Salty Matters

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

Salt Dissolution (2 of 5); Salt caves

John Warren - Saturday, September 30, 2017

 

Properties of evaporite karst

At first sight, medium-scale evaporite karst surface landforms, such as dolines, polje-like depressions, subsidence bowls and collapse dolines, appear near-identical to those found in and on carbonate karst (Part 1). Likewise, all the smaller-scale intracavern features, both dissolutional (pipes and cavities) and constructive calcite speleothems), can be preserved as karst zones in and above an evaporite host, even when the formative hydrology becomes inactive and the system is buried (palaeokarst). But, the much higher rates of dissolution of halite and gypsum compared to limestone or dolomite means there are significant differences in rates of formation and cave geometries compared to carbonate caves in nonevaporitic hosts. One cannot simply take models of caves developed by studies of carbonate karst and use them to unreservedly interpret evaporite caves.

Differences reflect the inherently higher solubility and flowability of the host salts at earth surface conditions and mean most evaporite-hosted caves show peculiarities related to higher dissolution rates compared to carbonate karst. For example, one most significant in a vadose halite cave is that evaporation of a near-saturated brine is aided by airflow and means many halite stalactites tend to curve into the direction of air flow (anemolites) rather than the subvertical dripstone mechanism that controls the shape of most calcite stalactites and stalagmites. Forti (2017) is an excellent summary of evaporite speleothems and formative mechanisms in both gypsum and halite caves.

The high solubility of subsurface evaporite beds atop deeply-circulating pressurised and jointed aquifers (confined hydrologies) means deep-phreatic chimneying (vertical shafts) and deep dissolution with upward-stoping are a common deep cavity-forming mechanism in and atop buried salt beds. This, in turn, means high volumes of overburden sediment can be swallowed quickly, especially in suprasalt regions with strong roof spans covered by loose sediment (Figure 1). In such settings, large at-surface natural sinkholes and circular structures (breccia chimneys) can daylight in days or even hours and so constitute significant geohazards (see Parts 4 and 5). This type of rapid point-sourced vertical stoping atop evaporites doesn’t just happen in continental settings. Similar circular structures atop breccia-filled vertical shafts or fault-bound chimneys form beneath deep waters of the Eastern Mediterranean via substratal breaching of the Messinian evaporite, followed by upward stoping through the overlying sediments (Bertoni and Cartwright, 2005).


Worldwide, much of the dissolution action in uplifting salt beds and masses begins well below the watertable. It takes place hundreds of metres or more below the surface at bathyphreatic depths where; 1) halite is leaching, 2) anhydrite is reconverting to gypsum and 3) where caprocks and other dissolution residues are forming atop diapiric salt features. Early dissolution is greatest at contacts between the salt edges and joints or fractures in adjacent aquifers. Thus patterns of jointing or fracturing control the extent and style of caves in a salt host and typically creates high-density maze caves. With halite, the only setting where a NaCl mass makes it to the surface is as an active or recently active diapir crest or namakier. Before its dissolution, the salt matrix hosting a cave-hosting diapiric halite is largely impervious, flowing and re-annealing. Hence, most of the karst action accessible for study in halite caves is vadose and tied to perched water tables; this is most obvious near the margins of namakier salt sheets and located well away from locations of active salt fountains.

If phreatic caves ever do form near to rising salt fountain domes, these early cavities are quickly closed by the pressurised salt flow needed to bring salt to surface. Rather than being telogenetic features, many of the metre-scale blebs of finely-layered gravitationally-aligned laminar halite seen within diapirs intersected in deep salt mines, and floating in a matrix of flow foliated coarse-crystalline halite, are the result of mesogenetic salt precipitating in gas-filled open cavities within the diapir mass (see Salty Matters October 31, 2016). These laminar cavity filling growth-aligned halite textures are not telogenetic, neither are the entrained salt clasts preserving primary (relict) depositional texture of the mother salt. Instead, they indicate the existence of former or present-day gas-filled cavities (N2, CH4 or CO2), which are a hazard occasionally encountered during salt and potash mining, for example in the diapirs of NE Germany (Hedlund, 2012). Their presence and their ability to blow out mine walls show that pressurised gas pockets and associated cavern fill textures are not unusual mesogenetic features in a flowing salt mass. Away from deep mine intersections of mesogenetic gas-filled cavities, most of the shallow cavities and fill textures of a flowing salt body and studied by speleologists are telogenetic and mostly vadose.

Cave-forming processes in a telogenetic salt cave in bedded salt bodies are either vadose or phreatic (both shallow and deep), with many gypsum/anhydrite caves showing textural evidence indicating transit from one realm to the other. At any one time, the transition from phreatic to vadose landforms is tied to depth to watertable. This is seen in the gypsiferous cap karst to diapirs and allochthons of Triassic salt in the Betic Cordillera of Spain (Calaforra and Pulido-Bosch, 1999). The crests of the salt structures are dominated by collapse dolines (vadose), this passes radially out into the belt of solution dolines occupied by seasonal saline lagoons and further out into rim of saline springs that form wherever the watertable intersects the landsurface (phreatic). Some caves pass from phreatic to vadose a number of times in their histories in response to watertable fluctuations tied to varying climate and tectonics (Columbu et al., 2015.

Vadose processes characterise the uppermost part of a karst aquifer and have air in pores above a water surface or perched watertable. Water drainage is free-flowing under gravity; cave passages drain downslope and show strong gravitational orientations (numerous sub-vertical features). Because they form above the watertable, vadose cave walls are subject to surface seepage, evaporation and drying. Speleothems decorating vadose cave walls indicate varying combinations of gravity and airflow and include; stalactites, stalagmites, cave popcorn, helicites and flowstones.

Caves can form in a salt bed on its way down (eogenesis; syndepositional to early burial phreatic), on its way up (evaporite telogenesis, generally begins at bathyphreatic depths) and at the bottom of the bed’s burial history (mesogenetic; e.g., gas-driven bathykarst).


Gypsum caves

Active caves beneath gypsum karst landforms first formed as deep phreatic maze caves and are characterised by dense passage networks with numerous contemporaneously closed loops, typically in bedded evaporites. Vadose cross sections of the same system can be quite large due to the high solubility and relative homogeneity of the host, especially if formed in a thick capstone. Heimkehle cave in the Zechstein anhydrites in the Harz Mountains of Germany has an overall passage length of more than 2 km, with large rooms up to 22 metres high and 65 metres wide. It was large enough to be used in World War II to house a factory manufacturing parts for JU88 aeroplanes (Knolle et al., 2013).


Other gypsum caves, some more than 150-200 km long, are well documented in Miocene gypsum hosts in west Ukraine (Table 1; Figure 2; Klimchouk, 2000; Klimchouk and Andrejchuk, 2003; Andrejchuk and Klimchouk, 2004; Klimchouk, 2007) and the Gypsum Plain of West Texas and New Mexico (Stafford et al., 2008, 2009, 2017) and much of our current understanding of the gypsum karst process comes from these regions.

Gypsum cave walls in the vadose realm are relatively undecorated compared to carbonate caves. But calcitic speleothems can form in a gypsum cave and are well documented, as are alabastrine flowstones and selenite rinds on cave walls, for example; in northwest Texas (McGregor et al., 1963), in Alabaster Cave and others in the Blaine Gypsum in western Oklahoma (Johnson, 1996; Bozeman and Bozeman 2002), and in quarries intersecting the gypsum karst system in the Kirschberg Evaporite member near Fredericksburg, Texas (Warren et al., 1990).


Phreatic gypsum caves

Phreatic caves in bedded evaporites atop artesian aquifers typically grow as upward stoping and branching blind flow loop caverns (phreatic cupolas) and chimneys, which can begin to grow deep below the watertable as bathyphreatic or hypogenic karst. For example, deep artesian systems drive cupola karst, and blind chimney stopes in the Black Hills of Dakota, the Elk Point Basin of Canada and the Optymistychna Cave, western Ukraine (Figures 2, 4a). Deeper in a basin, where bathyphreatic caves are bathed by centripetal mesogenetic crossflows, water flow is slow and phreatic karst is influenced by the escape of H2S- and CO2-rich basinal waters, not necessarily by the confined meteoric head. Fluid flow at these greater depths is driven by pore water gradients that reflect potentiometric variations in temperature, organic maturation, pressure and salinity of basinal waters.

      

Passages in mesogenetic and telogenetic evaporite-hosted caves tend to first develop near fractures and joints in adjacent aquifers and then expand into maze cave networks. Once growing in the main evaporite mass, some phreatic maze passages can show internal upward-directed switchback gradients independent of jointing in adjacent aquifers. Cave orientations within the evaporite bed and away from the aquifer inflow level are tied to internal inhomogeneities in the host bed such as less soluble or more soluble intrasalt beds and changes in mineral proportions.

Phreatic tubes are the most common passage shape in smooth-walled blind-dissolution pockets and cupolas in bathyphreatic gypsum caves. Tube or channelway shapes range in cross section from near-circular (isotopic dissolution of a soluble host) to elliptical tubes to canyon-shaped keyholes (Figures 3, 4a, b; Klimchouk, 1992, 1996). Ornamentation is minimal where undersaturated crossflow drives the rapid dissolutional breakdown of the cave wall and the resulting passages are smooth, with local scalloped dissolution irregularities. Dense intersecting maze cave networks first form in the early stages of exhumation at the contact between a tight but soluble gypsum/anhydrite unit and a less soluble carbonate or siliciclastic aquifer (Figure 2, 3). This less soluble bed is the supply conduit for groundwater crossflows that dissolve the edges of the initially impervious anhydrite/gypsum bed. The greatest rate of water supply to the dissolving gypsum contact is along joints and fractures in the adjacent aquifer bed (especially with limestone aquifers). Accordingly, the meshwork of caves penetrates the gypsum bed and tends to follow joint and fracture patterns of the adjacent aquifer.

More stagnant phreatic portions of telogenetic CaSO4 cave systems can be saturated and so precipitate isopachous crusts and crystal rinds. Gypsum is the commonplace isopachous precipitate in voids in this setting, while anhydrite tends to dominate in cavities in the deeper mesogenetic realm (Garcia-Guinea et al., 2002). By the time phreatic maze caves are exhumed and enter a vadose (epigene) setting, where they are accessible for speleological study, much of the earlier phreatic ornamentation has already been dissolved by the increasing undersaturated throughflow and cavern interconnection. This is associated with uplift into the more hydrologically-active phreatic realm, which always precedes entry into the vadose realm. (Warren et al., 1990).

Some of the longest and most complex phreatic maze cave systems in the world are found in Miocene gypsum in west Ukraine. Optymistychna Cave, with more than 214 km of surveyed passages, is the world’s longest gypsum cave and the second or third longest cave of any type (Table 1; Figure 2; Klimchouk, 2007). The world’s longest cave, at 550 km, is the carbonate-hosted Mammoth Cave of Kentucky. The West Ukraine region contains the five longest known gypsum caves in the world, accounting for well over half of the total known length of known gypsum caves on Earth. By area and volume, the world’s largest gypsum caves are: Ozernaja (330,000 m2 and 665,000 m3; with 122 km of documented passages it is also the world’s 10th longest), Zoloushka (305,000 m2 and 712,000 m3), followed by Optimisticheskaja Cave (260,000 m2 and 520,000 m3). They are all complex joint-controlled maze caves, formed under confined aquifer conditions that existed from the Pliocene to the Early Pleistocene (that is karstification on the way up - telogenesis or gypsum exhumation, initially focused on the underside of the bed). Their growth patterns indicate upward-transverse phreatic groundwater circulation, with ultimate cavern fusion across the gypsum bed. All these west Ukrainian caves were fed by artesian crossflow in sub-gypsum and supra-gypsum aquifers that were sourced in the Carpathian Mountains.

High rates of dissolution in phreatic gypsum caves, relative to rates of water crossflow, are indicated by bevelled, faceted and “keyhole” cross sections in the Ukrainian caves along with a lack of vadose wall ornamentation (Figures 3a 4; Klimchouk, 1996; Pfeiffer and Hahn, 1976). Keyholes indicate density stratification and convectional circulation in cave-forming waters, with shapes sometimes complicated by lithological discontinuities in the gypsum bed (Figure 3b; Kempe, 1972). Convection in caves is most pronounced where sluggish artesian flow and low flow velocities dominate.

This was the case in the Pliocene to early Pleistocene history of the maze caves of the western Ukraine (Klimchouk and Andrejchuk, 2003). At that time the deeply buried gypsum dissolved via upward growing but blind, phreatic cavities, with a reflux of somewhat denser “spent” waters sinking toward the base of the cave. Spent water was replaced by less dense inflow waters supplied from the lower aquifer (Figure 3a). This set up a natural density-stratified convection, which maintained fresher (less dense) waters near the phreatic cave roof. Upward growing blind caves tend to expand more at their tops driving the transition from subcircular to keyhole caverns along the cave conduit (Figure 4b). After dissolving gypsum and increasing in density, a portion of the “spent” cave water sank all the way back into the underlying aquifer where it once again joined the regional throughflow in the lower aquifer. Once a stoping cave breached the top aquifer water flow direction in the cave was controlled by temperature, pressure and density contrasts between aquifers on either side of the gypsum bed. It seems that post-breach most of the cave water continued to rise through the cave system into the overlying aquifer (Figures 3a, 4) Similar keyhole and cupola morphologies are developed in low flow rate bathyphreatic sulphuric acid caves in carbonate hosts (e.g., early stages in the formation of the Carlsbad and Lechuguilla caverns).

Formation of keyhole passages is not an exclusively phreatic phenomenon in density-stratified gypsum caves. Keyholes in the vadose portions of many telogenetic carbonate caves indicate the transition of the cave passage from phreatic, with circular cross sections, to vadose with deepening drainage slots at the base of the passages (e.g. Calaforra and Pulido-Bosch, 2003). Phreatic stoping, followed by vadose cavern enlargement, probably explains the close correlation between caprock sinkhole distribution and position of underlying vadose passages in Permian gypsum subcrop in the Kungur Cave region in the Russian Urals, where the caprock thickness is a little as 25 m (Figure 5).


Vadose gypsum caves

A lowering of the watertable, either by uplift or climatic change tied to increasing local aridity, converts a former hypogene to meteoric phreatic cave into a vadose cave. In the latter, the walls become ornamented with gypsum and calcite speleothems. This is the recent history of the accessible portions of the gypsum maze caves worldwide, including documented examples in New Mexico, the western Ukraine and Saudi Arabia where the passage into the middle-upper Pleistocene marks the transition from phreatic to vadose in most of the accessible caves (Stafford et al., 2008). It is characteristic also of the very recent history of some natural phreatic caves where water tables were artificially lowered to allow quarrying of gypsum (Warren et al., 1990; Klimchouk, 2012).

Climate shifts or watertable fluctuations at the early end of the burial cycle (salt on the way down) can create alternating vadose-phreatic conditions in evaporite beds in the early stages of burial and so create an early watertable-associated karst level in the accumulating evaporites. That is, a gypsum bed hosting a vadose hydrology on its way down into burial, may pass through the watertable a number of times before its final burial and passage into the mesogenetic realm. This is the case today in the captured recharge playas in central Australia and about the edges of some halite-filled salars in the Andes and many Canadian salt lakes (Last, 1993; Warren, 2016), where fluctuating hydrologic conditions alternate between vadose and phreatic. Similarly, Quaternary climate shifts have variably karstified the gypsiferous sediments of many Sinic playas. For example, Yaoru and Cooper (1997) document Pleistocene lake basins in north-west China, such as in the Chaidamu Basin, where exposed gypsum beds evidence karst overprints that include: corroded flutes, fissures, small caves and associated collapse breccias and roof falls, followed by phreatic evaporite cement overprints. Watertable fluctuation is a hydrological overprint that is preserved as alternating ornamented surfaces in gypsum caves, disconformities and cave fills in many ancient lacustrine gypsum units.

Halite Caves

Because of its high solubility, halite does not make it into outcrop or shallow subcrop as easily as gypsum/anhydrite. Where halite is at the surface, it tends to be in regions of Pleistocene halite deposition (salt on its way down) or in zones of active diapirism (salt coming up very quickly). Chabert and Courbon (1997) noted caves in ancient rock salt in several regions: Algeria (in diapirs, mostly as vertical to sub-vertical shafts and short caves up to 28shafts0 m long), Chile (diapirs, with caves 250-500 m long), Israel (in Mt Sedom diapir as vadose caves and tube caves several hundred metres in length and subvertical vadose shafts that are metres across), Romania (in diapirs, with caves up to several hundred metres long), Spain (in diapirs, with caves up to 650 m long), Tajikistan (several caves 300 to 2,500 m long and up to 120 m deep), and in the namakiers of Iran and the offshore island relicts where cavern lengths range from several hundred metres to kilometres. The maintenance of landscape elevation, which can be hundreds of metres above the surrounding plains, facilitates the creation of vadose caves in the diapir crest (Table 2).


Halite in an active namakier rises and spreads rapidly, so any karst in an active salt diapir tends to be a feature associated with the immediate underside of a caprock. Karst processes cannot deeply penetrate while salt is flowing, even when the plug rises more than 300 metres above the surrounds. So, more extensive halite hosted caves are best developed about the margins of a namakier where halite’s susceptibility to rapid dissolution means the length of a cave developed below the caprock can be substantial. There is a report of a single salt halite cave (Cave 3N) on Qeshm Island, Iran, with a passage length of more than 9 km (Bruthans et al., 2002). Once the rate of diapir rise has slowed or ceased, the positive topography of the now inactive diapir controls the depth of development of doline collapse inn the diapir itself. That is, deeper collapse dolines can now form in the more central topographically higher portions of a salt structure, once the rate of rise has slowed (e.g. Calaforra and Pulido-Bosch, 1999).

The high solubility of halite means a halite cave system can form in a few hundred years rather than the thousands to tens of thousands of years needed to form carbonate karst (Bruthans et al., 2010). Not only is rate of cavern formation swift (dissolution/downcutting of 20 mm/year), these rapidly forming karst features are hosted in and atop a flowing rock mass. This means some of our notions of karst process and cave stability, related to the rate and density of cavern expansion, need to be modified when dealing with halite karst. For example, unlike gypsum and carbonate karst, jointing is not significant in controlling cavern orientation in namakier karst.

      

Dead Sea karst

Halite caves occur in the Mt. Sedom diapir, where halokinetic Miocene salt is sporadically exposed beneath a weathering and fractured gypsiferous caprock (Figure 6a,b; Frumkin, 1996; Frumkin and Ford, 1995; Frumkin 2009). Water enters the various caves in Sedom diapir through breaches in the caprock (Figure 7b). Most of the caves in the higher parts of Mt. Sedom salt are vadose inlet caves; these are meandering steeply inclined tubes and canyon slots located within or immediately below the caprock. They form where salt solution quickly carves out near vertical slots and shafts (typically < 2m broad and much deeper) that lead down from the surface, sometimes along pre-existing fractures and shears in the salt. Inlet caves in the central portions of the mountain can only be accessed through their sinks and appear to have no distinct outlet (Figure 6c). All terminate several tens of metres above the regional watertable (e.g. Karbolot Cave; Figure 6c, d; Frumkin, 1994a, b, 1996).

The lower parts of inlet caves often contain steep silt and clay banks with surge marks that indicate occurrence of low energy water ponding, with variable residence times. Silt and clay sediments settling at the bottoms of inlet caves impede infiltration, extending the residence time of pond water (Frumkin, 1994a, 1996). Three of the studied caves in northern Mount Sedom had perennial ponds throughout the period 1984-1995. The ponds are perched, without any lithologic control, tens of metres above the nearest potential outlet at the foot of the mountain (Figure 6c). The water level in each pond differs from the others by tens of metres. All pond waters are highly concentrated, up to 324 g/l, with solutes consisting mainly of sodium and chlorine. Fresh inflow waters reach halite saturation within a few hours of reaching the pond. Both dissolution and precipitation features form the pond edges, and their equivalents can be seen on cave walls wherever ponds have dried out. Dissolution is indicated by horizontal notches, which connote density stratification in the ponds when aggressive fresh flood waters are temporarily diluting the upper parts of the pond. Subsequent saturation of holomictic pond waters is indicated by the growth of cm-scale halite crystals on the bottom and sides of the ponds.


