events. The flow geometry and relationships between these breccia styles are interpreted by Ireland et al. (2004a) to reflect mass-flow initiation as clast-rich debris flows, with transformation via the elutriation of fines into a subsequent turbulent flow from which the turbidite and matrix-supported breccia facies were deposited.
All the Type 1 mass-flow facies contain clasts of the common and minor components of the in-situ laminated base-metal mineralised siltstone. Texturally these clasts are identical to their in-situ counterparts and are distinct from other sulphidic clasts that are of unequivocal replacement origin. In the boulder breccias, intraclasts may be the dominant clast type, and the matrix may contain abundant fine-grained sphalerite and pyrite. Dark-coloured sphaleritic and pyritic breccia matrices are distinct from pale carbonate-siliciclastic matrices, are associated with a high abundance of sulphidic clasts, and systematically occupy the lower parts of breccia units. Consequently, clasts that resemble in-situ ore facies are confirmed as genuine intraclasts incorporated into erosive mass flows before complete consolidation. Disaggregation and assimilation of sulphidic sediment in the flow contributed to the sulphide component of the dark breccia matrices. The presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows. That is, the presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows (Ireland et al., 2004a).
Type II breccia beds occur throughout the HYC Pyritic Shale Member but are most common in the upper half of the Member. Clasts are predominantly grey dololutite which occasionally contain radiating clusters of acicular crystal pseudomorphs (“coxcos”) indicative of tropical Proterozoic shelf carbonates. The clasts are similar to lithologies in the Emmerugga and Teena Dolomites and are considered to have been derived from these formations. A characteristic of this breccia type, which differentiates it from Type I and III breccias is the absence of green and red clasts, signifying that clasts in Type II breccias were not derived from the Tooganinie or lower formations, but mostly derived by erosion and collapsed of updip shallow-water cemented shelf carbonate layers. Type II breccias lack the well-developed grading seen in Type I breccias. Isopach maps (Figure 6c) and maximum clast-size plots of individual breccia beds show a close correlation and indicate the type II breccias dominate in the southeast of the HYC subbasin.
Type III breccia beds are confined to the uppermost breccia unit of the HYC Pyritic Shale Member in the HYC sub-basin and are equivalent to the Upper Breccia of Murray (1975). This unit consists exclusively of Type III breccias with the exception of several shale beds near the base. The top of the Upper Breccia is not exposed in the sub-basin, and the unit reaches a maximum known thickness of 210 m. Clasts within the breccias are completely chaotic, and there is no recognisable grading or sorting. Clasts range in size from a few millimetres up to several tens of metres. The fragment lithologies are identical to those in the Type I breccias with the notable exception that they also contain clasts of sandstone, quartzite and potash-metasomatized quartz dolerite—lithologies that are characteristic of the underlying Masterton Formation. The fragments are therefore considered to be derived from the McArthur Group (below the Barney Creek Formation) and the Masterton Formation. According to Walker et al. (1977), the most likely source of the clasts from the Masterton Formation is erosion uplifts and horsts in the Emu Fault Zone. But the same authors also state the exact source area and the direction of movement of the clasts could not be identified. In my opinion, Type III breccias are salt-ablation derived and so contain a variety of clasts lithologies plucked by the rising salt as it rose toward the surface to feed an at-seafloor allochthon.
More broadly, breccias of the updip Cooley Dolostone member, that interfinger and also overlie the HYC deposit (Figure 6a) are usually regarded as part of the Barney Creek Formation. The Cooley Dolostone is interpreted, historically, as a talus slope breccia (Walker et al. 1977, Logan 1979), containing clasts eroded from the Teena and Emmerugga Dolostones. Hinman (1995) regarded the Cooley Dolostone as a tectonic breccia, formed along reverse faults within the steep to overturned, brittle dolomitic lithologies of Teena, Mitchell Yard and Mara Dolostones(members of the Emmerugga Dolostone) as they were overthrust against and over Barney Creek Formation lithologies. Perkins & Bell (1998) interpret the Cooley Dolostone as an in situ alteration body, contiguous with, and derived from, the HYC sequence, rather than being separated from it by a thrust fault. I interpret much of the Cooley as a salt allochthon breccia derived from a salt-cored basin edge fault system, now evolved into a salt weld (Table 2).