Towards the periphery of Mt Sedom, the inlet caves lead down to laterally expanding vadose cave levels that drain onto the Dead Sea plain (Figure 6c,d, 7c). Sedom Cave, is the longest laterally expanding diapir cave, with an aggregate length between two subparallel conduits of 1.8 km. Malham Cave, another large perched and laterally expanding cave, lies a few hundred metres south of Sedom Cave (Figure 6a, c, d; Frumkin, 1996). It has an aggregate passage length of more than 5.5 km and reaches to some 194 m below the landsurface. There is an upper tier of mostly inactive passages and a lower active channel level. 14C dates on fossil wood in the upper cave level shows meteoric waters began to sculpt the upper cave more than 5,500 years ago. Ongoing uplift of Mt Sedom salt means the active channel level in Malham Cave is now downcut some 10-12 metres lower than when it began. Malham Cave passages quickly developed an open outlet through which floodwater escaped directly to the Dead Sea floor, proving that during this period some 4000 years ago rock salt had already risen above region hydrological base level at the Malham outlet point within the eastern escarpment (Figure 6c). Lashelshet Cave, an inlet cave on the highest point on the diapir cross section, has an even older age of more than 7,000 years since cave initiation. Caves in the northern part of Mt Sedom did not begin to form until some 3,000 years later (Frumkin, 1996).

Based on their study of the caves of Mt Sedom, Frumkin and Ford (1995) concluded cave passages develop in two main stages: (1) an early stage characterized by inlet caves with high downcutting rates into the rock salt bed, and steep passage gradients; (2) a mature laterally expanding stage characterized by lower downcutting rates and the establishment of a wider subhorizontal perched stream bed armoured with alluvial detritus. This style of cave tends to develop toward the periphery of the diapir mound. In the mature expanding stage downcutting rates are controlled by the uplift rate of the diapir and changes of the level of the Dead Sea.

Passages may aggrade to create wide flat bevelled passages and slots with thick sediment armoured bases (Figure 7c; Frumkin, 1998). A lack of a consistent phreatic level in the blind bottoms of perched water levels and the presence of the horizontal slots in the lower levels of Sedom Cave means dissolution in both types of caves is largely restricted to times of flooding and perched or backed up freshwater in the vadose zone. This explains the tapering passages of inlet caves and the widespread alternation of armouring and bevelling as well as formation of narrow horizontal meandering slots toward parts of the top of the meandering channel that is now Sedom Cave.

Mass balance calculations in the halite caves of Mt Sedom yield downcutting rates of 0.2 mm s-1 during peak flood conditions, this is about eight orders of magnitude higher than reported rates in any limestone cave stream (Frumkin and Ford, 1995). However, floods have a low recurrence interval in the arid climate of Mount Sedom so that long-term mean downcutting rates are lower: an average rate of 8.8 mm/ year was measured for the period 1986-1991, while Frumkin (2000) estimate the average regional vadose downcutting rate in the Mt Sedom karst region to be 20 mm/year. This is still at least three orders of magnitude higher than rates established for limestone caves and more than able to cope with the rate of supply of the diapiric salt.

The highly impervious nature of halite and its resupply in actively growing diapirs means that, unlike carbonate and gypsum caves, there is no real watertable level to define maximum cave development in a rising salt stem. Rather, the inlet caves are simple dissolution tubes where rainwater has accessed halite and sank until it was saturated and then dissolution stopped until the next flood. Toward the edge of the rising stem, these inlet caves breached the edge of the salt mound and vented their perched groundwaters to the surrounding plain (Figure 6; Sedom and Malham caves). This creates a downcutting and laterally expanding cave system, which is still some metres to tens of metres above the base level of the regional watertable in the surrounding plain. The expanding cave level is dominated by mostly horizontal growth, often with a sediment-armoured floor. It has numerous benches in the walls that probably reflect changes in the hydrological base level.

Dense anastomose cave networks that characterise gypsum caves are not found in the halite caves of Mt Sedom. This reflects the ability of diapiric halite to re-anneal and the fact that all exhumed halite that makes it to the surface is diapiric, not bedded. At-surface halite is not sandwiched between jointed aquifers above and below the dissolving layer. Rather it is a growing mound subject to dissolution at its top and sides.


Away from Mt Sedom, there are active collapse sinkholes and caverns forming in the alluvial fans and clastic aprons that overlie the bedded Quaternary lacustrine halite of the Dead Sea (Figure 8a, b; salt on its way down in the burial cycle). In the sediments around the lakeshore, the pace of karst collapse has accelerated in the last 60 years due to a drastic lowering of the circum-lake watertable and the associated lakeward migration of the saline-fresh water interface (Figure 9). For example, a series of collapse dolines 2-15 m diameter and up to 7 m deep, appeared in 1990 in the New Zohar area. In January 2001 a large sinkhole, some 20 m deep and 30 m wide, cut through the asphalt surface of the main road along the western shore of the Dead Sea. It was opened by the passage of a busload of tourists on their way from Ein Gedi to the Mineral Beach solarium. Existing tourist facilities, such as the Ein Gedi beachside parking, were shut down after the road was damaged and several buildings have since collapsed into sinkholes. Sinkholes have since developed in other areas about the Dead Sea Margin including Qalia, Ein Samar and Ein Gedi. The process began in the southern part of the Dead Sea coast and slowly spread northward along the Israeli coast. Collapse is more localised in the northern and southern regions on the Jordanian side, and across the region, continues to increase in frequency as the sea level falls (Ezersky and Frumkin, 2013).

Three main types of sinkhole or doline fill have been recognized atop the dissolving Holocene salt beds of Mt Sedom; 1) Gravel holes in alluvial fans, 2) mud holes in the intervening bays of laminated clay deposits between fans, and 3) a combination of both types at the front of young alluvial fans where they overlap mud flats. Fossil, relict sinkholes have been observed in the wadi channels cutting into some old alluvial fans, showing this is a natural and ongoing process. While lake levels continue to fall (Figure 9b,c), the potential for subsidence hazards related to karst collapse is ongoing.

Sinkholes and related subsidence have been the focus of much geological study of the halite caves, but Closson et al. (2010) pointed out that an even more significant and ultimately damaging environmental effect of the ongoing water level lowering is the hectometre and larger scale landslides along the retreating shorezone. In the 1990s, international builders created major tourist resorts and industrial plants along the Jordanian and Israeli shore while, during the same period, geological hazards triggered by the level lowering spread out. From the beginning of the year 2000, sinkholes, subsidence, landslides, and river erosion damaged infrastructures more and more frequently: dykes, bridges, roads, houses, factories, pipes, crops, etc. all suffered as a result.

There is evidence of an older set of widespread ground collapses, sinkholes and caves that are tied to an earlier substantial fall in the Dead Sea water level some 4 ka. It may even be that the events described in Genesis 14 in the Christian Bible took place at the time of a substantially lowered sea level. The described battle, which occurred prior to the fall of Sodom and Gomorrah, perhaps took place on the subaerially exposed flats of the Southern Basin of the Dead Sea. The “pits of slime” described in the fall perhaps were solution collapse sinkholes activated by the 4000 ka fall in the Dead Sea water level (Frumkin and Elitzur, 2002).


Halite karst in diapiric Hormuz salt, Middle East

Namakier outcrops in and about the Arabian Gulf range from structures actively extruding salt to those in ruins where salt has not flowed for tens of thousands of years (Figure 9). Likewise the halite caves developed in the namakiers of Iran, or their offshore island counterparts, show a broad range of ages and styles of salt cave development tied to the time since cessation of salt flow (Filippi et al., 2011; Bruthans et al., 2010; Talbot et al., 2009).

Surfaces of actively flowing namakiers on the Iranian mainland are characterised by karren flutes and pinnacles, with numerous small-medium dolines, collapse structures, swallow holes and small caves at their base. As at Mt Sedom, caves tend to be sediment-armoured meandering tube caves or subvertical canyon slots that are centred on joints in the salt beneath a thin suffusion mantle. In contrast, the halite caves in the diapiric cores of the many islands in the Arabian Gulf have a more mature bevelled meandering style with thicker sediment armouring on the cavern floor. Many of these caves breach the retreating edges of former namakiers and salt fountains.

Salt movement in the various diapiric cores of these islands is inactive, or is greatly reduced, compared to the Miocene when these structures were active namakiers. For example, Dragon Breath Cave on Hormuz Island is a linear meander tube cave fed by an ephemeral stream in a shallow valley filling with alluvium (Figure 11; Bosak et al., 1999; Filippi et al., 2011). The surrounding landscape is classic salt karst with numerous depressions, blind valleys, ponors and subrosion sinks. Together they form a highly pockmarked centripetally-ringed topography, which outlines those central parts of the island underlain by shallow subcropping Hormuz salt. The cave itself is one of a number of tube caves exiting about the edge of the zone of diapiric salt. Hosted in steeply dipping diapiric salt, it is around 100 m long with its main passage created by a minor ephemeral stream. Its near flat roof, with an average inclination of 5.4%, is a notable feature and reflects either joint-related dissolutional spalling of the roof or an earlier watertable slot related to backup of a freshened water body.

The current cave passage has cut down a metre or more into earlier cave floor sediments (sediment armour), which contain clasts up to 50 cm in diameter. The cave formed by initial ingress along a linear joint, which was then widened by salt dissolution, so allowing meandering of the stream trace within the salt. It is a cave system very similar to the mature stages of laterally expanding caves in Mt Sedom. The base level of the cave correlates with the surface of a widespread marine terrace, which is now uplifted some 20 metres above sealevel and defines much of the periphery of Hormuz Island. The raising of the marine terrace is related to the ongoing raising of the island via salt flow.


Bruthans et al. (2000, 2010) show that the style of karst landform developed in dissolving diapiric salt in the Arabian Gulf Islands reflects the thickness of the carapace that caps the dissolving salt core (Figure 10). They distinguished four classes of diapir cap, each with a particular association of superficial and underground karst forms, namely: 1) outcropping salt, 2) thin capping (0.5-2 m), 3) capping with moderate thickness (5-30 m), 4) capping with greater thickness (more than 30 m). Cap thickness controls or reflects: 1) the density of recharge points, with high densities of recharge points in the thinner caps; 2) the amount of concentrated recharge which occurs at each recharge point, with suffusion karst characterising thinner caps; 3) the rate of lowering the ground surface atop the salt, with the faster rates of lowering occurring beneath thinner caps, and 4) the amount of load transported by underground flood-streams into cave systems. The volume of sediment load tends to be locally higher and focused beneath the thicker caps, particularly when inflow streams abut the edges of a dissolving salt dome. The thickness of caps atop expanding halite caves does not appear to influence the shape or style of the cave developed within the salt mass; more important seems to be the thickness of cap in the recharge area of the cave and the type of recharge into the salt environment. That is, how much water is passing into the salt and is its flow ongoing or ephemeral?


Halite caves in the relatively mature salt stems of the various islands of the Arabian Gulf, unlike carbonate systems, can swallow and store huge volumes of clastic sediment, volumes that would clog the entrance to a carbonate system. The extreme solubility of halite enables the pace of dissolution/corrosion enlargement in a salt cave to keep pace with large amounts of sediment carried into the cave by external inflows (Figure 11b). Stream sediments arriving at the cave entrance, including boulders, move inside and are trapped within the salt itself. Sediment is not dumped outside the cave entrance, which is the typical situation in blind valley river mouths at carbonate caves (Bruthans et al., 2003). For example, coarse-grained sediment fractions are carried hundreds of metres into the cave by two large intermittent streams entering the upper part of the Ponor Cave (Hormuz Island). The clasts in the resulting intra-cave alluvial fan conglomerates range from cobble of several centimetres up to 1 m diameter boulders; only sand-sized particles make it to the lower part of the same cave.

Caves capable of storing such coarse alluvium within the cavern itself are halite-specific with no equivalent in a carbonate karst terrain. There the boulder-size fraction in the cave itself is the result of roof fall, and almost all stream-borne coarse alluvium is deposited outside the cave.

 

Evaporite Speleothems

Cave walls in zones of less intense dissolution and stream crossflow are decorated with halite, gypsum and anhydrite speleothems (Figure 12). Halite has a much higher potential to form macro- or mono-crystalline speleothems than calcite and gypsum (Forti 2017). Therefore, in most of the studied halite caves around the world, relatively large euhedral or hopper halite crystals have been observed as in the Iranian and Atacama caves (Forti, 2017; De Waele et al., 2009). The preferred location for these crystals are the pools in the cave entrances, where evaporation is sufficiently low to allow the development of euhedral crystals up to 10 cm in size (Figure 12, 13; Fillipi et al., 2011).

In the Iranian caves halite macrocrystals normally form also along streams, whereas in the Atacama Desert they are completely lacking. This is because they need time to develop and therefore the stream must remain active for at least a few days (Figure 12a; Filippi et al., 2011).

Monocrystalline stalactites are widespread in the Iranian caves; the most common of which are the skeletal forms (Figure 12a; Filippi et al., 2011). An idealised skeletal stalactite normally consists of a central rounded “stalactite” from which, at different heights, three smaller and shorter rounded branches develops, being equally spaced at an angle of 120° (Figure 12b). Moreover, each branch and the central stalactite form an angle of ~70°. These values show that the directions of the central columns and the side twigs correspond to that of the four cube diagonals. Finally, at the end of each twig, there is a small halite crystal with one of its diagonals perfectly coincident with the twig (Figure 12b).

It is therefore evident that the entire structure of these peculiar stalactites consists of a single crystal lattice, albeit with a fractal appearance, and this fact is also confirmed by the presence along the rounded column of evident crystal facets oriented in the same direction (Figure 12b). Air currents and other local perturbing factors may cause a deflection from the theoretical direction of both the main column and of the side twigs (anemolites; Figure 12a). Finally, the rounded structure of the central column and of the external twigs is normally covered by glazy halite suggesting that cycles of deposition and dissolution alternate, while no inner feeding tube is present within the central column.


By these observations, the genesis of these skeletal monocrystalline stalactites is induced as driven by solutions mainly coming from brines and sprays, that then flow via gravity and capillarity only on the external surface of the stalactites. The amount and the composition of these solutions must change in time, becoming sometimes slightly undersaturated, probably during rain falls. The location of these speleothems close to waterfalls, where sprays are easily formed support that interpretation (Figure 13; Filippi et al., 2011; Forti 2017).

Compared to the growth rates of calcite speleothems in carbonate caves the growth rates of evaporite speleothems are phenomenal. Halite stalactites, several metres long and curving into the direction of airflow, formed in the mouth of Dragon Breath cave in a few years rather than millennia needed for carbonate counterparts (Bokacs et al., 1999). In August 1997, a network of numerous halite stalagmites and stalactites blocked the entrance to Dragon Breath Cave. In March 1998 there were no remains of the speleothem meshwork, while in February 1999, the stalagmites had reappeared (Bosak et al., 1999). Similar halite structures occur in the caverns of Mt Sedom and on the wet roofs of some salt mines.

Telogenetic halite deposits forming in namakiers encompass a range of mechanisms and speleothem textures (Figure 13; Filippi et al., 2011): i) via crystallization in/on streams and pools, ii) from dripping, splashing and aerosol water, iii) from evaporation of seepage and capillary water, and iv) other types of evaporative deposits. The following examples of halite textures are distinguished in each of the above-mentioned groups: i) euhedral crystals, floating rafts (raft cones), thin brine surface crusts and films; ii) straw stalactites, macrocrystalline skeletal and hyaline deposits, aerosol deposits; iii) microcrystalline forms (crusts, stalactites and stalagmites, helictites); iv) macrocrystalline helictites, halite bottom fibres and spiders, crystals in fluvial sediments, euhedral halite crystals in rock salt, combined or transient forms and biologically induced deposits. The occurrence of particular forms depends strongly on the environment, in particular on the type of brine occurrence (pool, drip, splashing brine, microscopic capillary brine, etc.), flow rate and its variation, atmospheric humidity, evaporation rate and, in some cases, on the air flow direction. Combined or transitional secondary deposits can be observed if the conditions changed during the deposition. Euhedral halite crystals originate solely below the brine surface of supersaturated streams and lakes.

Macrocrystalline skeletal deposits occur at places with abundant irregular dripping and splashing (i.e., waterfalls, or places with strong dripping from the cave ceilings, etc.). Microcrystalline (fine-grained) deposits are generated by evaporation of capillary brine at places where brine is not present in a macroscopically visible form. Straw stalactites form at places where dripping is concentrated in small spots and is frequently sufficient to assure that the tip of the stalactite will not be overgrown by halite precipitates. If the tip is blocked by halite precipitates, the brine remaining in the straw will seep through the walls and helictites start to grow in some places.

Macrocrystalline skeletal halite deposits and straw stalactites usually grow after a major rain event when dripping is strong, while microcrystalline speleothems are formed continuously during much longer periods and ultimately (usually) overgrow the other types of speleothems during dry periods. The rate of secondary halite deposition is much faster compared to the carbonate karst. Some forms increase more than 0.5 m during the first year after a strong rain event; however, the age of speleothems is difficult to estimate, as they are often combinations of segments of various ages and growth periods alternate with long intervals of inactivity.

Anhydrite forms speleothems in preference to gypsum in those rare parts of a cave with very high salinity waters, But overall gypsum speleothems dominate Some of these gypsum speleothems can be quite large, up to a few metres long.

Unlike halite and gypsum caves, which are rich in halite and gypsum formations respectively, anhydrite caves do not host anhydrite speleothems at all. This is a direct consequence of the CaSO4 – CaSO4.2H2O solubility disequilibrium, which makes the hydrated mineral (gypsum) less soluble than the anhydrous one (anhydrite) at normal cave temperatures, thus totally hindering the development of secondary anhydrite formations. Most of the gypsum produced by hydration replaces anhydrite within the rock structure, and therefore anhydrite does not form any speleothems. Nonetheless, a minor part of this secondary gypsum may develop some small deposits. In the caves of the Upper Secchia Valley, small gypsum crusts and flowstones were observed where condensation water, after dissolving anhydrite, flows over the gallery roof or walls where air currents induce evaporation (Chiesi & Forti, 1988). In the same caves, when per ascensum capillary flow and evaporation are possible, euhedral aggregates of small gypsum crystals may develop on top of rock. The interested reader is referred to a comprehensive paper by Forti, 2017 dealing the great variety of halite and gypsum speleothems.

Anhydrite karst is known from several countries of the world, but in almost all cases it is located at depths that make direct exploration almost impossible (Ford & Williams, 2007). This is the reason why until present, only caves from two locations (South Harz in Germany and Upper Secchia Valley in Italy) were explored and their speleothems studied.

The Upper Secchia Valley anhydrite caves and their chemical deposits were already known when the first monograph on speleothems in gypsum caves was published (Forti, 1996). However, at that time these caves were incorrectly considered as formed in gypsum and therefore their deposits were described along with those hosted in classical gypsum karst.

The genesis and evolution of the German and Italian anhydrite caves are completely different; in fact, the first are hypogenic caves (Kempe, 2014) and lack any natural entrance, whereas the second ones are epigenic and often develop very close to the surface (Malavolti, 1949). Therefore, chemical deposits are different in the two locations and restricted to the peculiar environment that controlled the evolution of the caves.

Leaving aside the widespread secondary gypsum produced by the hydration of the host rock, anhydrite caves are extremely poor in chemical deposits. The lack of minerals in the hypogenic caves is because they were filled with near-stagnant water for most of the time during their development. In the epigenic caves, instead, the absence of cave minerals is mainly attributed to the strong increase in volume caused by hydration of anhydrite (that turns into gypsum), which makes the wall and the ceiling of these cavities extremely fractured. In this latter setting, the rather continuous breakdown normally inhibits the development of even small chemical deposits, which, in any case, are easily washed away by the frequent floods that characterise the Upper Secchia Valley. Despite all these restrictions, the anhydrite caves proved to be interesting not only from a mineralogical point of view, as they host one cave mineral (clinochlore, Chiesi & Forti, 1985) restricted to this environment, but also for the presence of a unique gypsum/anhydrite speleothem, i.e., the huge “leather like sheets” of Barbarossa Cave (Figure 14).

      

Implications

Part 1 and Part 2 of this set of articles dealing with evaporite dissolution emphasise the importance of rapid rates of volume loss in creating a unique set of karst landforms and speleothems. This rapidity creates cavities in a hydrological milieu of contrasting brine salinity and temperature interfaces and permeability contrasts.This inherent association of voids in a setting with abrupt chemical interfaces facilitates the enrichment levels of economic commodities (part 6) and drives rapid bed stoping and foundering that forms zones of significant natural and anthropogenically enhanced geohazards in the landscape (parts 4 and 5).

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Dissolving salt (1 of 5): Landforms

John Warren - Thursday, August 31, 2017

 

Introduction

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

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


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

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

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

Evaporite solution dolines

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

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

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


Suffosion dolines

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


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


Subsidence Dolines

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


Collapse dolines

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


Sinkhole Classification

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


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

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

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


Broader-scale evaporite landforms

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

Subsidence basins

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

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

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

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

Salt Breccia Chimneys

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

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

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

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


Delaware Basin, West Texas

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

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


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

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

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


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

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


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


Western Canadian karst

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

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

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

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

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

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

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


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

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

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


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

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

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

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

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


Holbrook Anticline, Arizona

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

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

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

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

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

Implications

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

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

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

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Lithium in saline settings

John Warren - Sunday, July 30, 2017

 

Introduction

Historically, until the electronic revolution, lithium located near the top of the periodic table, was of little economic interest. With an atomic number of 3 and an atomic weight of 6.9, lithium is exceptionally small and light, with a high charge/radius ratio. Estimates of the average crustal abundance of lithium vary, but it is likely to be approximately 17–20 parts per million (ppm). In igneous rocks, the abundance is typically 28–30 ppm, but in sedimentary rocks, it can be as high as 53–60 ppm (Evans, 2014; Kunasz, 2006).Lithium-prone hypersaline brines can attain values as high as 6000 ppm, but such high levels are unusual.

Lithium compounds are consumed in the production of ceramics, lubricants, glass, and primary aluminium (Figure 1). Its high specific heat capacity makes lithium ideal in heat transfer technology where it is used in welding and metallurgical applications. Its light weight and its high electrochemical potential (it is most electronegative metal known) and its high electrical conductivity make it amenable to battery applications requiring low weight and high storage potential.

Compared to nickel metal hydride batteries, the type of battery currently powering most hybrid electric vehicles, Li-ion batteries are lighter, less bulky, and more energy efficient. Lithium batteries have three times the energy of nickel hydride at one-third of the weight, and they operate at very low temperatures with a longer battery life. The use of Li-ion batteries in electric cars and electronic devices has increased the global demand for lithium, a trend that is likely to continue. Currently, lithium‐rich saline brines are the most economically recoverable Li source on the planet. (Kesler et al., 2012; Grosjean et al., 2012).


Economically recoverable forms of lithium occur in four types of deposits (Table 1; Figure 1)

(1) Pegmatites,

(2) Continental saline brines

(3) Hydrothermally altered saline lacustrine clays,

(4) Geothermal and basinal brines.

Figure 2 plots know occurrences of lithium in saline deposits. There are four main clusters of hypersaline brine-related lithium occurrences, 1) Andean Altiplano, 2) Tibetan Highlands, 3) Qaidam Basin and 4) Playa brines in the Basin and Range of the south-west USA. Then there are the lesser volumes, as yet un-economic lithium accumulations, associated with lacustrine clays near Hector, California and in the Jadar Valley, Serbia. Basinal (oilfield) brines are known to entrain elevated levels of lithium in the Smackover Fm, USA and the Fox Creek region of Canada. Geothermal brines below the Salton Sea can also contain elevated levels of lithium.


Lithium‐rich continental brine sources account for about three‐fourths of the current world lithium production with the remainder from pegmatites (Figure 3; U.S. Geological Survey, 2017). Only the geology of hypersaline brine sources and associated saline sediment hosts, basin brines and clay replacements are discussed in detail in this article. For information on pegmatites and oilfield brines the interested reader is referred to Garrett (2004), Kesler et al.(2012) and Evans (2014). All three articles contain a broader discussion of occurrences of lithium raw materials and their processing.

In natural brines, most lithium salts are highly soluble and tend to stay in solution until lithium concentrations approach and exceed 6000 ppm. Lithium can be to be absorbed by saline playa clays at lower concentrations, as in the hectorite beds in Clayton Playa, Nevada or hydrothermally from hypersaline saline waters as formed jadarite nodules in the Jadar Valley, Serbia. Actual lithium carbonate precipitates are highly soluble and so very rare in sedimentary basins; lithium carbonate (zabuyelite) is a natural precipitate from the high-altitude hypersaline waters of Lake Zabuye on the Tibetan Plateau.


Lithium carbonate brines

Production from hypersaline pore brines in South American salars dominates current world lithium brine production, with Chile and Argentina producing some two-thirds of the documented world brine production (Figure 2). Chile has emerged as the largest lithium carbonate producer from a lake brine, largely through the exploitation of brines in Salar de Atacama. China and Argentina are the other main producers of lithium from saline lake brines (Figure 3).

Lithium in salar brines of the Andes

Salar de Atacama lies on the Tropic of Capricorn at an altitude of 2,300 m in the Desierto de Atacama, some 200 km inland from Antofagasta. In its more central portions this salt-encrusted playa contains a massive halite unit (nucleus) that is more than 900 m thick, with an area ≈ 1,100 km2. Fringing saline muds, with an area ≈ 2,000 km2, surround the nucleus (Figure 4a, b).

The current salt crust atop this halite nucleus contains a sodium chloride interstitial brine that is rich in Mg, K, Li, and B ((Figure 4c; Figure 5; Alonso and Risacher, 1996). Lithium contents of the pore brine range from 200-300 ppm in the marginal zone, some 500-1,600 ppm in the intermediate zone and 1,510-6,400 ppm in the salt nucleus). The nucleus zone averages 4,000 ppm lithium and is asymmetric with respect to the salar centre due to the sump offset via ongoing faulting. Main inflows to the salar drain volcanic formations of the Andean Highlands located to the east of the basin.


Salts dissolved in inflow waters have a double origin. Weathering of volcanic rocks supplies K, Li, Mg, B and, to a lesser extent, Na and Ca. Leaching of ancient halokinetic evaporites sourced in a mother salt layer beneath and piercing the volcanic formations provides additional amounts of Na, Ca, Cl, SO4 to the most saline inflow waters. The mass-balance of the upper halite nucleus in the salar shows a strong excess of NaCl with respect to the bittern solutes Mg, K, Li, B. According to Alonso and Risacher (1996), this suggests that the nucleus did not originate from evaporation of inflow waters similar to the present groundwaters. Rather, the excess of NaCl is due to NaCl-rich inflow waters that formerly drained the Cordillera de la Sal, a Tertiary-age evaporitic ridge along the western rim of the present-day salar (Figure 4b).


Although annual salt accretion rates in a salar salt nucleus facies in an Andean salar can be as high as 5-6 cm/year (Ruch et al., 2012), the average sedimentation rate of halite in the Atacama lake centre is ≈ 0.1 mm/year, based on the age of an ignimbrite interbedded with the salt. This slow aggradation rate implies a climatic setting of long dry periods and inactivity alternating with short wet periods during which large amounts of water, and so large amounts of salt are first recycled and then accumulate in the halite crusts of the basin sump. The lack of peripheral lacustrine deposits and the high purity of the Atacama salt also suggest that the main salt unit is not the remnant of an ancient deep saline lake, but originated mostly from evaporation of waters supplied by subsurface and subterranean saline seeps.


Once Li-rich lake brines are pumped to the surfac,e they flow into a series of evaporation pans where three main economic products (halite, potash and lithium salts) are recovered. To achieve this controlled salt-series production, the brines are first pumped from 30 metre deep boreholes that penetrate the porous salt nucleus layer into a series of solar evaporation ponds (Figure 6a). Over the successive passage through the concentrator pons, liquors are concentrated by a factor of 25, generating a final brine strength of 4.3% Li (Figure 6b). During evaporation and processing for production of halite, potassium chloride and potassium sulphate from lithium precursor brines, the ion ratios are continuously monitored and adjusted to avoid the precipitation of a lithium potassium sulphate salt. This combination of solar concentration and brine processing, proceeds as follows:

1) Sodium chloride (common salt) precipitates first. If required, this salt can be scrape-harvested as a by-product.

2) At the appropriate level of concentration, the brine is transferred to a second set of ponds in which a mixture of sodium chloride (salt) and potassium chloride (potash, in the form of sylvinite) is precipitated. These salts usually are harvested and the two components separated in a flotation plant.

3) The remaining brine is piped to another set of evaporation ponds where it remains until the concentration increases to 6000 ppm Li (essentially the saturation point of lithium chloride - saturated brines typically show a green colour as visible in Figure 6b). Ripe brine is then transferred to a recovery plant where impurities such as magnesium and boron are removed. When soda ash (sodium carbonate) is added to the ripe brines, lithium carbonate drops out. Brine with low magnesium levels is the preferred feed brine as this makes for simpler processing.

The high initial lithium content of the Atacama brines and the extremely arid setting (3200 mm pan evaporation and <15 mm precipitation) means that only 90 hectares of evaporation ponds are required in one of the current brine operations on the salar, this is only 5% of the area required at Clayton Valley, Nevada with its milder climate and lower Li concentration in the feeder brine (Figure 7). Borate (as perborate) is recovered at levels of 0.84 g/l during lithium extraction at Atacama. Increasing volumes of lithium are also produced by new salar brine processing facilities in nearby Salar de Hombre Muerto, Argentina and Salar de Uyuni, Bolivia. All these salars have lower levels of Li in the primary brine feed than Atacama.

Lithium brines in the USA

Clayton Valley is host to the only commercially producing lithium project in North America, Albemarle’s Silver Peak brine evaporation pond project (Figure 7). Historically, the Clayton Valley playa produced about one-third of the US lithium requirements, but its economic viability suffered from fierce market competition, especially from South America, and a largely depleted brine supply. Originally, the central valley area contained 100–800 ppm Li, and the discovery well at 229 m depth contained 678 ppm when pumped at 450 gpm (Garrett, 2004). The average brine analysis when commercial production of lithium carbonate began in 1966 was about 400 ppm (Figure 7). Since that time the feed concentration of lithium has been slowly declining, and in 1998 the concentration was about 100–300 ppm Li (averaging 160 ppm, Harben and Edwards, 1998).


The Silver Peak Playa has an area of 50 km2 and an elevation of 1300- 1400 m (Figure 7). It lies in the rain-shadow of the Sierra Nevada, with an annual rainfall ≈130 mm and an evaporation rate of ≈1380 mm. Near-surface sediments consist of a mixture of clays (smectite, illite, chlorite, kaolin) and salts (halite and gypsum) and widespread pedogenic calcite. Lithium in the brines is derived from weathering and leaching of volcaniclastics in the Tertiary Esmeralda Formation and Quaternary ash-fall tuffs (Davis et al., 1986). Lithium content is highest on the eastern side of the playa adjacent to the outcropping marls of the Esmeralda Fm. Before it is leached, lithium is held in the clay fraction of the playa sediments and is probably part of the clay structure (hectorite is a widespread but minor component in the Clayton Valley clays - see later)).

Lithium-rich brine feed to the plant averages 0.023% (230 ppm) lithium in a background NaCl concentration of 200,000 ppm, is pumped from depths of 100-300m in the Clayton Valley (Silver Peak) playa via a number of gravel-packed wells. The lithium (and potassium) in the deposit probably originated from hot springs along the Silver Peak Fault, with the current brine composition being a blend of evaporated water from these springs and surface and ground water that drains into the basin (Garrett, 2004). Modern saline spring outflows contain 9280–10,000 ppm Na, 786–826 ppm K and 24–43 ppm Li. Unusually high brine temperatures in some areas of the deposit (up to 44°C at fairly shallow depths ≈ 25m) tend to support a volcanic/geothermal origin for the lithium. Some of the brine feeder wells show elevated levels of radon gas.

Pumped brine progresses through a series of fractionating evaporation ponds (Figure 7; Zampirro, 2004). Lithium concentration in the liquor increases to 6,000 ppm over the course of 12 to 18 months in the solar evaporation pans. When the lithium chloride level reaches optimum concentration, the liquor is pumped to a recovery plant and treated with soda ash to precipitate lithium carbonate, which is then removed by filtration, dried, and shipped.

Lithium from brine, when the Clayton Valley first produced product in the 1970s, was considered a unique deposit. Its operations established the technology and economic viability of lithium recovery from saline brine, which led to the development of brine production from the salars of South America that now dominate world production of lithium from brine.

Lithium brine in Chinese salt lakes: Zabuye (Zhabei) and Qaidam basins

The lithium brine resource of China is mostly stored in two saline lake regions in high altitude zones, Lake Zabuye region in the Alpine tundra climate zone on the Tibetan Plateau and four salt lakes in the cold arid steppe climate region of the Qaidam Basin on the Mongolian Plateau (Figure 2). Something like 80% of brine lithium resource found in China is contained in the four salt lakes of the Qiadam: Bieletan, DongTaijinaier, XiTaijinaier, and Yiliping (Figure 8; Yu et al., 2013). Zabuye lake on the Tibetan Plateau is probably the most geologically interesting as the Li content of the lake waters are so elevated that it is the only known lacustrine location where lithium carbonate, zabuyelite, is a natural brine precipitate (Figure 9. Nie et al., 2009; Gao et al., 2012).


Qaidam Lakes

Detailed sedimentological and hydrological work in the Qaidam by Yu et al. (2013) has shown that: (1) Some 748.8 tonnes of lithium is transported annually into the lower catchment of the four salt lakes via the Hongshui-Nalinggele river (H-N river in Figure 8), which is the largest river draining into the Qaidam Basin, (2) Li-rich brines are formed only in those salt lakes in the Qaidam that are associated with inflowing rivers with Li concentrations greater than 0.4 mg/l, and (3) the water's Li concentration is positively correlated with elevated levels in both the inflowing river and the associated subsurface brine. Their findings show that long-term input of Li+ from the Hongshui-Nalinggele river controls the formation of lithium brine deposits. They conclude that the source of the lithium in the lake brines is ultimately from hydrothermal fields, where two active faults converge in the upper reach of the Hongshui River. These hydrothermal fields are associated with a magmatic heat source, as suggested by the high Li+ and As3+ content water in geysers in the geothermal field. Based on the assumption of a constant rate of lithium influx, they estimate that the total reserves of lithium in the Qaidam were likely formed since the postglacial period.

Field mapping and coring indicate that lithium reserves in each of the four salt lakes depend on the influx of Li+-bearing water from the H-N River. The data also suggest that during the progradation of the alluvial Fan I, the Hongshui-Nalinggele drained mostly into the Bieletan salt lake, until the Taijinaier River shifted its watercourse to the north and began to feed the salt lakes of the DongTaijinaier, XiTaijinaier and Yiliping salt lakes, while also driving Fan II progradation (Figure 8).

One of the You et al. (2013) major findings in terms of lithium enrichment models is the importance of the contrasting hydroclimatic conditions between the high mountains containing ice caps and the terminal salt lakes. The greater than 4000 m of relief in the watershed enables a massive amount of ions, such as K+, to be weathered and transported, together with detrital material from the extensive, relatively wet alpine regions to the concentration sumps in hyperarid terminal salt lakes, where intense evaporation rapidly enriches the lake water, resulting in evaporite deposition and associated K- and Li-rich brines. It is no surprise that a saline lake at the foot of the nearby Golmud River fan is one of the few places in the modern world where carnallitite is found (Casas and Lowenstein, 1992).


Lake Zabuye

Lake Zabuye is located some 1000 km west of Lhasa, the Tibetan capital, and lies in the ET Köppen high altitude climate zone of the Tibetan Plateau (Figures 2, 9). The lake is perennial, and water levels can vary by metres each year; in 2008 the water level was some 4422 m above sea level. At this level, the lake’s area is approximately 247 km2. Salinity varies from 360 to 440 ‰, depending on seasonal differences in water input and evaporation rate. The volume of lithium product at the lake is currently limited by the sulphate-rich nature of the primary lake brine, prior to concentration in solar pans (Gao et al., 2012).

When concentrated, the crystallisation sequence of salts from highly concentrated Zabuye lake brine at 25°C is (Figure 10a; Nie et al., 2009):

halite (NaCl) --> aphthitalite (3K2SO4•Na2SO4) --> zabuyelite (Li2CO3) --> sylvite (KCl) --> trona (Na2CO3•NaHCO3•2H2O) and thermonatrite (Na2CO3•H2O)

The lake’s brine is naturally supersaturated with NaCl and other salts, so millions of metric tons of halite, potash, trona, and other minerals have accumulated on the bottom of the lake in the past few thousand years (Zheng and Liu, 2010). Lithium carbonate and sylvite precipitate, via a combination of brine concentration and cooling, and higher levels of lithium carbonate precipitation in the end brine can be induced by the addition of soda ash, as is done in the South American salars (Figure 10b, c).


The problem with the natural lake chemistry of the Zabuye salt lake is that a lithium sulphate salt Li2SO4.3Na2SO4. 12H2O precipitates naturally in the early stages of the low-temperature evaporation process, so reducing the levels of lithium carbonate in the end-stage brines. If the brine concentration series in the pans can be artificially held at mirabilite concentration, then the amount of lithium lost to the sulphate salt is reduced, so levels of lithium in the end-stage brines improve (Gao et al., 2012).

Zabuye Lake is of significant economic value as it is a new type of exploited saline lacustrine deposit (compared to the salars of South America) in that contains it precipitates lithium and borate salts in addition to significant volumes of potash, halite, natron and Glauber’s salt. Lake waters also retain elevated levels of caesium, rubidium and bromine.

Lithium in minerals soaked in saline brines

Two saline minerals in sedimentary basins known to have significant lithium contents are hectorite and jadarite. Hectorite [Na0.33(Mg,Li)3Si4O10(F,OH)2] is a clay mineral of the smectite group, where the replacement of aluminium by lithium and magnesium is essentially complete. It has a lithium content of more than 1%, a hardness of 1–2 on Mohs scale, and a density of 2–3 kg/m3. To date, an economically viable technology for extracting lithium from hectorite, rather than from brines that enclose some of these clay deposits, has yet to be developed (Evans, 2014). Jadarite [LiNaB3SiO7(OH)], is a newly recognised mineral with up to 5.7% Li and 14.7% B. Jadarite is a white porcellanous borosilicate mineral with a Moh hardness of 4-5, and a density of 2.45 gm/cc. It is associated with borate salts such as colemanite in the Oligocene-Pliocene lacustrine host sediments in its type area in the Jadar Valley in Serbia (Stanley 2007). Hectorite is probably associated with crossflows of moderate salinity hydrothermal waters, while jadarite requires a bath of hypersaline hydrothermal waters to form.

Hectorite

Hectorite has a soft, greasy texture, a candlewax-like appearance and feels like modelling clay when squeezed between the fingers. As a colloid, hectorite’s unique thixotropic properties for emulsion stabilising, gelling, suspending, binding, bodying and disintegrating, means it sells for more than US$2,000 a ton, generally as a lubricant to the oil and gas industry. Associated authigenic clays include stevensite and saponite, and in its type area at Hector California lies adjacent to a colemanite deposit.

Hectorite is mined periodically (not as a lithium source) in its type area, the Hector Mine, near Barstow, California. There, hectorite is the main clay mineral in a sequence of altered volcanic ash beds that are interbedded with lake sediments and travertines along an 8 km fault zone (Figure 11; Ames et al., 1958). The hectorite is thought to have formed through hydrothermal alteration of the ash by saline fluids moving up the fault zone (Sweet, 1980). Lithium-bearing volcanic rocks that probably formed in the same way have also been described from Arizona, and the Clayton and King Valleys in Nevada (Brenner-Tourtelot and Glanzman, 1978; Kesler et al., 2012). Hectorite is not considered to be a prime lithium resource in any of these occurrences. It is, however, considered of co-indicator of the former, or current, presence of Li-rich saline brines and as such is considered a pointer mineral to a possible lithium brine resource.