Regional-scale potassic alteration of Tawallah Group dolerites and sediments were documented by Cooke et al. (1998), Davidson (1998, 1999). These authors describe fluids responsible for this alteration as oxidised, low-temperature (100˚C), saline (> 20wt % NaCl equiv), Na-K-Ca-Mg-rich brines, and argue that the high salinities and the presence of hydrocarbons are consistent with brine derivation from nearby evaporitic carbonates during diagenesis.
I suggest that saline fluids feeding these haloes came not from the dissolution of evaporites in adjacent bedded carbonate hosts, but from the decay of former fault-fed thick salt allochthon tongues in positions that now are indicated by salt allochthon breccias. These breccias tie back to what were salt-lubricated fault and salt welds. The presence of salt and diagenetic haloes in these features focused tectonic movement and fluid supply in both initial extensional and subsequent compressional stages. As such, this interpretation supports a salt dissolution origin of the brine origins proposed by both Logan (1979) and Hinman (1995). The difference with their interpretations is that I envisage the brine being derived during salt flow emplacement and dissolution, tied to focused fault conduits in a mobile, suprasalt fault complex, atop or adjacent to the now-dissolved flowing and tiered salt mass. I do not think the nearby platform carbonates (with coxcos and smooth-walled cherts) ever contained significant volumes of primary evaporites.
Worldwide and across deep time, most halokinetic basinwide evaporite associations are typified by an initial extensional and loaded set of diapirs evolving into salt-cored fault welds, with subsequent reactivation of these features in compression (Warren, 2016; Chapter 6). Such a framework typifies long-term salt tectonics with inherently changing structural foci across most Phanerozoic halokinetic salt realms, as in the North Sea, the Persian (Arabian) Gulf and most circum-Atlantic salt basins. It is indicative of continental plate-edge evaporites caught up in the Wilson cycle (Warren, 2010).
Near the HYC deposit, Mn-enrichment, particularly of dolomite and ankerite in the W-fold Shale beneath the ore zone, is considered to be related to exhalation of Mn-bearing brines, associated with rifting and basin deepening, before the onset of zinc-lead mineralisation (Large et al. 1998). This too, is consistent with the salt-focused mineralisation hydrology of diagenetic ferroan and Mn-bearing hydrologies of the modern Red Sea halokinetic deeps (Schmidt et al., 2015) and the Danakhil depression in the Quaternary, when it was a marine-fed saline system (Bonatti et al., 1972).
In the area to the east of to McArthur River HYC basin, a number of currently sub-economic Zn-Pb-Cu deposits occur, typified by the nearby Ridge and Cooley deposits (Figure 6a; Walker et al. 1977; Williams 1978). Both are similar to the Coxco deposit, being described as MVT deposits mainly hosted by dolomitic breccias, but with minor, shale-hosted concordant mineralisation in the Ridge II deposit (Figure 8; Williams 1978). Likewise, the Coxco deposit contains several million tonnes at 2.5% Zn and 0.5% Pb, in coarse-grained, stratabound galena-sphalerite-pyrite-marcasite, hosted by dolomitic breccias containing clasts of the Mara Dolostone Member, Reward Dolostone, and the Lynott Formation of the McArthur Group, within the Emu Fault Zone (Walker et al. 1977, Walker et al. 1983). Mineralisation comprises veins, “karst” and dissolution breccia fill likened to Mississippi Valley Type (MVT) mineralisation (Walker et al. 1977).