Hectorite is thought to be a result of the combination of three distinct geological processes: 1) the alteration of volcanic ash or glass; 2) precipitation of authigenic phases from saline lacustrine pore waters; and/or 3) the incorporation of lithium into existing smectite clay deposits (Asher-Bolinder, 1991). To form hectorite, all three processes require an arid environment and are associated with lithium-enriched saline alkaline waters, volcanic rocks and hot springs that can also co-precipitate travertines and fine-grained amorphous silica (Zientek & Orris, 2005).

The same association of processes explains the lithium-rich hectorite clays in King Valley (Nevada Lithium prospects) Nevada. There, layers of hectorite occur in a sequence of sedimentary and tuffaceous rocks in moat sediments along the western side of the McDermitt caldera (Figure 12; Kesler et al., 2012). Volcanic activity at the McDermitt caldera complex has yielded precise 40Ar/39Ar ages of 16.5 to 16.1 million years ago and was characterised by extrusion of early metaluminous and peralkaline rhyolite, followed by the eruption of a voluminous ignimbrite with peralkaline rhyolite to metaluminous dacite compositions (Carew and Rossi, 2016). After collapse, the central part of the caldera complex was the site of resurgence, and a moat-like lake formed between this resurgent dome and the caldera walls. The lake was the site of deposition of volcaniclastic sediments that now form a nearly continuous ring within the caldera and host the various hectorite lenses(Figure 11).


Hectorite layers ranging from 1 to 90 m in thickness and have been recognised over a length of about 20 km. Individual layers or groups of layers extending for several km and are annotated as stage 1-5 lenses. The Stage 1 lens of the Lithium Nevada deposit (informally known as the King Valley deposit) has proven and probable reserves of 50 million tonnes, averaging 0.312% Li (Carew and Rossi, 2016). As in the type area in the Hector Mine in California, hectorite in the various lenses is the main lithium-bearing clay mineral in a sequence of altered volcanic ash beds. These ash beds are interbedded with saline lake sediments and travertines, and are hosted in the sedimentary moat facies adjacent to an 8 km fault zone. That is the hectorite formed through hydrothermal alteration of volcaniclastic ash in regions where moderately saline hydrothermal fluids moved up a fault zone.

Hectorite clays are also found in the Sonora Lithium Project, 11 km south of Bacadehuachi in the Sonora state of north-west Mexico. The resource statement, in an April 2016 report, lists 839,000 tonnes of contained lithium in the indicated category and a further 515,000 tonnes in the inferred category, within two distinct lacustrine clay units situated below basaltic caprocks (Pittuck and Lepley, 2016). A pre-feasibility study has been completed, which proposes a two-phase open-pit mine with lithium carbonate processing facility and a mine life of 20 years. A pilot plant has also been constructed, and discussions have commenced regarding possible off-take agreements.

None of these hectorite occurrences are currently mined as a lithium resource.

Jadarite

Jadarite was discovered in 2007 by Rio Tinto and the Jadar deposit, near the town of Loznica, and at that time was estimated to contain an inferred resource of 125.3 million tonnes at a weighted average of 1.8% Li2O, in addition to an inferred resource of boron minerals. Jadarite has so far only been identified in significant amounts within the 20-km long Jadar Basin of Serbia. The Jadar Basin entrains oil shales, dolomicrites, pyroclastic sediments and evaporites which are believed to have accumulated in an intermontane lacustrine environment.


The jadarite occurs both in massive form, several metres thick, and also as small nodules within a fine-grained carbonate matrix (Stanley, et al., 2007). At the main Jadar deposit, a layer containing nodular colemanite (Ca2B6O115H2O) overlies three separate layers or lenses containing jadarite LiNaB3SiO7(OH). Jadarite likely formed via a hydrothermally-facilitated interaction between saline brine and clastic/evaporitic sediment, either in a tuffaceous or clay host (Kesler et al., 2012).

In May 2017, Rio Tinto announced that the Jadar area contains one of the largest lithium deposits in the world, lifting their estimate for Lower Jadar's deposits to 138 million tonnes. Extraction is scheduled to begin in 2023, with a projected underground exploitability of 50 years. As of June 2017, construction of a mine has not begun. A jadarite processing plant is also planned next to the mines, that plant will process the ore into lithium carbonate and boric acid.


Summary

Characteristics that appear to be essential to define a potential lithium carbonate brine resource are; i) an arid climate and, ii) a closed, tectonically active basin, with significant elevation and tectonically activity, which can entrain brines with elevated lithium contents (Figure 15; Bradley et al., 2013; Yu et al., 2013; Warren, 2010, 2016). Sources of lithium can be deeply circulated magmatic or recycled basinal fluids. Magnesium levels in the brine should not be too high as this complicates brine processing during lithium carbonate extraction. A co-occurring potash resource, extractable from the same brine, if present can reduce processing costs.

Another possible requirement—or at least a favourable characteristic—is elevated heat flow, as evident from young volcanoes or hot springs and the associated increase in Li-rich juvenile waters flushing the surrounding drainage basin, as is occurring beneath the Andean Altiplano. Volcanogenic source rocks in the lake drainage, such as felsic, vitric tuffs that have abundant and readily leached lithium are favourable, but perhaps not essential, since lithium is present in most crustal rocks at tens of parts per million. Worldwide all the exploited salt lakes have lithium levels in their lake brines that are well above typical (Figure 14).

Another possible favourable indication of a lithium brine is the existence of hectorite or jadarite in associated clays in the bajada rims.

All known and potential lithium brine deposits are located in arid tectonically active areas, typically in subduction or collision belts with deep-faulted suture systems (Figure 2). At the world scale, lithium-prone saline deposits are latitudinally restricted to cool arid Koeppen climate belts within endorheic brine sumps surrounded by high altitude drainage basins (Figures 2, 15b; Warren, 2010). Borates as evaporite salts are generally tied to the same setting (Warren, 2010).


Active faulting appears to be involved in forming a suitable spring-fed hydrology for all known economic lithium brine basins. Fault-related subsidence also creates accommodation space, without which only a thin veneer of arid basin sediments and brines can accumulate. Thus, a thick basin fill is needed to provide an aquifer of sufficient volume to hold a viable lithium brine resource.

In contrast, saline lakes atop shallow, superficial basins in intracratonic regions such as the Sahara Desert and most inland Australian deserts largely lack active fault control and associated rapid subsidence, and are not known to be prospective for lithium brines.

Salt fills in some lithium basin lacustrine sumps are cut by active intrabasinal faults (known from boreholes and seismic) but have no surface expression due to rapid infill and levelling of the accommodation space by salt precipitation. Significantly, the brine pools in Clayton Valley, Salar de Atacama and the Qaidam sumps are localized along active intrabasinal faults, which also control the distribution of aquifers and influence groundwater movement patterns, as well as the position of maximum stacking of concentrates and brines in the halite nucleus, along with porosity retention levels in the subsurface halite host (Zampirro, 2004; Jordan and others, 2002).

Porosity levels in a host halite aquifer are a major constraint on the potential economics of any salar or salt lake lithium brine resource. Most halite units lose their effective porosity and permeability by depths of 50-60 metres (Warren, 2016; Chapter 1). Thus, most Quaternary lithium brine operations hosted in a halite bed/aquifer will have an economic basement to brine recovery at around this depth. It is unlikely that recovery operations in Salar de Atacama and planned projects in Salar de Uyuni can recover economic brine volumes at much greater depths. There may be a 900m thick halite-dominant succession infilling Salar de Atacama and a number of halite beds to a depth of 120m in Salar de Uyuni, but economic porosities in the halite will likely only be present in the upper portions of the halite fill in both salars.

Postulating likely lithium resources, in a salar of a salt lake, to depths greater than 50-60m should only be done after salar-hydrology-aware drilling has established the presence of economic permeabilities in the hosting halite aquifer and this is likely related to the presence of active faults-. Such measurements require drilling and sampling equipment that facilitates reliable “in-situ” determinations of porosity and permeability in the halite mass and Li measurements in the brine that are related to actual content at the level of measurement, with minimal contamination by waters from outside the measured horizon.

References

Alonso, H., and F. Risacher, 1996, Geochemistry of the Salar de Atacama. 1. Origin of components and salt balance [Spanish]: Revista Geologica de Chile, v. 23, p. 113-122.

Ames, L. L., L. B. Sand, and S. S. Goldich, 1958, A contribution on the Hector California bentonite deposit: Economic Geology, v. 53, p. 22-37.

Asher-Bolinder, S., 1991, Descriptive model of lithium in smectites of closed basins, in G. J. Orris, and J. D. Bliss, eds., Some Industrial Mineral Deposit Models: Descriptive Deposit Models, United States Geological Survey Open Report 91-11A, p. 53-54.

Bradley, D., L. Munk, H. Jochens, S. Hynek, and K. Labay, 2013, A Preliminary Deposit Model for Lithium Brines: USGS Open-File Report 2013–1006.

Brenner-Tourtelot, E. F., and R. K. Glanzman, 1978, Lithium-bearing rocks of the HorsevSpring Formation, Clark County, Nevada: Energy, v. 3, p. 255-262.

Carew, T. J., and M. E. Rossi, 2016, Independent Technical Report for the Lithium Nevada project, Nevada, USA, in G. Arseneau, ed., SRK Project Number 2CL020.000, Vancouver Canada, SRK Consulting.

Carmona, V., J. J. Pueyo, C. Taberner, G. Chong, and M. Thirlwall, 2000, Solute inputs in the Salar de Atacama (N. Chile): Journal of Geochemical Exploration, v. 69, p. 449-452.

Casas, E., T. K. Lowenstein, R. J. Spencer, and P. Zhang, 1992, Carnallite mineralisation in the nonmarine, Qaidam Basin, China; evidence for the early diagenetic origin of potash evaporites: Journal of Sedimentary Petrology, v. 62, p. 881-898.

Davis, J. R., I. Friedman, and J. D. Gleason, 1986, Origin of the lithium-rich brine, Clayton Valley, Nevada: US Geological Survey Bulletin, v. 1622, p. 131-138.

Evans, R. K., 2014, Lithium (Chapter 10), in A. G. Gunn, ed., Critical Metals Handbook Chichester, UK, John Wiley & Sons Ltd.

Gao, F., M. Zheng, P. Song, L. Bu, and Y. Wang, 2012, The 273.15-K-isothermal evaporation experiment of lithium brine from the Zhabei Salt Lake, Tibet, and its geochemical significance: Aquatic Geochemistry, v. 18, p. 343-356.

Garrett, D. E., 2004, Handbook of Lithium and natural Calcium Chloride; Their deposits, processing, uses and properties Amsterdam, Elsevier Academic Press, 476 p.

Grosjean, C., P. H. Miranda, M. Perrin, and P. Poggi, 2012, Assessment of world lithium resources and consequences of their geographic distribution on the expected development of the electric vehicle industry: Renewable and Sustainable Energy Reviews, v. 16, p. 1735-1744.

Harben, P., and G. Edwards, 1998, Lithium: Times are Changing: Can. Inst. Mining and Metal., Special Volume, v. 50, p. 172-182.

Jordan, T. E., N. Muñoz, M. Hein, T. K. Lowenstein, L. Godfrey, and J. Yu, 2002, Active faulting and folding without topographic expression in an evaporite basin, Chile: Geological Society of America Bulletin, v. 114, p. 1406-1421.

Kesler, S. E., P. W. Gruber, P. A. Medina, G. A. Keoleian, M. P. Everson, and T. J. Wallington, 2012, Global lithium resources: Relative importance of pegmatite, brine and other deposits: Ore Geology Reviews, v. 48, p. 55-69.

Kunasz, I., 2006, Lithium Resources, in J. E. Kogel, N. C. Trivedi, J. M. Barker, and S. T. Krukowski, eds., Industrial Minerals and Rocks, SME (Soc. Mining Metallurgy and Exploration), p. 599-614.

Nie, Z., L. Bu, M. Zheng, and W. Huang, 2011, Experimental study of natural brine solar ponds in Tibet: Solar Energy, v. 85, p. 1537–1542.

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Stanley, C. J., 2007, Jadarite, LiNaSiB3O7(OH), a new mineral species from the Jadar Basin, Serbia: European Journal of Mineralogy, v. 19, p. 575-580.

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Aeolian Gypsum and Saline Pans - an indicator of climate change

John Warren - Friday, June 30, 2017

Introduction

Evaporites deposited as aeolian dunes are not commonplace in Quaternary successions and not yet documented in any pre-Quaternary succession (Table 1). These eolian deposits are deposited above the water table in a vadose setting, generally in a degrading playa or salt lake hydrology. Consequently, there is an inherent low preservation potential for this style of evaporite; most documented examples are less than a few tens of thousands of years old.

 

Even though relatively rare as an evaporite type, the presence of eolian evaporites, usually as gypsum dunes or lunettes with associated soils and saline mudflats, does indicate particular climatic and hydrological conditions. Eolian gypsum deposits may have possible counterparts in the Martian landscape (Szynkiewicz et el., 2010).

Over the Quaternary and across the Australian continental interior, increased aridity is expressed by episodes of dune reactivation, lake basin deflation with eroded sediment accumulating downwind in transverse dunes or lunettes (Bowler, 1973; Fitzsimmons et al., 2007), Deposition is tied to increased dust mobility (Hesse and McTainsh, 2003) and reduced river discharge and channel size (Nanson et al., 1995). Such responses to increasing landscape aridity in saline groundwater sumps are seen in most arid to semi-arid regions of the world where water tables are falling, usually driven by increasing aridity.

This article focuses on eroded subaerial evaporites as a response to increasing aridity, especially the formation of gypsum dunes and lunettes (Table 1; Figure 1).


Gypsum dune styles and saline pans

Figure 1 and Table 1 plot documented occurrences of eolian gypsum across the world, overlain on a Koeppen climate base (Figure 1a). Figure 1b plots the latitudinal occurrences of documented gypsum dunes versus elevation and Koppen climate type. Figures 1c and 1d plot the detail of these same occurrences for the USA and Australia, where individual deposits are better documented. At the worldscale, there is an obvious tie to the world's desert belts with occurrences consistently situated in regions of the cool dry descending cells of northern and southern hemisphere Hadley cells (positions indicated by light blue rectangles in Figure 1b - See also Salty Matters article from Jan. 31, 2017). Many occurrences are also situated in Late Pleistocene to Holocene climate transition zones, marked by aridification at the transition from Late Pleistocene to Holocene climates, and in many case tied to transitions from perennial saline lakes and mega-lakes to continental saltflats to dunes and interdunal pans, An example of a quartz sand erg association (downwind of a gypsiferous strandzone) is seen in the transition area into the southern Kallakoopah Pans from the northern margin of Lake Eyre, Australia and its megalake precursor (Figure 2).


At the local scale, gypsum dunes generally occur downwind or atop a saline pan or playa that is, or was, recently subject to a lowering of its lacustrine watertable. In many situations the elongation of individual pan shapes line up in an orthogonal direction to the dominant wind and so also show an eolian control, like the associated gypsum dune position and alignment (Figure 3). Wind-aligned lakes and sumps and oriented-pans are much more numerous with a broader climatic range than gypsum dunes (Goudie and Wells, 1995; Goudie et al., 2016). When present, eolian bedforms associated with oriented pans lacking evaporites are dominated by clay pellets or quartz sand.


Many of the pan edge dunes show crescent shapes and so are termed lunettes. (Figure 3; Bowler, 1973). Lunette sediments range in composition from quartz-rich to sandy clay, gypsiferous clay to nearly pure gypsum. Pure quartz dune lunettes likely formed under lake-full conditions, and so show a distinct hydrology from that of the clay pellet or gypsum-rich varieties, which form by deflation of subaerially-exposed adjacent lake floors. The flocculation of suspended clays into pellets requires some degree of salinity but is less than that required to precipitate gypsum.

Lunette sediments range in composition from quartz-rich, sandy clay, through gypseous clay to nearly pure gypsum. Pure quartz dunes formed under lake-full conditions and are distinct from that of the clay and gypsum-rich varieties, which formed by flocculation and deflation from adjacent subaerially exposed lake floors. (Bowler, 1986). Gypsum and pelleted clay dunes (lunettes) line the edges of many salt lakes and playas in southeastern, southern and southwestern Australia; Prungle Lakes and Lake Fowler (gypsum lunettes), Lake Tyrell (clay lunette with occassional gypsum enrichment) and Lake Mungo (quartz sand lunette). All these lunettes are lake or pan-edge relicts from the Late Pleistocene deflationary period, when the lacustrine hydrology changed from perennial water-filled lakes to desiccated mudflats. Likewise, there are gypsum dunes in deflationary depressions in Salt Flat Playa and the Bonneville/Great Salt Lake region of Utah (Figure 4; Table 1).


Internal sedimentary structures in many of these lake-edge gypsum dunes or lunettes show tabular cross beds with consistent bedform orientation. Many lack abundant trough or festoon cross beds, suggesting consistent wind directions (Jones 1953; Bowler, 1973, 1983). Grain constituents clearly indicate deflation of former lake sediments, which were mostly vadose prior to deflation and passage into the dunes (Figure 4).

Gypsum dunes are part of a much broader lake-edge eolian sandflat association with the lakes often supplying large volumes of quartzose eolian sediment into adjacent sand seas or ergs (Figure 2; Warren, 2016). As mentioned pan-edge dunes described as ‘lunettes’ have a characteristic crescentic shape, other lake edge dunes may show more linear or longitudinal outlines, sometimes with parts of large sand seas or ergs being fed by the deflation of the salt lake or pan as at the southern edge of the Simpson Desert in Australia where it is in contact with the expanding and contracting edge of (Lake Eyre Figure 2).

Hydrological transitions from downwind evaporite dunes and lunettes

The role of salts, groundwater oscillations and the associated lake water levels/watertables are critical in creating eolian evaporites. Typically, once seasonal drying of an increasing arid lake floor sump begins, remaining surface waters with suspended clay become saline enough for the clay to flocculate and sink to the bottom of the desiccating water mass. If surface water concentration continues and the water surface sinks into the sediments to become a saline water table, then secondary gypsum prisms and nodules grow within the capillary zone of already-deposited sediment. In waters that are increasingly saline but not saturated with gypsum or halite, pelletization can continue to occur in the capillary fringe of clayey surface sediment (Figure 5).


Ongoing seasonal aridity further lowers the watertable in a saline mudflat, so the upper part of the vadose sediment column leaves the top of the capillary zone. It then deflates, leading to an accumulation of sand-sized sediment in adjacent eolian lunettes. If there is a prevailing wind direction, this builds significant volumes of dune sediment in a particular wind-aligned quadrant of the saline pan edge. Whether clay pellets or gypsum crystals are the dominant lunette component depends on the humidity inherent to the pan climate. In hyperarid situations, halite can be an eolian component in the lake hydrology (Salar de Uyuni; Svendsen, 2003).

In some lunettes, the mineralogy changes according to climate-driven changes in the hydrogeochemistry of the lake waters sourcing the lunette. For example in the Lake Tyrell lunette in semi-arid southwest Australia, the sediments in a layer range from clay pellets (75%) and dolomite (25%) in somewhat humid times of deflation to layers, with gypsum making up >90%, indicative of a more arid hydrochemistry. Lunettes associated with the shrinkage and deflation of Late Pleistocene Estancia megalake (New Mexico, USA) show similar variations in the proportions of clay pellet and gypsum sands in lake margin deposits around the edges of up to 120 blowout depressions. These blowouts define the former extent of the shrinking megalake and encompass both shoreline and lunette sands (Allen and Anderson, 2000)

Thus, the presence of an active gypsum lunette-field at a saline pan or playa edge is tied to landscape instability and a change from more humid to more arid conditions. To form a lunette requires a change in climate and an associated change in pan or playa hydrology and it hydrological base level and lake edge water table level, over time frames typically measured in hundreds to thousands of years.

 

Not just sand and dust-sized particles

Coarser than sand-sized gypsum crystals are transported in in lake margin mounds under hyperarid windy conditions that typify ephemeral pans and saline mudflats in parts of the Andean Altiplano and even higher elevations in the alpine tundra climatic zones. Salar Gorbea is a type example for this type of coarse-grained eolian transport (Figure 6; Benison, 2017). Whirlwinds, dry convective helical vortices, can move large gypsum crystals in their passage over the saline muflat. The transported gravel-sized crystals are entrained on the saline pan surface, after they first grew subaqueously in shallow surface brine pools. Once the pools dry up the crystal clusters disaggrate and then are transported as much as 5 km to be deposited in large dune-like mounds.

The dune gravel is cemented relatively quickly by gypsum cement precipitating from near-surface saline groundwater, resulting in a gypsum breccia. This documentation marks the first occurrence of gravel-sized evaporite grains being moved efficiently in air by suspension and provides a new possible interpretation for some ancient breccias and conglomerates, and improves understanding of limits of extremity of Earth surface environments.

 

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Calcium Chloride (CaCl2), Article 2 of 2: CaCl2 minerals in evaporites

John Warren - Wednesday, May 31, 2017

 

Introduction

Found in association with some highly-saline calcium chloride brines are four naturally occurring CaCl2 minerals with evaporite associations, namely; 1) tachyhydrite CaMg2Cl6.12H2O), 2) antarcticite (CaCl2.6H2O), 3) sinjarite (CaCl2.2H2O) and 4) chlorocalcite (KCaCl3). For any CaCl2 salt to precipitate in a saline setting requires a Ca-enriched mother brine completely depleted in sulphate (Table 1). Otherwise, gypsum precipitates and in so doing removes all calcium from the concentrating brine, well before the bittern stage.

 

Of the four, tachyhydrite is perhaps the most common solid phase, sinjarite or chlorocalcite the least. Bischofite is sometimes a co-precipitate with tachyhydrite and indicative of a mother brine with an elevated MgCl2 content (Figure 1). Even so, tachyhydrite is still a rare bittern mineral phase across the Phanerozoic of Earth, although as we shall see later, the situation may be a little different on the martian surface. Along with carnallite and bischofite, tachyhydrite typifies highly saline marine bittern assemblages in only a few ancient potash-rich evaporite systems. These same systems are also MgSO4 poor, and tend to define times in Earth history of MgSO4-depleted seas. Occurrences include: Silurian potash in Michigan Basin; Carboniferous potash in the Canadian Maritimes and the Paradox Basin; Permian Zechstein-2 Stassfurt potash seams in Stassfurt, Germany; Triassic Saharan potash of North Africa; Cretaceous potash evaporites in Brazil, Gabon and Thailand and Oligocene potash in the Rhine Graben. Tachyhydrite-bearing successions are much less common within a broader MgSO4-poor group of potash deposits (see Warren 2016, chapters 2 and 11; Warren, 2017).


To form antarcticite (CaCl2).6H2O) at the bittern stage, a CaCl2 brine must be depleted in Mg, otherwise tachyhydrite forms, along with carnallite and bischofite. Carnallite, bischofite and tachyhydrite are not part of the bittern suite precipitated from today’s MgSO4-rich oceans, where sulphate and magnesium levels are too high relative to calcium. At times in the past, these three salts were more common precipitates in the bittern stages of ancient CaCl2 or MgSO4-depleted oceans. These tended to be at times when rates of seafloor spreading, and hence volumes of through-circulated seawater, were greater than today (Lowenstein et al., 2014). It seems that most the highly saline calcium chloride brines form in sedimentary basins with a thick evaporite near the base of the sediment fill (see article August 11, 2015). But, not all are basins with highly saline CaCl2 brines are co-associated with bedded tachyhydrite occurrences in the buried and dissolving evaporite units.

Antarcticite

Antarcticite (CaCl2.6H2O) is extremely hygroscopic, colourless with perfect basal cleavage and good to perfect prismatic cleavage, is brittle with a hardness between 2 and 3, and has a specific gravity of 1.715 ± 0.010 (14 °C), and a density of 1.700 g/cm3 (Dunning and Cooper, 1969). The lowest temperature for antarcticite to precipitate is -50 °C and it’s melting point is 30 °C.

There are only two known modern hypersaline saline deposits where the pure calcium chloride mineral antarcticite (CaCl2.6H2O) is present as thin crystal beds and efflorescences; they are, Bristol Dry Lake (BWh) in California and Don Juan Pond (EF) in the Wright Valley, near Lake Vanda in Antarctica (Dunning and Cooper, 1969; Horita, 2009). Both saline lakes possess similar CaCl2-enriched, MgCl2-depleted chemistries and are fed in part by hydrothermal waters. In Bristol Dry Lake the concentration process is driven by solar evaporation, while cryogenesis is the primary driver in Don Juan Pond. Only in Bristol Dry Lake is a CaCl2 brine is commercially extracted from Quaternary sediments. There is also a minor occurrence as isolated small crystals in the Kunteyi Lake of the Tibetan Plateau (Kezao and Bowler, 1986; Zheng and Liu, 2009)


Don Juan Pond, Antarctica

Don Juan Pond is a perennial free-standing water body, approximately 300 m long and 100 m wide, surrounded by glacial moraine deposits in the Wright Valley, Antarctica (Figure 2). It defines the outcrop portion of a liquid water drainage sump that retains a 10-cm-deep CaCl2 brine at the surface (Torri and Ossaka, 1965; Marion, 1997; Burt et al., 2003) (Figure 2). This pool contains some of the saltiest free-standing perennial waters on the earth’s surface, with concentrations ranging up to 40% (by mass; Dickson et al., 2013). Mean annual precipitation the Don Juan Pond region is 5 to 10 cm and air temperatures range from –55°C to +10°C (Marion, 1997). The pond does not freeze because it is a eutectic brine of H2O, CaCl2, and NaCl, with a freezing point of -52°C and a density of approximately 1.4 (Marion, 1997; Burt et al., 2002). Crystals of antarcticite were first found sticking out of the brine at the bottom of the pond, with halite and gypsum fringes at the edges of the pond (Torri and Ossaka, 1965).

Precipitation of antarcticite in this cryogenic environment is the result of the arid climate, ongoing cold temperatures and a CaCl2 brine chemistry (Torri and Ossaka, 1965). The CaCl2 brine is in part derived from springs fed by a deep, confined, basaltic aquifer, with a groundwater similar in composition to the surface brine (Harris et al., 1979; Marion, 1997; Burt et al., 2002). The origin of the CaCl2 is still contentious, but is most likely be related to mineral weathering and periodic deliquescence of calcium chloride held in the sediments(Dickson et al., 2013).


Brine in the Don Juan Pond is predominantly from CaCl2 solutions (Figure 3; Marion, 1997). Only Na of the remaining cations contributes significantly to the overall charge balance. Both Mg and K are present at very low concentrations, and SO4 is usually not detectable. In contrast to low ionic variability in seawater composition, levels of the dominant ions (Ca, Na, and Cl) in the Pond are highly variable (Figure 3). The concentration of the brine varies seasonal, controlled mainly by a complex hydrologic system including groundwater discharge, melting of nearby permafrost, rare precipitation, and evaporation or sublimation (Harris et al., 1979; Dickson et al., 2013). Despite the changes in the absolute concentration of Ca and Cl across an annual cycle, the Ca/Cl molal ratio remained relatively constant over the 23-year sampling period; it varies between 0.432 and 0.506 with an average of 0.462 ± 0.018 (Figure 2; Marion, 1997).

Bristol Dry Lake, California

Bristol Dry Lake is a 155 km2 continental playa located within a closed Basin and Range valley in the Mojave Desert of southern California (Figure 4a). It is the second known location of natural antarcticite and is thought to precipitate from lake brines via a lowering of brine temperatures (Dunning and Cooper, 1969). No antarcticite was identified in the nearby Cadiz or Danby dry lakes. Lake brines first concentrate by solar evaporation beyond halite precipitation Mountains surrounding Bristol Dry Lake are composed mostly of pre-Cambrian and Mesozoic granite, some Palaeozoic carbonates, and Tertiary and Pleistocene volcanic rocks including Quaternary olivine basalts that erupted from nearby Amboy Crater. Alluvial fans are present between the mountains and the lowest parts of the valley (Dunning and Cooper, 1969; Handford, 1982).


Bristol Dry Lake playa contains more than 300 m of interbedded halite and salt-bearing sands, silts and clay. Halite and gypsum are currently deposited from ephemeral bodies of water under evaporative conditions, with a gypsum pavement surrounding more saline halite-saturated sumps in the central lake depression (Figure 4a; Dunning and Cooper, 1969; Handford, 1982; Rosen and Warren, 1990). The primary evaporite minerals at Bristol Dry Lake are gypsum, anhydrite, and halite with rare antarcticite, celestite (SrSO4), calcite, and possibly chlorocalcite (KCaCl3). Bristol Dry Lake is mined for its chloride minerals and has been since the 1900s, and so is covered with pits and trenches (Figure 4b).

Antarcticite was identified in one of these previously excavated trenches approximately 9 kilometres south of Amboy, near the road from Amboy to Twenty-nine Palms. Dunning and Cooper (1969) examined this and other trenches during the winters of 1961, 1962, and 1968. Only one trench had antarcticite. The trench was 3 m deep, 20 m long, and contained a 10 to 25 cm deep pool of CaCl2 brine extending along the length of the trench. At the base of the pit was a massive halite deposit, with acicular groups of antarcticite protruding outward from the halite and within the brine. Antarcticite is not present on the playa surface as no surface brines are present. The only CaCl2 brines exposed to solar radiation are subsurface brines that have seeped into, and are now exposed within, the excavated trenches.


Dunning and Cooper (1969) suggested that CaCl2 brines at Bristol Dry Lake originated from chlorine of volcanic origin interacting with calcium carbonate in surrounding alluvial deposits (Figure 5). Lowenstein et al. (2009) argue that the nearby active magma chamber drives the formation of Ca-Cl brines at elevated temperatures, as well as the thermally-driven transportation of these brines to the surface. Other evidence for the magma chamber is the Amboy crater and its associated recent lava flows, which occur directly North of Bristol Lake. Evaporative concentration of this calcium chloride brine, driven by an arid climate and typical surface temperatures below 30°C, encourage antarcticite precipitation (Dunning and Cooper, 1969).

Sinjarite (CaCl2.2H2O)

Sinjarite is a modern ephemeral precipitate in surficial alluvial fan sediments near Sinjar in Iraq (Figure 6). No natural occurrences of either antarcticite or sinjarite have been documented in ancient salts, except as a volumetrically insignificant mineral phase in brine inclusions in both sediments and igneous rocks (Polozov et al., 2016), even in sedimentary basins containing highly saline CaCl2-rich pore fluids.


Sinjarite is an extremely rare pink-tetragonal-hygroscopic mineral that is highly soluble with 117 g dissolving in 100 g of water at 20 °C (Table 1; Aljubouri and Aldabbagh, 1980). The lowest temperature for sinjarite to precipitate is 42.3 °C, while its melting point is 175.5 °C (Wardlaw, 1972). Theoretical composition of sinjarite is 27.26% Ca, 48.24% Cl, and 24.50% H2O but the actual chemistry for sinjarite present at its type locality in Iraq is 25.84% Ca, 46.64% Cl, 26.55% H2O, and 0.85% Na with 226 ppm K and trace amounts of Mg, Sr, and Fe (Aljubouri and Aldabbagh, 1980). Properties include; good cleavage, average density of 1.81 g/cm3, a very deliquescent habit, hardness of 1.5, vitreous to resinous lustre, white streak, and occurs naturally in massive to granular forms.

Little research has been performed on sinjarite and knowledge of the environment of formation comes from the only place it was discovered. Sinjarite occurs in detritus in a wadi cutting through the Sinjar anticline near the town of Sinjar, Iraq (west of Mosul). Sinjarite precipitates via slow evaporation of groundwater saturated with Ca and Cl ions. CaCl2 in the groundwater solution must be 3.5 times greater than NaCl for sinjarite to precipitate instead of halite. The extremely deliquescent nature of sinjarite means that the mineral is ephemeral and quickly dissolves in wet seasons or changes to the hexahydrite antarcticite when the temperature is less than 30°C (Aljubouri and Aldabbagh, 1980).


CaCl2 salts and liquid water flows on current-day Mars

On the current Martian surface it has been recently proposed that aqueous solutions form seasonally, via the deliquescence of hygroscopic salts in contact with atmospheric water vapor. Regions of these hygroscopic salts are thought to be indicated by occurrences of surface features known as recurring slope linae - RSL (Figure 7a, b; Chevrier and Rivera-Valentin, 2012). Older studies largely focused on perchlorate species as the most likely set of hydroscopic salts driving the periodic formation of RSLs, but another Mars-relevant set of salts, with similar low eutectic temperatures, relevant to ambient conditions with a significant deliquescence potential, are the calcium chloride hydrates, antarcticite and sinjarite. Gough et al. (2016) propose hydrated calcium chloride salts are linked to RSL formation on Mars. They also note deliquescence of CaCl2, not perchlorate, is also known to be responsible for the only terrestrial RSL analog known thus far: these are the seasonal water tracks in the McMurdo Dry Valleys, especially in the vicinity of the Don Juan Pond, as documented by Dickson et al., 2013 (Figure 8).

Soluble regolith salts like perchlorate and calcium chloride salts with low eutectic temperatures are likely to deliquesce at low relative humidity (RH) values at a wide range of temperatures (Gough et al., 2016). Deliquescence is the process by which a solid crystalline phase absorbs water vapor to form a saturated aqueous (liquid) solution. This phase transition from solid salt to liquid brine occurs at a deliquescence relative humidity (DRH), the value of which is specific for each hydration state of each salt and often varies with temperature. A salt is expected to be aqueous whenever the temperature is above the eutectic temperature of the salt and the relative humidity is above the DRH, although at much higher RH values ice may form. The low temperature deliquescence of perchlorates has been extensively studied (Gough et al., 2011, 2014; Nuding et al., 2014; Zorzano et al., 2009). Many perchlorate species have deliquescence relative humidity (DRH) values below 40% RH, and so should deliquesce and be stable or metastable liquids under Martian surface conditions (Chevrier et al., 2009), although slow reaction kinetics may limit their formation (Fischer et al., 2014). Not as much attention has been paid to other soluble, deliquescent salts, especially the calcium chloride salts, that may be present in the martian regolith.

Chlorine has been detected by rovers and landers in every Martian soil sample analyzed to date and is found at similar concentrations (≈0.2 to 1%) in all locations (Glavin et al., 2013). The form of the chlorine (i.e. chloride vs. perchlorate) is unknown at some sites, although measurements of the regolith at the Phoenix landing site confirmed chloride, perchlorate (Hecht et al., 2009) and probably chlorate (Hanley et al., 2012). The identity of the associated cation(s) is also generally unknown, although calcium, sodium, magnesium and potassium are the most likely candidates (Hecht et al., 2009). Regionally and globally widespread chloride deposits have also been detected by orbiting spectrometers (Figure 7c; Osterloo et al., 2010; Keller et al., 2006).

The global distribution of chloride deposits across the Martian surface is similar to that of recurring slope lineae (RSL), prompting hypotheses of a relationship between the two (Figure 7c; McEwen et al., 2011; Stillman et al., 2017). Chevrier and Rivera-Valentin (2012) suggest that CaCl2 is one of the best candidates for the formation of RSL, the narrow, dark features that appear and grow seasonally on Mars and appear to be caused by flowing liquid.


Two hydrated chloride salts, MgCl2 and CaCl2, may be the most appropriate salts because the seasonality of observed RSL formation best matches the seasonality of the melting of ice associated with these salts (Chevrier and Rivera-Valentin, 2012). Specifically, these chloride salts have eutectic temperatures less than or equal to the threshold temperature of 250 K that seems to mark the start of RSL activity (McEwen et al., 2011), but the eutectic temperatures of these salts are not so low that brines on Mars would be permanently liquid. Additionally, calculations show that CaCl2 in particular will cause seasonal melting of water ice throughout the top 20 cm of the martian regolith, therefore providing greater potential fluid flow than other salts (Chevrier and Rivera-Valentin, 2012). More recently, hydrated chloride salts (although not specifically calcium chloride) were observed to be present in RSL but not in surrounding regions (Ojha et al., 2015). It is not generally believed that deliquescence could be fully responsible for RSL formation because the small amount of water vapor in the martian atmosphere would limit the condensed phase water that could form. It is certainly possible, however, that salt deliquescence may be involved in the formation of RSL or in their appearance or behavior (McEwen et al., 2011; Ojha et al., 2015; Dickson et al., 2013; Stillman et al., 2017).

In summary, determining whether liquid water exists on the Martian surface is central to understanding the hydrologic cycle and potential for extant life on Mars (Ohja etal., 2015; Stillman et al., 2017). Recurring slope lineae, narrow streaks of low reflectance compared to the surrounding terrain, appear and grow incrementally in the downslope direction during warm seasons when temperatures reach about 250–300 K (-23°C to +26°C), a pattern consistent with the transient flow of a volatile material. Brine flows (or seeps) associated with seasonal deliquesence of hydrated salts (possibly hydrated calcium chloride salts or perchlorates) are proposed to explain the formation of recurring slope linea. As yet, no direct evidence for either liquid water or actual hydrated salt mineralogies has been found.


Tachyhydrite

Tachyhydrite (CaMg2Cl6.12H2O) (occasionally spelled tachydrite or tachhydrite) is a yellow, transparent to translucent, trigonal-rhombohedral mineral with very high solubility. Pure tachyhydrite is composed of 7.74% Ca, 9.39% Mg, 41.10% Cl, 37.09% O, and 4.67% H (Wardlaw, 1972). Its chemical formula is most accurately expressed as [Mg(H2O)6]2[CaCl6] because the structure consists of Mg(H2O)6 octahedra and CaCl6 octahedra loosely linked together by hydrogen bonds. Physical properties include; good cleavage, vitreous to greasy lustre, white streak, massive habit, bitter taste, deliquescent habit (dissolves or liquefies upon exposure to air), a density of 1.66 g/cm3, a hardness of 2, and it is hygroscopic (Braitsch, 1971).

Tachyhydrite is extremely soluble with 160 g dissolving in 100 g of water at 20 °C, and its solubility increases with increasing temperature (D’Ans, 1961; Wardlaw, 1972). Its hygroscopic and extremely deliquescent nature means at earth surface temperatures a crystal becomes liquid by absorbing moisture in the air. On exposure, tachyhydrite quickly alters to a residue of bischofite (MgCl2.6H2O) and a CaCl2.nH2O phase (Wardlaw, 1972). This is why tachyhydrite is typically documented in natural occurences where a crystal is still encased in halite. Tachyhydrite has a wide thermal stability field from 21.95°C to at least 167°C (Braitsch, 1971; Clark et al., 1980). The lowest temperature for tachyhydrite precipitation is 21.95°C at a concentration of 92.7 mol CaCl2/1000 mol H2O (450 g CaCl2/liter (D’Ans, 1961; Braitsch, 1971; Wardlaw, 1972). Its lowest temperature of formation increases by 0.0162ºC for every atmosphere of pressure (D’Ans, 1961).

Modern tachyhydrite occurrences

Natural tachyhydrite is documented only in a few modern hypersaline settings and it never forms a bed with primary precipitation textures (Figure 9). It is found in greater quantities in some ancient potash deposits. Tachyhydrite occurs seasonally as a minor interstitial cements and efflorescences, along with antarcticite (the hexahydrate form of calcium chloride), in the modern ephemeral halite crusts, atop sabkhas, of the Gavkhoni Playa (BWk), southeast of Isfahan, Iran (Pakzad and Ajalloeian, 2004), in mine wates in the Salar de Pedenales in the andean Altiplano, and perhaps as minor salt effloresecnes in the uppermost parts of the Abu Dhabi sabkha (Wood et al., 2005).