According to Williams (1978), the Emmerugga Dolostone hosts the discordant mineralisation of Cooley II deposit, while Cooley Dolostone breccias contain the Ridge II deposit (Figure 8). The Emmerugga Dolostone at Cooley II consists of massive to laminated dolostone and contains carbonaceous matter, stromatolites, oncolites, and ooids, indicating that it was deposited in a shallow-water normal-marine environment with high biologic productivity. Similarly, the Cooley Dolostone host at Ridge II is a breccia composed of randomly oriented dolostone clasts varying in diameter from a few millimetres up to several tens of metres. Some clasts have near-identical lithologies to those comprising the Emmerugga Dolostone, whereas others contain coxcos and were likely derived from the fragmentation of Teena Dolostone. The Cooley Dolostone breccia contains little depositional matrix. Clast boundaries are marked by sudden changes in features such as dolostone type and bedding-core angles, indicating that the breccia was mostly clast-supported at the time of formation. Most interestingly, drilling in the vicinity of the deposit (DDHR210) intersected a large clast of “out of sequence” dolerite (Figure 8a). Similar large salt-buoyed clasts (up to 100’s meters across) composed of Eocene dolerite occur in the salt allochthon breccias at Kuh-e-Namak-Qom (Salty Matters blog, March 10, 2015).
Two major phases of crosscutting brecciation in the area are recognised by Williams (1978) in drill core samples of discordant mineralisation from both the Emmerugga and Cooley Dolostone hosts. First generation breccias, formed during the earlier phase of brecciation, consist of angular clasts of dolostone (< 1 mm to at least 1 m in diameter) in a dark colored matrix of tiny ( < 1 µm to 20µm) anhedral dolomite grains, disseminated euhedral pyrite crystals (<50 µm in diameter) and reddish brown carbonaceous matter). The identical nature of the first generation breccias in both the Emmerugga and Cooley Dolostone hosts suggests that brecciation occurred simultaneously in both, via the same mechanism (Williams, 1978). At the time this interpretation was made, there was no “data” (paradigm) available to determine whether the brecciation in the Cooley Dolostone occurred in situ or whether it took place in the dolostone before its removal from the Western Fault Block. Today, we would likely interpret these features as reworked salt ablation breccias on the deep seafloor with infiltrated suspension clays and early-diagenetic pyrite.
Second generation breccias, formed during a later phase of brecciation, consist of angular clasts of first-generation breccias (< 1 mm to at least 10 cm in diameter) in a matrix of either veins filled with sulphide minerals and dolomite, or fine-grained (10 µm to 100 µm in diameter) anhedral dolomite grains, disseminated to massive sulfide minerals, small (on the average 500 µm x 20 µm) interlocking laths of barite or dolomite pseudomorphs after barite, and brown carbonaceous matter (Williams, 1978). Second generation breccias, although coincident with the first generation breccias, are less widespread than the earlier breccias. Again, according to Williams (op. cit.), the similarity of the second generation breccias in both the Emmerugga and Cooley Dolostones suggests a common origin. Again, they concluded there was no “data” (paradigm) available to establish the time of this brecciation relative to the deposition of the Cooley Dolostone. I would argue these “second generation” breccias represent a less distally reworked salt ablation breccia, possibly with interspace anhydrite and gypsum at the time they formed. These calcium sulphate phases facilitated the shallow subsurface emplacement of metal sulphides via bacterial or thermochemical sulphate reduction, in a way not too dissimilar to the mechanisms emplacing Pb-Zn at Cadjebut or Bou Grine ores in Tunisia (Warren and Kempton et al., 1997; Warren 2016; Chapter 15).
The origin of the HYC deposit and adjacent subeconomic mineralised accumulations is still somewhat controversial and equivocal (Figure 6a; Ireland et al. 2004a,b; Perkins and Bell, 1998; Logan, 1979; Walker et al., 1977). Large et al. 1998 summarised the alternative models: 1) a sedimentary-exhalative (‘sedex’) model was proposed by Croxford 1968 and Large et al. 1998; while, 2) a syndiagenetic subsurface replacement model was introduced by Williams 1978; Williams & Logan 1986; Hinman 1995 and Eldridge et al. 1993, the latter based on sulphur isotopes. In my opinion, a third factor, namely a now-dissolved salt allochthon system, should be considered in interpretations of ore genesis and associated breccias. I interpret ore-hosting laminites of HYC deposit as DHAL laminites, and the Ridge II and Cooley II were hosted in updip regions once dominated by salt tongues and salt ablation breccias within a fault-fed salt allochthon complex surrounded by updip normal-marine shoal-water platform carbonates (Figure 9).