Much of the elevated ionic content of various Iranian playas in the Great Kavir is due to salt dissolution of the crests of nearby at- or near-surface diapirs and namakiers, where the mother salt sourced in halokinetic Miocene marine salts (Warren, 2008). These outcropping diapirs have carried carnallite and sylvite remnants into the namakiers that now reside at the surface (Rahimpour-Bonab and Kalantarzadeh, 2005). Tachyhydrite is also found as white feathery efflorescences within waste piles at an abandoned borate working along the southwestern margin of Salar de Pedernales, Chile (a BWk Koeppen climate location very near the ET climate boundary; Ericksen et al., 1989).

It seems that tachyhydrite’s high solubility means it has little or no preservation potential as a solid salt in any modern at-surface depositional setting; in the Peruvian case, it is a winter precipitate that disappears with the end of winter. In Abu Dhabi it, along with other highly soluble salts in the salty surface efflorescences, can disappear in morning dew (Warren, pers. obs.). In both the Iranian and the Peruvian settings the climate is cool and elevated. It seems the arid desert environments where bedded subaqueous-textured tachyhydrite accumulated in the drawndown Aptian seepage basins of the opening Atlantic were very different to its occasional rare occurrence in the evaporite settings of today.

Cretaceous of Brazil & W. Africa

Lower Cretaceous (Aptian) evaporite deposits in Sergipe, Brazil, and Gabon and Congo in western Africa contain significant amounts of tachyhydrite in halite-carnallite beds, along with other SO4-poor bitterns (Figure 10; Wardlaw, 1972; Borchert, 1977; de Ruiter, 1979; Hardie, 1990; Garrett, 1995; Zhang et al., 2017). These evaporite basins formed during the early Cretaceous rifting of Africa and South America. In each basin, potash-rich halite-dominant evaporites are transitional between older continental pre-rift and rift siliciclastics and younger post-rift marine shales and carbonates (Wardlaw, 1972, Borchert, 1977; de Ruiter, 1979; Szatmari et al., 1979; Hardie, 1990; Garrett, 1996).


The basic salt cycle of the Gabon and Congo basins includes from bottom to top: (1) thin black shale, (2) halite, (3) combination of halite and carnallite (carnallitite, and (4) bischofite and/or tachyhydrite (Figure 10;de Ruiter, 1979; Zhang et al., 2017). Although variations exist, a similar sequence is present in Sergipe, Brazil comprising the Ibura Member of the Muribeca Formation. The Ibura Member includes from bottom to top: (1) halite, (2) carnallite, (3) tachyhydrite, (4) sylvite (KCl) and halite, and (5) anhydrite (CaSO4) (Wardlaw, 1972; Borchert, 1977; Garrett, 1996). Tachyhydrite is located within the central and deepest portions of the Sergipe basins (Wardlaw, 1972; Borchert, 1977; Szatmari et al., 1979).

The origin of the evaporite sequences in Sergipe, Gabon and Congo is controversial with some authors suggesting a marine origin (Wardlaw, 1972; de Ruiter, 1979; Szatmari et al., 1979) or diagenetic origin (Borchert, 1977). More recent literature discussing Sergipe, Gabon, Congo, and other similar tachyhydrite deposits cite a non-marine/hydrothermal origin based on geochemical and textural studies (Lowenstein et al., 1989; Hardie, 1990; Garrett, 1996; El Tabakh et al., 1999). Yet other recent authors indicate a marine source based on its dominance in beds with a CaCl2 oceanic timing (Warren, 2016; Zhang et al., 2017).

The brine that tachyhydrite precipitated from must have been high in Ca because in waters with more SO4 than Ca present, the Ca is used-up by gypsum or anhydrite. Only if the Ca level is greater than SO4 in the early stages of precipitation will Ca stay in the brine allowing for potential precipitation of tachyhydrite (Wardlaw, 1972; Hardie 1990). The CaCl2-rich brine that produced the tachyhydrite and other potash salts is thought to have formed either by the interaction of hot meteoric groundwater with rift-related sediment and bedrock, or an influx of seawater at a time of a MgSO4-depleted ocean. Based on figure 1, Hardie (1990) suggests heated seawater could also have interacted with sediment and/or bedrock to produce a CaCl2 brine. This hot CaCl2 brine was perhaps driven to the surface by thermal circulation (Hardie, 1990). Alternatively, without arguing that the proportions of major ions in seawater change over time, the fact that modern marine water cannot precipitate tachyhydrite precludes it being the dominant mother brine in the Sergipe, Congo, and Gabon Basin evaporite basins,  (Lowenstein et al., 2014; Warren  2016).

Textures recovered in core, and visible in mine walls in Brazil, indicate Sergipe tachyhydrite was a primary bedded precipitate in those circum-Atlantic Cretaceous evaporite basins with stratiform tachyhydrite units in potash beds (Figure 10; Wardlaw, 1972). The Cretaceous world ocean at the time this tachyhydrite precipitated was a MgSO4-free marine bittern. Figure 9 shows tachyhydrite does not occur as a bittern phases in many other potash sequences deposited in basinwide settings at other times of CaCl2 oceans (e.g. the Devonian of Canada, the Silurian of the Michigan Basin). This has led to the argument that stratabound tachyhydrite in potash beds indicates an additional hydrothermal/basinal source for the mother brine. If so, the presence of tachyhydrite is not solely indicative of a MgSO4-depleted marine feed to the basin (Hardie 1990).

Cretaceous occurrences in Thailand

Another significant tachyhydrite occurrence is within slightly older mid Cretaceous (Cenomanian) evaporites of the Maha Sarakham Formation on the Khorat Plateau of northeastern Thailand (Figure 11). The majority of the Maha Sarakham Formation consists of alternating anhydrite and halite, separated by mudstone and siltstone “redbeds”. A 10- to 75-m-thick salt unit overlies a lower halite unit and underlies mudstone red beds of the lower clastic unit. Carnallite is most abundant in this salt unit, but sylvite and halite are also locally abundant. Tachyhydrite is present with halite and carnallite, but not with sylvite and is perhaps concentrated more in former basin centre or sump positions (El Tabakh et al., 1999; Warren 2016). Hite and Japakasetr (1979) estimated that tachyhydrite comprises less than 30% of the total carnallitite deposit, but tachyhydrite can form pure layers up to 16 m thick. Tachyhydrite crystals are euhedral and average 1 mm in size (El Tabakh et al., 1999). Sylvite and halite are also associated with accessory amounts of hilgardite [Ca2BCl(OH)2] and boracite (Mg3ClB7O13).


The origin of the Maha Sarakham Formation evaporites is still contentious; Hite and Japakasetr (1979) advocate a purely marine origin (highly restricted and concentrated sea), Hardie (1990) advocating an exclusively non-marine origin (hydrothermal CaCl2 brines), and El Tabakh et al. (1999) advocate an alternating marine and non-marine origin. They envisage an inland basin that was periodically inundated by the rising sea and at the time of tachyhydrite deposition received an additional influx of hydrothermal CaCl2 brines. Hardie (1990) states that formation of Maha Sarakham evaporites may have been similar to that of tachyhydrite formation in the Cretaceous basins of Brazil and western Africa, although he offers no evidence for the origin of the hydrothermal CaCl2 brines. El Tabakh et al. (1999) suggested that hydrothermal CaCl2 waters entered the restricted marine basin and created the right conditions for tachyhydrite precipitation and they cite contemporaneous granitic intrusions as possible evidence of thermal activity during the time of tachyhydrite formation.

Hydrothermal oceanic CaCl2 brine

Today, large volumes of relatively dilute, oceanic CaCl2 brines form via hydrothermal circulation and thermally-driven alteration of mid-ocean ridge (MOR) lithosphere (see part 1). These CaCl2 waters occur in and near active fracture zones, wherever seawater interacts with labile basalt (oceanic crust) at elevated temperatures and converts the circulating fluid from a Na-Mg-Cl water into a low-salinity Na-Ca-Cl brine, depleted in Mg and SO4. Similar hydrothermally-driven alteration of continental basalts via deeply circulated seawater interactions forms modern CaCl2-rich brine seeps, for example, within the thermally active continental Danakil rift valley (Hardie, 1990).

Tachyhydrite is a 'Swift Water" mineral

There may be an additional factor at work in terms of our recognition of tachyhydrite’s basin-by-basin distribution across the Phanerozoic. That factor is tachyhydrite’s high solubility in most subsurface waters and drilling muds. Even the name, tachyhydrite, was chosen from the Greek tachy hydros for “swift water” to indicate its rapid deliquescence on exposure to atmospheric conditions. This property means that any ground-up remains (rock chips) of tachyhydrite in a return mudstream in a conventional drilling operation typically do not make it to the surface. So, as most oil companies prefer not to core evaporites when drilling for hydrocarbons trapped beneath salt, there may be more tachyhydrite occurrences in deeply buried basin evaporites than the few currently documented.

Ancient tachyhydrite has only been found encased in halite in recovered cores and mine walls, never in drill cuttings, and its presence or absence can be easily misinterpreted in wireline logs, especially when drilling through thick salt masses in the search for oil and gas (companies tend to run a more limited log suite in thick salt units).

 

Tachyhydrite’s lack of shear strength and extremely high solubility makes mining any associated potash beds difficult; it presence leads to head beam and floor stability problems in conventional underground mines in Brazil and West Africa. Holle, a West African potash mine in the 1970s, was never economical and was lost to flooding after a few years of conventional mining and the Brazilian mine has ongoing floor and wall stability problems ties to tachyhydrite-induced heave. Today, solution mining is proposed as a more efficient means to exploit areas of Cretaceous potash associated with tachyhydrite beds, as in Aptian halite-potash associations on both sides of the Atlantic. High levels of tropical humidity mean at-surface potash evaporation pans are not viable in either area, so brine concentration and processing will probably require a more expensive option involving motorised dryers.

 

Not only is tachyhydrite highly soluble, but it is also highly reactive whenever in comes into contact with subsurface bitterns. Based on its occurrence in the Permian Stassfurt series in Germany, D’Ans (1961) suggested that much of the widespread polyhalite found in the CaCl2 brine-rich parts of the Zechstein Basin was a reaction product indicating former tachyhydrite. He noted that, at temperatures higher than room, there is a reaction between gypsum/anhydrite, sylvite and tachyhydrite-bischofite, or their dissolution-related brines, which removes tachyhydrite to form diagenetic polyhalite and calcium chloride brines;

CaMg2Cl6.12H2O+4KCl+8CaSO4--> 2CaSO4Mg(SO4)4(H2O)2 + 5CaCl2 + 4H2O

This was the mechanism suggested by Manheim (1974) to explain the lack of tachyhydrite, and the formation CaCl2-rich brines that characterise hydrothermal pore waters circulating in sediments associated with buried Miocene evaporites along the deep axial trough of the Red Sea.

Terrestrial CaCl2 minerals across time

The high solubility of CaCl2 minerals means that, with the exception of some parts of the Aptian Atlantic Potash association, even in past times of likely CaCl2-rich oceans, calcium chloride minerals are relatively rare as primary-textured occurrences in solid salt beds. On the other hand, as we saw in part 1, CaCl2 brines are commonplace in basinal or formation waters in many Phanerozoic sedimentary basins entraining thick sequences of dissolving ancient salt. For example, since 1914, CaCl2 brines have been extracted from Silurian strata adjacent to Salina Salt of the Michigan Basin USA, yet there are no significant volumes of tachyhydrite documented in the Michigan Basin(Garrett, 2004).

Historically, before micro-inclusion studies of chevron halite showed that the ionic proportions of seawater likely varied across the Phanerozoic, the various CaCl2 basinal brines occurrences in Silurian and Cretaceous strata were explained as an indicator of widespread dolomitisation and other diagenetic reactions, which preferentially extracted magnesium from pore waters. Since then, some authors have argued that CaCl2 enrichments in many ancient basinal brines, including brines in the Detroit group are partial leftovers of primary seawater chemistries (Lowenstein and Timofeeff, 2008). That is, Ca-enriched (MgSO4-depleted) pore brines are indicators of ancient CaCl2 oceans, with the pore brines being remnants from time the enclosing evaporitic and marine sediments were deposited (relict or connate brines).

Others, such as Houston et al. (2011), conclude this is not necessarily so, they agree that there are two end-members typifying highly saline subsurface brines. But they conclude that end-member chemistries relate to either substantial subsurface halite dissolution, or to preservation of early reflux-related seawater. Houston et al. (2011) go on to argue that CaCl2-enriched formation water chemistries from many basins worldwide, including the Michigan Basin, do not support an interpretation of variation in ionic proportions in seawater across the Phanerozoic. They find that CaCl2-rich brines formed either by dissolving bittern salts in the subsurface, or simply lost water in the subsurface after significant rock-fluid interaction had taken place. Water loss might be achieved by interaction with a gas phase at the elevated temperatures of deep burial or, alternatively, water may have been lost to clays. Both these mechanisms would have the effect of dehydrating (concentrating) the brine.

However, whatever the mechanism, it does seem that there is a Quaternary continental/pre-Quaternary marine dichotomy in the nature and distribution of CaCl2 brines and the associated antarcticite/tachyhydrite minerals, much as there is for the world’s potash deposits (MgSO4-enriched versus MgSO4-depleted). On Earth, there are no documented examples of pre-Quaternary bedded antarcticite. The situation may be different on Mars.

Today terrestrial antarcticite either precipitates in the hot-arid Bristol Dry Lake or the icecap-polar setting of Don Juan Pond, neither is associated with brine sourced in seawater or derived by the recycling of older (“connate”) marine evaporite successions. Rather the occurrenceof the calcium chloride minerals is in part a reponse to low temperatures and regolith cycling. Other present-day examples of regions with somewhat enriched levels of CaCl2 surface waters, lack both tachyhydrite and antarcticite precipitates, and contain higher levels of Mg that are tied to deeply circulated marine/hydrothermal waters and variable interactions with MgSO4-enriched marine salts (e.g., Lake Asal and the Dallol Salt Pan). CaCl2 brines of the Dead Sea and the Red Sea show even more elevated levels of MgCl2 as they are derived, at least in part by the fractionated dissolution of bedded and halokinetic Miocene marine evaporites. The Mg-rich clay palygorskite is co-associated with these CaCl2 brines in the subsurface of the Red Sea and the Dallol.

In contrast, tachyhydrite in some ancient marine-fed basinwide evaporite associations is found as somewhat rare, but at times bedded, units in the bittern-rich portions of the halite succession, as in the Cretaceous basinwide evaporites of Brazil, the Congo and Thailand, and as a cementing phase in the Permian Stassfurt 2 in the Zechstein (where it is commercially extracted in association with MgCl2 brines).

All the ancient tachyhydrite examples mentioned above, are associated with the presence of widespread potash salts within adjacent salt beds. However, there are many other even larger and richer ancient potash deposits, such as the intracratonic Alberta basin (Devonian) and the Kama basins (Permian) where no tachyhydrite is documented. It seems that the terrestrial precipitation of bedded tachyhydrite is not just favoured by times of CaCl2 oceans, it also requires additional input from saline hydrothermal/basinal waters. Such settings are most likely in the transition phases in an actively-opening hydrothermally-influenced continental rift as passes into the marine seepage realm at a time when the adjacent ocean was a MgSO4-depleted system.

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Calcium Chloride (CaCl2), Article 1 of 2: Usage and brine chemistry

John Warren - Sunday, April 30, 2017

 

Introduction

Calcium chloride minerals in the natural state are rare and only found in a few specific evaporite associations. On the other hand, calcium chloride-rich brines are commonplace in the burial diagenetic realm, especially in deep high-salinity basinal brines and in a number of hypersaline lake waters, especially in rift settings. In the subsurface, these brines also play a significant role in the formation of a number of metal ores. Occurrences of both the brine and the minerals have significance in modelling rock-fluid interactions and seawater chemistry across geological time.

At earth-surface temperatures, calcium chloride can exist in the solid state as the anhydrous form (CaCl2) as well as in four levels of hydration – CaCl2.H2O; CaCl2.2H2O; CaCl2.4H2O; CaCl2.6H2O (Table 1). Of this group, CaCl2 occurs naturally as two rare minerals; antarcticite and sinjarite. All of the early studies on calcium chloride and its hydrates were done with laboratory-prepared samples of brines and hydrates, since CaCl2 was not produced on a commercial scale until after the ammonia–soda process for the manufacture of soda ash (Solvay Process) was in operation. Before its industrial uses were discovered, calcium chloride was considered a waste product of brine production. Today, its primary industrial use is predicated on the very high enthalpy change of solution, indicated by considerable temperature rise accompanying dissolution of the anhydrous salt in water (Table 1 – Heat of solution in water). This property is the basis for its largest scale application, namely road de-icing.

 

In the natural state, most CaCl2 occurs in solution in basinal waters in sedimentary basins and modified pore waters in specific hydrothermal associations. Calcium chloride, in a mineral state in the natural world, occurs as the rare evaporite minerals; sinjarite (CaCl2.2H2O) and antarcticite(CaCl2.6H2O). The related potassic and magnesian calcium chloride minerals, chlorocalcite (KCaCl3) and tachyhydrite (calcium magnesium chloride, CaMg2Cl6•12H2O) are also rare in the sedimentary realm, and have particular evaporite associations and implications (see Part 2).

Outside of an industrial byproduct of the Solvay Process, most CaCl2 is derived from the processing of hypersaline basinal brines. The only current commercially, exploited natural CaCl2 surficial brine source is in Bristol Dry Lake, California (Figure 1). In the USA, for example, basinal brines are the primary commercial source of calcium chloride. Some of these brines in Michigan, Ohio, West Virginia, Utah, and California contain >4% calcium, with the Michigan Basin as the dominant producer. In the USA, a former commercially important source of calcium chloride was as a by-product of the Solvay Process used to produce soda ash. Because of environmental concerns and high energy costs, the Solvay Process has been discontinued as a source of CaCl2 in the USA.

This article will focus on the utility and geological significance of CaCl2 brines, while the next will focus on the geochemical significance of various calcium chloride minerals in particular evaporitic settings.


Usage

Calcium chloride depresses the freezing point of water, and its principal use is to prevent ice formation, especially on winter roads. Calcium chloride released to the environment is relatively harmless to plants and soil in diluted form. As a de-icing agent, it is more efficient at lower temperatures than sodium chloride. Solutions of calcium chloride can prevent freezing at a temperature as low as -52 °C (Figure 2). Hence, more than 50% of world CaCl2 usage is for road de-icing during winter. The second largest application of calcium chloride brine exploits its hygroscopic properties and the tackiness of its hydrates. In summer, it is used for roadbed stabilisation in unsealed roads and as a dust palliative. When sprayed onto the road surface, a concentrated CaCl2 solution maintains a cushioning layer on the surface of dirt roads and so suppresses formation of road dust. Without brine treatment dust particles blow away, eventually larger aggregate in the road also begins to shift around, and the road surface breaks down. Using calcium chloride reduces the need for grading by as much as 50% and the need for fill-in materials as much as 80%.


Calcium chloride’s low-temperature properties also make it ideal for filling agricultural implement tyres as a liquid ballast, aiding traction in cold climates. It is also used as an accelerator in the ready-mix concrete industry, although there is concern about its usage because of possible long-term chloride-induced corrosion of steel in highways and buildings. Calcium chloride is also widely used to increase mud fluid densities in oil- and gas-well drilling. It is also used in salt/chemical-based dehumidifiers in domestic and other environments to absorb dampness/moisture from the air.

Calcium chloride is used in the food industry to increase firmness of fruits and vegetables, such as tomatoes, cucumbers, and jalapenos, and prevent spoilage during processing. Food-grade calcium chloride is used in cheese-making to aid in rennet coagulation and to replace calcium lost in pasteurisation. It also is used in the brewing industry both to control the mineral salt characteristics of the water and as a basic component of certain beers.


Calcium Chloride brines production

Generally, CaCl2 brines are found in permeable strata either below, adjacent to, or above evaporite deposits, gradually becoming more dilute as brines approach the surface, and modified somewhat in proportion to distance from a potash or salt layer (Figure 1; Table 2). Other natural calcium chloride brines are derived from hydrothermally-modified marine waters. Dilute calcium chloride brines are also occasionally found in coastal aquifers, and some oil or gas formation waters that have been formed from seawater, possibly by a dolomitization reaction supplemented by the leaching of certain types of rocks (Garrett, 2004).

Basinal brines are chemically similar to CaCl2 brines forming hydrothermally at modern mid-ocean ridges, where seawater is being converted by interaction with basalt at elevated temperatures into low-salinity Na-Ca-Cl brines, depleted in Mg and SO4. These CaCl2 waters occur in and near active fracture zones, wherever seawater interacts with labile basalt (oceanic crust) at elevated temperatures and converts the circulating fluid from a Na-Mg-Cl water into a low-salinity Na-Ca-Cl brine, depleted in Mg and SO4. Similar hydrothermally-driven alteration of continental basalts via deeply circulated seawater interactions forms modern CaCl2-rich brine seeps, for example, within the thermally active continental Danakil rift valley (Hardie, 1990; Warren, 2016).

 

Calcium chloride is produced in commercial amounts using a variety of procedures: 1) refining of natural brines, typically with heating to increase concentration, 2) reaction of calcium hydroxide with ammonium chloride in Solvay soda ash production, and 3) reaction of hydrochloric acid with calcium carbonate. The first two processes account for over 90% of the world’s total calcium chloride production. Historically, natural brines sources are the dominant CaCl2 source. There is currently an excess of capacity in the calcium chloride industry, a situation which is only expected to become more acute as synthetic and byproduct capacity increases.

As we shall see in the next article, calcium chloride crystals are relatively rare as primary-textured occurrences in solid salt beds. On the other hand, CaCl2 brines are commonplace in basinal or formation waters in many Phanerozoic sedimentary basins entraining thick sequences of dissolving ancient salt. For example, since 1914, CaCl2 brines have been extracted from Silurian strata adjacent to Salina Salt of the Michigan Basin USA (Figure 1; Ludington; Oxy Chemicals, Formerly Dow Chemicals)). These brines are recovered from Detroit Group sediments that overlie the Silurian Salina evaporites. Based on fluid inclusions in primary halite chevrons in the Salina Salt, the Silurian was a time characterised by a CaCl2-enriched MgSO4-depleted ocean (but no tachyhydrite is known from Silurian strata, here or elsewhere). Industrial production of CaCl2

CaCl2 brines in the Michigan Basin (commercial)

This Silurian halite/potash basin has many aquifers with calcium chloride brines, both above and below the Silurian Salina Group’s halite and potash levels (Figures 3, 4; Garrett, 2004). Major aquifers are the overlying Devonian carbonate and sandstone beds, with many lesser aquifers. In the first porous bed above the potash (Sylvania Sandstone Formation) there is an extensive area of rich calcium chloride brine sitting directly above the potash deposit and extending to the south-southeast. Brine concentrations at nearly the same concentration as potash end liquor in fractures in the intrasalt carbonates. The less voluminous sandstone of the Filer Formation to the northwest contains a similar, but slightly more dilute CaCl2 brine. Several thinner and less abundant aquifers also occur under the potash beds with equally strong, or stronger calcium chloride brines (Figure 3).


The porous 0-90 m thick Sylvania Sandstone lies at the base of the Detroit River Group and is the main source of CaCl brine production. It is in direct erosional contact with the salt succession (Figure 4). The remainder of the Group consists of 0-350 m of variably porous carbonates (Garrett, 1995, 2004). Both sandstones and carbonates contain CaCl2-rich brines and extend across some 40% of the Michigan Basin at depths of from 300-1,400m. Brine concentration and the relative amount of CaCl2 increases with depth. Typically, the brines are only considered to be of economic importance below about 880 m depth. In carbonate hosts, the CaCl2 content varies from 3-23% and KCl from 0.2-2.9%, usually increasing in concert with concentration, as the NaCl content decreases. CaCl2 content in the Sylvania Sandstone varies from 14-22%, KCl from 0.6-2.1%, and both are more uniformly concentrated compared to the carbonate-hosted brines (Garrett, 2004).


Each aquifer entrains roughly the same ratio of ions, but pore waters become progressively more dilute as beds approach the surface about the basin margin. It seems likely that in this basin, a potash liquor originally seeped through and under the potash deposit (and reacted with calcite) was much later forced from its original sediments into the overlying porous strata into the overlying porous strata as they were compressed by deep burial, possibly aided by load-induced pumping induced by the waxing and waning of thick glacial ice that formed over this basin (McIntosh et al., 2011). Variable ionic content, as seen in Table 2, results from their considerably different migration history and variable dilution by meteoric or other groundwater (as is strongly indicated by the brine’s deuterium and 18O analyses), precipitation (such as gypsum), and their different contact with rocks that they could partially leach or react with. However, in the Michigan Basin these reactions were limited, since the porous carbonate strata (average, 20%) contains fairly pure carbonates, and the sandstone strata fairly pure silica (quartz arenites) cemented by dolomite or quartz (Martini, 1997).

There is a general synclinal structure to the strata under the Michigan Basin, and examples of the specific stratigraphy to the southeast of the centre of the basin at Midland are shown in Figure 4. The Detroit River Group consists of 0–350 m of variable porosity carbonates, and at its base there is 0–90 m of porous sandstone called the Sylvania Formation. Each of these formations cover about 40% or more of the Michigan Basin, and contain strong calcium chloride brines at depths of 300–1400 m. Their brines have been commercially recovered in the past, and were generally only considered to be economic below about 880 m. The brines’ total dissolved solids (TDS) and the amount of CaCl2 increases fairly consistently with depth from 3 to 23% CaCl2, and the NaCl and MgCl2 concentrations vary inversely with the CaCl2 . In the Sylvania Formation, the CaCl2 usually ranges from 14 to 22% (Figure 3). Additional information on the brine in other aquifers and the various reactions and changes that have occurred with them are discussed in Martini a(1997), Wilson and Hewett, (1992) and Wilson and Long (1992).

The Michigan Basin brines’ very low pH (4.5 to 5.3) helps to explain an ability to leach and react with other rocks, as is indicated by their high contents of strontium, barium and metals, much of the Sr and Ba probably came from the reaction with calcite. Geothermal water also probably mixed with some of the formations, as indicated by the variable presence of iodine, boron, lithium, caesium, rubidium and other rare metals. With most of the brines, the calcium concentration is somewhat higher than its magnesium equivalent in seawater end-liquor from a potash deposit, and the potassium a little lower. Wilson and Long (1992) speculated that this occurred by the conversion of the clays kaolinite and smectite to illite: Small amounts of glauberite (CaSO4.Na2SO4) and polyhalite (2CaSO4.K2SO4.MgSO4; have also been found in the basin. Finally, some of the calcium chloride aquifers have a slightly elevated ratio of 87Sr/86Sr (range from 0.7080 to 0.7105; seawater is 0.70919), further indicating that there was some rock leaching during burial(Martini, 1997).

What does a CaCl2 basinal brine indicate?

Pore fluids in the deeper parts of many sedimentary basins, especially if they contain a significant unit of evaporite, tend to be CaCl2 brines, entraining large volumes of hypersaline brine and in places, hydrocarbons (e.g., Michigan Basin, the U.S. Gulf Coast, European North Sea Basin, Western Canada Basin and Volga Basin).

Worldwide, one of the principal geochemical characteristics of saline waters in sedimentary basins is the progressive shift in their major ion composition from Na–Cl to Na–Ca–Cl to Ca–Na–Cl dominated waters with increasing chlorinity or salinity (Hanor and McIntosh, 2006). Such basinal brines (also called oil-field brines or formation waters) with significant calcium chloride contents have salinities that typically range up to 300,000 mg/l (Hanor, 1994; Lowenstein et al., 2003). The majority of these basinal brines are chemically distinct in their high Ca concentrations, separating their hydrogeochemistries from modern seawater and other common surface and near-surface waters which tend to be Na-Cl-SO4, Ca-HCO3, or Na-CO3 types (Drever, 1997).

Calcium levels in a CaCl2 basinal brine typically exceed the combined concentrations of SO4, HCO3, CO3 ions, (specifically, mCa > ∑(mSO4 + 1/2mHCO3 + mCO3); Lowenstein et al., 2004). And yet, the evaporative concentration of modern seawater leads to brines depleted in Ca, as required by the principle of chemical divides (CaCO3 and CaSO4 divides) for any evaporating water (Hardie and Eugster, 1970). Explanations for the origin of CaCl2 basinal brines remain problematic.

There is no simple pathway by which modern seawater, and most other surface and near-surface waters trapped in sedimentary deposits, can be converted to CaCl2 basinal brines during burial, without invoking significant rock-fluid interaction. Historically, before micro-inclusion studies of chevron halite showed that the ionic proportions of seawater likely varied across the Phanerozoic, the various CaCl2 basinal brines occurrences, for example in in Silurian and Cretaceous age strata were explained as an indicator of widespread dolomitisation and other diagenetic reactions, which preferentially extracted magnesium from pore waters (Garrett, 2004).

Since the mid-1990s, others have argued that CaCl2 enrichments in many ancient basinal brines with thick evaporites in the stratigraphy, including brines in the Detroit Group, are partial leftovers from times of CaCl2-enriched seawater chemistries (Table 3; Lowenstein and Timofeeff, 2008; Lowenstein et al., 2014). That is, Ca-enriched (MgSO4-depleted) pore brines adjacent to thick evaporites are indicators of ancient CaCl2 oceans, with the pore brines being remnants from the time the enclosing evaporitic and marine sediments were deposited (relict or connate brines).


Other authors, such as Houston et al. (2011), conclude this is not necessarily so; they agree that there are two end-members typifying highly saline subsurface brines (MgSO4 depleted or enriched). But, they conclude that end member chemistries relate to either substantial subsurface halite dissolution, or to preservation of early reflux-related seawater. Houston et al. (2011) go on to argue that CaCl2-enriched formation water chemistries from many basins worldwide, including the Michigan Basin, do not support an interpretation of variation in ionic proportions in seawater across the Phanerozoic. They find that CaCl2-rich brines formed either by dissolving bittern salts in the subsurface, or simply lost water in the subsurface after significant rock-fluid interaction had taken place. Water loss might be achieved by interaction with a gas phase at the elevated temperatures of deep burial or, alternatively, water may have been lost to clays. Both these mechanisms would have the effect of dehydrating (concentrating) the brine.

Calling upon a CaCl2 seawater source as an explanation for the origin of basinal brines was criticized also by Hanor & McIntosh (2006) who pointed out that no matter what the starting composition of a paleoseawater, significant diagenetic alteration must have occurred to produce the present major ion chemistry of Illinois and Michigan basin brines specifically, and basinal brines in general. In their view, the diagenetic mineral–brine interactions that occur during burial mask any original compositional variations in the starting seawater.

Hanor & McIntosh (2006) also argued that due to ongoing fluid escape and crossflow, it is difficult, if not impossible, to assign specific ages to basinal brines in a sedimentary basin. If the age of a basinal brine is not known, then the possible parent seawater, whether CaCl2 or MgSO4 type, cannot be determined. Hanor and McIntosh (2007) illustrated further complications in the interpretation of the timing of the origin of basinal brines. They showed that some brines in the Gulf of Mexico basin were not formed during the Middle Jurassic, contemporaneous with deposition of the Louann Salt, but formed during the Cenozoic from the dissolution of the Louann salt.

Interestingly, many marine potash deposit end-liquor brines have a high to medium–high lithium content, such as the Angara-Lena basin, Russia’s 1600–1900 ppm, the Paradox Basin’s 66–173 ppm Li, the Michigan Basin’s Sylvania Formation’s 36–72 ppm and the English Zechstein Formation’s 7– 65 ppm, etc. (Garrett, 2004)). However, some end-liquors have only a nominal lithium contents, such as from the Saskatchewan, Canada potash deposits. A few calcium chloride lakes also have medium–high values, such as the Don Juan Pond’s 235 ppm, Bristol Lake’s 30–108 ppm, Cadiz Lake’s 20–67 ppm, and Lake Vanda’s 27 ppm (Figure 1). We shall come back to this topic in a future article that will focus on lithium-rich brines.

Origin of CaCl2 brines

So, currently, there are two schools of thought used to explain the origin of CaCl2 basinal brines in evaporitic basins. One school assumes that the chemistry of the world’s ocean and its ionic proportions have remained near constant across the Phanerozoic. To form a CaCl2 enriched basinal brine then requires substantial subsurface rock-fluid interactions, utilising mechanisms and processes that include nonmarine parent waters, diagenetic alteration, pervasive dolomitization of carbonates, or bacterial sulphate reduction. All these mechanisms can reduce the proportion of Mg, HCO3 and SO4 relative to Ca in subsurface pore waters (Hanor and McIntosh, 2006). Proponents of this school tend to base their argument on basin-scale variations in the hydrogeochemistry of pore fluids.

The other school (mostly based on the micro-inclusion chemistry of chevron halite) argues for long-term changes in the major ion chemistry of seawater (Table 2). For example, Upper Jurassic, Cretaceous, and Cenozoic seawater records a systematic, long-term (>150 My) shift from the Ca2+ - rich, Mg2+ - and SO42- - poor seawater of the Mesozoic (“CaCl2 seas”) to the “MgSO4 seas” (with higher Mg2+ and SO42- > Ca2+) of the Cenozoic (Lowenstein and Timofeeff, 2008; Lowenstein et al., 2003, 2014). Over that period, the Mg/Ca ratio of seawater rose from 1 in the Early Cretaceous, to 2.3 in the Eocene, and 5.2 in present-day seawater.

Suggested drivers of long-term variation in the major ion chemistry of seawater include; fluctuations in the volume of discharge of hydrothermal waters from the global mid-ocean ridge system (Hardie, 1996), changes in the rates of volcanic activity and weathering processes, and variations in the amount of dolomite formed in the oceans (Holland and Zimmermann, 2000).

In my mind, much of the conflict between to two schools of thought as to the origin of CaCl2 basinal brines stems from the source of evidence. One approach utilises micro-inclusion chemistry in halite chevrons to define the evolution of Phanerozoic seawater. This data is extracted from an intra-salt textural association that, due to its long lack of permeability (‘locked up in halite’), likely preserves the chemical composition of the original depositional setting (e.g. Zambito et al.). The other school focuses on pore fluid hydrochemistry in subsurface waters, generally using water samples in boreholes, collected from pores and fractures in a carbonate, sandstone or shale host. The nature of fluids in these non-salt sediments, some which have been permeable since deposition, mean fluids experienced ongoing re-supply via crossflow and rock-fluid interaction as the ambient temperature, pressure and salinities evolved in the burial environment. This process shutdown once matrix permeability was lost (Warren et al., 2014).

In the next article, we shall expand our discussion of the significance of CaCl2 brines with a close look at where and how particular calcium chloride minerals can precipitate and be preserved and why some types of calcium chloride salts are more common in particular evaporitic settings.

References

Drever, J. I., 1997, The geochemistry of natural waters: Surface and groundwater environments: New Jersey, Prentice-Hall Inc., p. 327-351.

Garrett, D. E., 1995, Potash: Deposits, processing, properties and uses: Berlin, Springer, 752 p.

Garrett, D. E., 2004, Handbook of Lithium and atural Calcium Chloride; Their deposits, processing, uses and properties Amsterdam, Elsevier Academic Press, 476 p.

Hanor, J. S., 1994, Origin of saline fluids in sedimentary basins: Geological Society, London, Special Publications, v. 78, p. 151-174.

Hanor, J. S., and J. C. McIntosh, 2006, Are secular variations in seawater chemistry reflected in the compositions of basinal brines?: Journal of Geochemical Exploration, v. 89, p. 153-156.

Hardie, L. A., 1990, The roles of rifting and hydrothermal CaCl2 brines in the origin of potash evaporites: a hypothesis: American Journal of Science, v. 290, p. 43-106.

Hardie, L. A., 1996, Secular variation in seawater chemistry: an explanation for the coupled secular variation in the mineralogies of marine limestones and potash evaporites over the past 600 m.y.: Geology, v. 24, p. 279 - 283.

Holland, H. D., and H. Zimmermann, 2000, The Dolomite Problem Revisited: Int. Geol. Rev., v. 42, p. 481-490.

Houston, S., C. Smalley, A. Laycock, and B. W. D. Yardley, 2011, The relative importance of buffering and brine inputs in controlling the abundance of Na and Ca in sedimentary formation waters: Marine and Petroleum Geology, v. 28, p. 1242-1251.

Lowenstein, T., B. Kendall, and A. D. Anbar, 2014, Chapter 8.21. The Geologic History of Seawater, Treatise on Geochemistry (2nd Edition), Elsevier, p. 569-621.

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

Lowenstein, T. K., and M. N. Timofeeff, 2008, Secular variations in seawater chemistry as a control on the chemistry of basinal brines: test of the hypothesis: Geofluids, v. 8, p. 77-92.

Martini, A. M., 1979, Hydrogeochemistry of Saline Fluids and Associated Water and Gas, Michigan Basin: Doctoral thesis, University of Michigan, 236 p.

McIntosh, J. C., G. Garven, and J. S. Hanor, 2011, Impacts of Pleistocene glaciation on large-scale groundwater flow and salinity in the Michigan Basin: Geofluids, v. 11, p. 18-33.

Warren, J., C. Morley, T. Charoentitirat, I. Cartwright, P. Ampaiwan, P. Khositchaisri, M. Mirzaloo, and J. Yingyuen, 2014, Structural and fluid evolution of Saraburi Group sedimentary carbonates, central Thailand: A tectonically driven fluid system: Marine and Petroleum Geology, v. 55, p. 100-121.

Warren, J. K., 2015, Seawater chemistry (1 of 2): Potash bitterns and Phanerozoic marine brine evolution, Salty Matters blog, www.saltworkconsultants.com.

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

Wilson, T. P., and T. A. Hewett, 1992, Geochemistry and isotope chemistry of Michigan Basin brines: Devonian formations: Applied Geochemistry, v. 7, p. 81-100.

 

Salt, Oil, Gas & Metals: What Drives the Link?

John Warren - Friday, March 31, 2017

Introduction

This article is based on a review written for CSPG Reservoirs and presented orally at the AAPG 2016 meeting in Calgary. It endeavours to give an up-to-date synopsis of how and why ancient salty basins tend to contain elevated levels of oil, gas and metals. It begins with an overview of hydrologies, then places ancient salt bodies in their climatic and tectonic context, and lastly looks at industrial associations and predictors. For brevity, details of many regions, deposits and references are summarised as tables, more complete discussion and referencing can be found in appropriate sections in Warren (2016), or feel free to contact me (jkwarren@saltworkconsultants.com or via www.saltworkconsultants.com) with requests for more comprehensive documentation of particular examples.


Evaporite styles reflect intrabasin brine hydrology

If we accept the definition of an evaporite as “A salt rock originally precipitated from a saturated surface or near-surface brine in hydrologies driven by solar evaporation,” then the greater volume of saline mineral salts in the earth’s sedimentary realm are the product of solar heating of brine. There are other sets of mineral salts in the depositional and diagenetic realm with the same mineral composition as evaporite salts, but these salts result from cryogenic, hydrothermal or burial re-equilibration processes (Warren, 2016). For a water molecule to escape into the vapour phase in an evaporitic setting, and so increase the salinity of an enduring brine body or pore water in the capillary zone of a sabkha, the water molecule must; 1) absorb heat energy, 2) be located near the liquid surface, 3) be moving in the proper direction, 4) have sufficient energy to overcome liquid-phase intermolecular forces and, 5) pass through the surface tension inter-face (Figure 1).