That is, all three deposits are related to the ongoing and time-transgressive dissolution of shallow halokinetic salt tiers. The salt tongues periodically shed mass flow deposits, triggered by seafloor instability created by the interactions of salt flow, salt withdrawal and the dynamic nature of salt and fault welds. In my opinion, the lack of equivalent breccias, DHAL laminites and halo evidence in otherwise similar deepwater sediment in Barney Creek Formation in the Glyde River Basin, some 80 km to the south-east of HYC, is why this basin lacks economic levels of base metal mineralisation (Figure 7).
Assuming that the first and second generation breccias in Type 1 and III breccias in all of the stratigraphically discordant deposits (allochthon and weld breccia), first defined by Walker et al., 1977 (Table 2) had shared salty origins, the wider distribution of the first generation breccias suggests that they formed via seafloor reworking processes acting across the whole region as a rim to discordant mineralisation (Williams 1978). Therefore, Williams (op cit.) argued geologically reasonable causes of the brecciation in the Cooley Dolostone include; movement on the Western and Emu faults, slumping of debris off the Western Fault Block, and stratal collapse due to the dissolution of evaporite minerals. I would argue for all of the above, but add that the whole Cooley Dolostone breccia system at the time the first generation breccias formed was a massive salt-flow fault-feeder system that was salt-allochthon cored and salt-lubricated. Situated at and just below the deep seafloor, salt tongue dissolution created salt-ablation breccias, while the halokinetic-induced seafloor instability instigated periodic mass flows into a metalliferous brine lake; as occurs today in the modern Red Sea deeps, the Orca basin in the Gulf of Mexico and the various brine lakes (DHAL's) of the Mediterranean Ridges (Table 2).
Breccia textures in a halokinetic salt ablation system are always two stage (Warren, 2016); the first stage of brecciation occurs as the salt tongue is inflated and spreading over the surrounds, even as its edges dissolve into ablation breccias reworked by further salt tongue movements and accumulations of contemporary salt-carapace materials (Figure 9). This first stage is typified by mass wasting piles related to the debris rims accumulating about the salt tongue edges, as debris slides downslope across the top of a continuously resupplied salt mass. The friction along the underside of the expanding salt sheets drives overturn, contortion, and brecciation of the underlying deep seafloor bed, this ultimately creates subsalt thrust overfolds (known as gumbo zones beneath the salt allochthons of the Gulf of Mexico). The second stage of brecciation is related to the dissolution of the salt itself once the salt supply is cut off by salt withdrawal and overburden touchdown.
Because allochthons are set up in the expansion stage of salt movement across the seafloor, Stage 1 breccias tend to be more widespread at the landsurface than stage 2 breccias. Stage 2 breccias form once the mother salt supply to the salt tongue or tier is cut off, the salt tongue then dissolves and final brecciation occurs, often with significant roof collapse features in any overburden layers. Similar two-stage allochthon breccias outcrop and subcrop in salt namakier provinces across Iran (Warren 2016, Chapter 7). However, unlike Iran the HYC laminites and associated breccias accumulated in a local deeper marine anoxic sump within a dominant subaqueous normal-marine carbonate shelf setting. There are also partial analogies with salt-cored Jurassic shelf carbonates and allochthon breccias in the paleo Gulf of Mexico, or the Cretaceous mineralised and ferruginised shelf-to-slope halokinetic-cored depositional system that now outcrops in the Domes Region of North Africa (Warren, 2008; Mohr et al. 2007).
Based on the sedimentology of the HYC ore host (Figure 9), I conclude that the HYC deposit accumulated as classic DHAL deposit in a salt allochthon-floored sump. Initial ore accumulation took place as metalliferous laminites in a local salt withdrawal basin. The anoxic brine-filled DHAL sump sat atop a deflating salt allochthon sheet with one of the tiers indicted by salt dissolution breccias at the Myrtle-Mara contact.
The following observations further support this conclusion; 1) the scale and deepwater setting of the deposit, 2) the fault-bound brine-fed margin to the deposit, 3) the rapid local subsidence of the sediments in the deeper water anoxic portion that constitutes the Barney Creek Fm host (HYC Pyrite member), 4) the syndepositional nature of the inter-ore polymict mass flow breccias, 5) the presence of syndepositional barite and Mn haloes from a diagenetically imposed oxidised saline set of pore waters hosted in what were formerly normal-marine sediment pore fluids.
Salt flowing from an allochthon sheet into salt risers in the Emu-Western fault region drove fault-bound rapid subsidence that created local deeper-water anoxic brine-filled sumps in an otherwise healthy marine carbonate shelf (see Salty Matters blog, April 29, 2016, for a salt-controlled structural analogy in the Red Sea). The fault-controlled salt risers allowed brine to escape onto the seafloor at Barney Creek time and to flow across the seafloor into the large DHAL sump that is today the HYC deposit (Figure 9). With time, the salt risers evolved in salt welds and ultimately into fault welds with salt-ablation breccia textures.
The characteristic Fe-Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the fault-conduit aquifer focus to the suprasalt brine flow and the level of hypersaline brine intersections. There are also transitions into more-typical more-oxidised marine pond and pore water masses in the upper levels atop the DHAL waters and around the edge of its brine curtain.
Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree and would argue that a later DHAL mineralisation focus, during the creation of a later generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as the mother salt supply was lost (Table 2).
Williams (op. cit.) noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite, presumably brought in from an outside source, in the matrix of the second generation breccias suggest that the later breccias formed by solution collapse following the introduction of mineralizing solutions into the porous, first generation breccias. I am in complete agreement with this conclusion. In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession that we classify as the sedex brine pool stage that is forming the HYC deposit. With time and salt dissolution/source depletion, we pass to the next generation of breccias, which are linked to a fault weld, evaporite-collapse sub-economic set of MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).
In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones. Salt withdrawal from allochthon sheets emplaced below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits, as defined by thickness and mineralogical/ colour changes in the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor depression, which was the HYC DHAL sump, it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault-weld). The stratigraphic level of the withdrawal is indicated by the allochthon collapse breccia seen at the top of the Myrtle Shale.
The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride rich brines circulating in the buried sediments of the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALs, some of the same metal-bearing brines exploited the presence of fractionally dissolved interclast calcium sulphate within diapir collapse breccias. So a similar set of redox interfaces drove discordant mineralisation in second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and salt collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks breccia and peripheral Cretaceous seafloor DHAL laminites in the Bahloul Formation of Northern Africa (see Warren 2016; Chapter 15).
The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after the main episode of laminite Pb-Zn ore formation. This interpretation of both inter-ore “sedimentary” and Cooley Dolostone member breccias across the region reconciles what were seen as previously conflicting primary versus time-transgressive relationships (e.g., Williams 1978; Perkins & Bell 1988).
The characteristic Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the aquifer and the level on hypersaline brine intersections with the more typical more oxidised marine water mass and pores water at levels atop the brine lake.
Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree, and would argue that the later mineralisation focus, during the creation of the second generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as any ongoing mother salt supply was lost. Williams (op. cit.) in a discussion of the Ridge and Cooley deposits noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite in the matrix of the second generation breccias, presumably brought in via fluids with an outside source. He suggests that later breccias formed by solution collapse following the introduction of mineralising solutions into the porous, first generation breccias. I agree also with this conclusion but would also place it in the typical saline baryte ore association seen in many salt diapir provinces such as the Walton-Magnet Cove region of Nova Scotia, or the Oraparinna Diapir in the Flinders Ranges, South Australia (see Warren 2016, Chapter 7 for detail on theses and other similar baryte deposits).
In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession we classify as the sedex brine pool that is the HYC deposit, to the next generation of breccias that are linked to a fault weld, evaporite-collapse sub-economic set of smaller scale MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).
In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones were deposited. Salt withdrawal below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits defined by the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor that was the HYC DHAL sump it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault- weld) with the stratigraphic level of the withdrawal indicated by the allochthon collapse breccia at the top of the Myrtle Shale.
The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride-rich brines circulating in the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALS, some of the same metal-bearing brines exploited diapir collapse breccias and drove discordant mineralisation and second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks and Cretaceous seafloor laminites of the Bahloul Formation of Northern Africa (see Warren 2016 Chapter 15).
The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after ore formation.
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