Simple physics of molecular escape (rate and intensity) during solar evaporation essentially controls the potency of re-maining brine. Bed textures and mineralogies entomb evidence of concentration levels in the brine and its hydrological setting/stability (Figures 2, 3). This makes evaporites excellent gauges of climate present and past. In combination, relative brine density and specific heat capacity of adjacent brine masses control density and thermal stratification in saline brine bodies located at or near the earth’s surface (Figure 2). Subsequent burial alteration controls later textural evolution by interactions with regional shallow and deep phreatic crossflows (Warren, 2016).

Specific heat is the amount of heat needed to raise one gramme of a substance by 1 °C. For a given amount of heat input, a unit volume of hypersaline water will show a greater increase in temperature than a less salty water (Figure 2c). Given the same degree of insolation, this means density-stratified water bodies tend to be heliothermic, with the lower, denser bottom brine layer being warmer than the somewhat less saline, less dense, upper water layer (Figure 2a). The level in any water column where a marked change in temperature occurs is the thermocline, and in a density-stratified brine mass corresponds to the halocline, also termed a chemocline (Figure 3a). The combination of high temperature, high salinity and lower oxygen levels in the lower brine mass in a heliothermal brine system means only a specialised biota can survive there, often with specialised bacterial populations living in waters just above a saline thermocline. A modern example is the purple sulphur-oxidising community flourishing immediately above the halocline in Lake Mahoney, British Columbia. Fluctuating salinity and nutrient levels endemic to many evaporite depositing regions encourage preservation of elevated levels of organic matter in a variety of hypersaline settings past and present (“feast and famine” associations; Warren, 2011).

As any brine concentrates, its density increases (Figure 2b). The overlying water body must be holomictic for bottom-nucleated salts to accumulate across the subaqueous floor of a brine mass and for dense, saturated brine to sink (reflux) into underlying sediments (Figure 3a). Holomixis means a near homogenous distribution of brine density, temperature, and salinity throughout the brine mass, with internal mixing being ongoing and mostly maintained by wind movement. In contrast, a meromictic brine body is internally stratified, with a lower more-saline, denser, warmer water mass separated across a halocline from an upper, less saline, less-dense, cooler water mass. A longterm halocline hinders chemical or physical changes in the underlying denser waters, so shutting down bottom nucleation, as well as slowing and ultimately stopping brine reflux. A permanently stratified system is ectogenic, while a brine column that is temporarily stratified is endogenic.

Over decades, saline water masses can change from ectogenic to endogenic. In February 1979, salinity equalisation drove the mixing of upper and lower water masses in the Dead Sea, resulting in a holomictic water body. Since then, aside from short episodes of storm-flood-driven freshening of the upper water mass in 1980 and 1994 (meromixis), the Dead Sea has been holomictic, and halite has been accumulating on the deep lake floor (Gertman and Hecht, 2002). Before 1979, the Dead Sea had been a stratified system for at least 400 years and only pelagic carbonate laminites with minor gypsum, not halite, accumulated on the deep lake floor, beneath a 370-380 m deep brine column.


Holomixis permits deposition of a coherent salt layer across the whole basin floor, beneath both shallow and deep brine columns. Density stratification allows evaporitic salts to crystallise only in the upper water mass or at the upper brine - lower brine interface, so bottom nucleation tends to occur on the shallower lake floor, where it lies above the halocline (Figure 3a). That is, long-term (ectogenic) column stratification mean bottom nucleation of salts can only occur where the upper salt-saturated brine mass intersects the sediment bottom, with a pelagic settling of salts occurring deeper out in the depositional basin, as in the Dead Sea prior to February 1979 (Figure 3a). The bottom growth of crystals cannot occur on a deep bottom located beneath a density-stratified system, as there is no mechanism to drive ongoing supersaturation in the lower water mass. For the same reason, constant brine reflux driving sinking of a dense brine into sediments beneath the floor of the evaporite basin can only proceed if significant regions of the overlying brine mass are holomictic. Deposition of capillary salts (sabkha deposits) occurs in subaerial settings, wherever the saline capillary zone inter-sects the land surface (Figure 3b).

When salts are accumulating beneath a holomictic brine mass, textures in bottom nucleates is controlled by the stability of the overlying brine column (Figure 3b). When the overlying column is deep (>30-100m) then, other than areas on the deep bottom of local phreatic spring-fed outflows, there is no general hydrochemical mechanism to drive fluctuations in bottom-brine chemistry. The resulting deep bottom precipitates tend to be monomineralogic crystal clusters, possibly encased by re-transported material washed in from the shallower surrounds (Figure 3a, 3b). In contrast, when the overlying brine column is shallow (<30m and typically <5-10m) then the chemistry and stability of the brine varies on a shorter term (daily-weekly) basis, so more layered bi-mineralogic bottom-nucleates can accumulate as layered to laminated salt beds. In addition, all evaporite sediments can be reworked by bottom currents, with similar textures to those that characterise siliciclastic and mechanically-modified carbonate sediments (Figure 3b).

"Now" versus "then" in evaporite deposition

Uniformitarianism is an essential tenet of geological understanding. Yet, when we look at evaporite volumes and depositional settings across deep time, we see that the diversity of modern evaporite analogues is constrained by a deficit in two conditions, specifically; 1) the current lack of greenhouse eustasy; contemporaneous atmospheric conditions and sea levels are controlled by the earth’s current icehouse climate mode and have been for the last 10-12 million years, and 2) the current lack at the plate-edge scale of marine seepage into large hydrographically-isolated oceanic sump basins (Warren, 2010). Both situations circumscribe different hydrologies and eustasies compared to continental-fed in-flows that typify the world's current larger evaporite basins. Today, and across the Quaternary, the largest and thickest salt stacks, with areal extents up to 10,000 km2 and thickness up to 900 m, tend to precipitate in the lower parts of suprasealevel intermontane lacustrine sumps located in tectonically active parts of continental interiors, such as Salar Atacama and Salar di Uyuni in the Andean Altiplano (Figure 4a). Most ancient evaporites are marine-fed and were deposited in huge hydrographically-isolated subsealevel marine-seepage sumps located in intracratonic basins or within rifts or compressional sutures. Often, the areal extents of these ancient systems were more than 250,000 km2, this is more than two orders of magnitude larger than any Quaternary evaporite deposit. Bedded (pre-halokinetic) thicknesses could be more than a kilometre.


Ancient marine saline giants (megahalites and megasulphates) accrued in either of two plate-scale settings, which at times merged into one another, namely; 1) Platform evaporites (Figure 5) and 2) Basinwide evaporites (Figure 6). The first major contrast with nonmarine continental dominance in Quaternary evaporite settings is the fact that platform evaporites require greenhouse eustasy, the second is that basinwide evaporites require tectonically- and hydrographically-isolated widespread subsealevel depressions, typically found along plate edges with continent-continent proximity (Figure 5).

Neither condition is present on the current earth surface. For basinwides, suitable hydrologic conditions were last present during the Messinian Salinity Crisis in the Mediterranean region, and platform evaporite settings were last present on earth across large parts of the Middle East carbonate platform during the Eocene (Tables 1, 2). There is a third group of ancient evaporite deposits; it encompasses all nonmarine lacustrine beds past and present (Table 3). This group has same-scale modern-ancient counterparts, unlike ancient marine platform and basinwide evaporites (Figure 4a; Warren, 2010, 2016).


Platform evaporites

Are made up of stratiform beds, usually <50 m thick and composed of stacked <1 to 5 m thick parasequences or evaporite cycles, with a variably-present restricted-marine carbonate unit at a cycle base (Table 1). Salts were deposited as mixed evaporitic mudflat and saltern evaporites, sometimes with local accumulations of bittern salts. Typically, platform salts were deposited in laterally extensive (>50-100 km wide), hydrographically-isolated, subsealevel marine-seepage lagoons (salterns) or evaporitic mudflats (sabkhas and salinas). These regions had no same-scale modern coun-terparts and extended as widespread depositional sheets across large portions of hydrographically isolated marine platform areas, which passed seaward across a subaerial seepage barrier into open marine sediments (Figure 5). In marine margin epeiric settings, such as the Jurassic Arab/Hith and Permian Khuff cycles of the Middle East or the Cretaceous Ferry Lake Anhydrite in the Gulf of Mexico, these platform evaporites are intercalated with shoalwater marine-influenced carbonate shelf/ramp sediments, which in turn pass basinward across a subaerial sill into open marine carbonates. Landward they pass into arid zone continental siliciclastics or carbonate mudflats.


Platform evaporite deposition occurred in both pericontinental and epicontinental settings, at times of low-amplitude 4th and 5th order sealevel changes, which typify greenhouse eustasy (Figure 5; Warren, 2010). Platform evaporites also typify the saline stages of some intracratonic basins. Platform evaporites cannot form in the high-amplitude, high-frequency sealevel changes of icehouse eustasy. The 100m+ amplitude oscillations of icehouse times mean sealevel falls off the shelf edge every 100,000 years, so any evaporite that had formed on the platform is subaerially exposed and leached. Fourth order high-amplitude icehouse eustatic cycles also tend to prevent laterally-continuous carbonate sediment barriers forming at the top of the shelf to slope break and so icehouse evaporite systems tend not to be hydrographically isolated (drawdown) at the platform scale. Rather icehouse eustasy favours nonmarine evaporites as the dominant style, along with small ephemeral marine-margin salt bodies, as seen today in the bedded Holocene halites and gypsums of Lake Macleod in coastal West Australia.

Ancient platform evaporite successions may contain halite beds, especially in intracratonic basinwide settings, but periplatform settings, outside of intracratonic basins, are typically dominated by 5–40 m thick Ca-sulphate beds intercalated with normal-marine platform carbonates (Table 1). The lateral extent of these epeiric platform sulphate bodies, like the Middle Anhydrite Member of the Permian Khuff Fm. of Saudi Arabia and the UAE, with a current area of more than 1,206,700 km2, constitute some of the most aerially-extensive evaporite beds ever deposited.


Basinwide evaporites

Are made up of thick evaporite units >50–100 m thick made up of varying combinations of deepwater and shallow wa-]ter evaporites (Table 2). They retain textural evidence of different but synchronous local depositional settings, including mudflat, saltern, slope and basin (Figures 6). When basinwide evaporite deposition occurs, the whole basin hydrology is evaporitic, holomictic, and typically saturated with the same mineral phase across vast areas of the basin floor, as in the Dead Sea basin today. The Dead Sea has a more limited lateral scale than ancient basinwides but currently has halite forming simultaneously as; 1) decimeter-thick chevron-dominated beds on the saline-pan floor of the shallow parts around the basin edge in waters typically less than 1-10 metres deep, and 2) as coarse inclusion-poor crystal meshworks of halite on the deep basin floor that sits below a halite-saturated brine column up to hundreds of metres deep (Figure 3a). Ancient basinwide successions are usually dominated by thick massive salt beds, generally more than 100-500 m thick. Deposits are made up of stacked thick halite beds, but can also contain substantial volumes of thick-bedded Ca-sulphate and evaporitic carbonate, as in the intracratonic basinwide accumulations of the Delaware and Otto Fiord Basins (Table 2).

Owing to inherent purity and thickness of the deposited halite, many halite-dominant basinwide beds are also remobilized, via loading or tectonics, into various halokinetic geometries (Hudec and Jackson, 2007). Some basinwide systems (mostly marine-fed intracratonic settings) entrain significant accumulations of marine-fed potash salts, as in the Devoni-an Prairie Evaporite of western Canada. In contrast, all Quaternary examples of commercial potash deposits are accumulating in continental lacustrine systems (Warren 2016; Chapter 11).

Basinwide evaporite deposits are the result of a combination of tectonic and hydrological circumstances that are not currently active on the world’s surface (Figure 4b). They were last active in the Late Miocene (Messinian), in association with soft-suture collision basins tied to the Alpine-Himalaya orogenic belt, and in Middle Miocene (Badenian) basins developed in the early rift stages of the Red Sea. Basinwide systems will be active again in the future at sites and times of appropriate plate-plate interaction, when two continental plate edges are nearby, and the intervening seafloor is in or near a plate-edge rift or suture and is both subsealevel and hydrographically isolated (Figure 6).


Lacustrine (nonmarine) evaporites

Quaternary continental playa/lacustrine are constructed of stratiform salt units, with the greater volume of saline sediment accumulating in lower, more-saline portions of the lacustrine landscape. Beds are usually dominated by nodular gypsum and displacive halite, deposited in extensive evaporitic mudflats and saltpans with textures heavily overprinted by capillary wicking, rather than as bedded bottom-nucleated layers on the subaqueous floors of perennial brine lakes (Figure 3b; Ruch et al., 2012). In ancient counterparts, the total saline lacustrine thickness ranged from meters to hundreds of meters, with lateral extents measured in tens to hundreds of kilometres (Table 3; Figure 4a). Lacustrine salt beds are separated vertically, and usually surrounded by, deposits of lacustrine muds, alluvial fans, ephemeral streams, sheet floods, eolian sands, and redbeds. As today, ancient lacustrine salts accumulated in endorheic or highly restricted discharge basins, with perennial saline water masses tending to occur in the drainage sumps of steep-sided drainage basins (Warren, 2010, 2016). Saline lake basins accumulating gypsum, or more saline salts like halite or glauberite, typically have a shallow water table in peripheral saline mudflat areas and so are dominated by continental sabkha textures. Nearby is the lowermost part of the lacustrine depression or sump where deposition is typified by ephemeral ponded brine pan deposits, rather than permanent saline waters.

Saline lacustrine mineralogies depend on compositions of inflowing waters, so depositional sumps in regions with non-marine ionic proportions in the feeder inflow, accumulate thick sequences of nonmarine bedded salts dominated by trona, glauberite, and thenardite. In contrast, nonmarine areas with thalassic (seawater-like) inflows tend to accumulate more typical sequences of halite, gypsum, and anhydrite.


Across the Quaternary, less-saline perennial saline-lake beds tend to occur during more humid climate periods in the same continental-lacustrine depressions where saline-pan beds form (e.g., Lake Magadi, Great Salt Lake, Lake Urmia). On a smaller scale, in some modern saline lake basins, parts of the lake floor can be permanently located below the water surface (Northern Basin in the Dead Sea or Lake Asal). In some modern saline sumps dominated by mudflats, a perennial saline lake water mass is located toward the edge of a more central salt-flat zone, forming a perennial water filled “moat” facies surrounding a seasonally desiccated saline pan (as in Salar, de Atacama, Salar de Uyuni, Lake Magadi, Lake Natron). These permanent to near-permanent saline water “moat” regions are typically created where fresher inflows encounter saltier beds of the lake centre, dissolve them, and so form water-filled peripheral depressions. Bottom sediment in the moats tend to be mesohaline carbonate laminites, which can contain TOC levels as high as 12%.

High-water stage perennial saline lacustrine sediments tend to be carbonate-rich or silica-rich (diatomaceous) laminites. Ancient examples of large saline lacustrine deposits made up of alternating humid and desiccated lacustrine units include the Eocene Green River Formation of Wyoming and the Permian Pingdiquan Formation of the Junggar Basin, Chi-na (Table 4). Evaporites deposited in a suprasealevel lacustrine basin (especially Neogene deposits) have numerous same-scale Quaternary analogues, unlike the more voluminous ancient marine platform and basinwide evaporites (Figure 4a).


Salt punches above its weight, but why? (facilitator for economic accumulations of oil, gas and metals)

In terms of total mass in sedimentary basins, the proportion of evaporite across the world’s Phanerozoic basins is rarely more than 2% (Figure 7; Ronov, 1980). Today we have comprehensive documentation that salt horizons, their brines, associated dissolution and alteration conduits control significant economic associations of oil, gas and metals (Warren, 2016):

  • 50% of world’s carbonate reservoirs (seals, traps and source rocks)
  • All the world’s supergiant oil and gas fields in thrusts (seals and structural traps).
  • All supergiant sedimentary copper deposits (halokinetic brine focus)
  • 50% of world’s giant SedEx deposits (halokinetic brine focus).
  • 80% of giant MVT deposits (sulphate-fixer & brine interface)
  • World’s largest Phanerozoic Ni deposit (meta-igneous – Noril’sk).
  • Many larger IOCG deposits (meta-evaporite, brine and hydrothermal).
  • This enrichment runs counter to a proportion of 2% of the world's Phanerozoic sediments. The exact why or how of these associations is still not well understood. Most geologists working with oil, gas or metal buildups in a salt-rich basin will come to have a suspicion, and for some, the conviction, that salt or its subsurface alteration plays a role in defining the position or enrichment level of the commodity of interest. Evaporite masses in the subsurface, especially if halite-dominant, enable both physical and chemical alterations, which tend to improve economic prospectivity. The unique properties of salt in the diagenetic realm tend to facilitate, focus and stabilise processes that lead to elevated levels of accumulations of various commodities (Figure 8). That is, evaporites in a basin tend to enhance the volume of oil, gas or metals in an accumulation, but are not necessarily the direct cause of the accumulation/precipitation.


    Evaporite-hydrocarbon association

    When the reservoired hydrocarbon below a salt seal is oil, little or no leakage can take place through a laterally continuous evaporite. Even when the reservoired hydrocarbon is methane, little or no loss occurs, even by diffusion (Ehgartner et al., 1998). The much greater efficiency of evaporite seals, compared to shales, is clearly seen in the total hydrocarbon volumes held back by the two lithologies. Total worldwide shale sediment volume in the Earth's sedimentary crust is more that an order of magnitude greater than that of evaporites, yet the split between reservoired hydrocarbons below a shale or an evaporite seal is roughly 50:50 (Grunau, 1987). Likewise, typical volumes of salt-sealed giant fields dis-covered in the last decade are much grander than that held in mudrock-sealed systems (Bai and Yu, 2014).

    Of the 120 giant oil and gas fields discovered in the period 2000 - 2012 some 54.6 % are hosted in marine carbonates and 12% in lacustrine carbonates, meaning less than a third of new giant discoveries are in siliciclastic reservoirs (Figure 9a). Some 56% of these oil and gas giants have an evaporite seal, with 82% of the marine carbonates having an evaporite seal and 91% of the lacustrine carbonates having an evaporite seal (Figure 9b, c). Clearly, carbonate reservoirs with evaporite seals constitute most of the giant oil and gas discoveries across the period 2000-2012, and the proportions of this association are likely to increase in conventional discoveries across the next decades (Bai and Yu, 2014).


    Of these 120 oil and gas giants, one is a megagiant field with a recoverable reserve of 50 Bboe (Billion barrels oil equivalent) or more, and 6 are supergiant fields with recoverable reserves of 5 Bboe or more. The seven largest fields are: Galkynysh (aka Yolotan or Osman) gas field in the Ama-Darya Basin of Turkmenistan with an original 2P reserve of 67.1 Bboe, making it likely the second largest known gas field in the world; Kashagan oil field (18.1 Bboe) in the North Caspian basin; Kish 2 gas field in the Arabian Basin (Iran): Lula (previously known as Tupi), Franco and Libra oil fields in the Santos Basin of offshore Brazil; and Sulige gas field (5.7 Bboe) in the Ordos Basin, China. In this listing, the geology of Galkynysh is not yet reliably published, but of the remaining 6, only Sulige does not have an evaporite association.

    Typical bedded evaporite seals, especially beds composed of monomineralogic assemblages such as massive nodular anhydrite or massive halite and have measured entry pressures more than 3000 psi. Most impure evaporite beds have entry pressures greater than 1000 psi, as do many evaporite-plugged reflux dolomites. Contrast this with most shales, which tend to be water-bearing (mostly bound and structural water), with typical entry pressures between 900 and 1500 psi (Sneider et al., 1997). Although such shales are respectable seals, over time shale allows substantial diffu-sive leakage of methane and even liquid hydrocarbons via inherent microporosity, less so if the shales are organic-rich. Even ignoring halite's ability to reanneal and flow under stress, any evaporite seal has much lower intrinsic permeability than shale, and this helps maintain its seal integrity. With permeabilities of 10-7 md, a hydraulic gradient of 0.01, and a porosity less than  0.01, a brine would take somewhere between 3 and 30 million years to flow 1 metre into an unfractured halite seal. For anhydrite, which has permeability some 100 times higher than halite, a brine would take between 30,000 and 300,000 years to flow 1 metre into the seal (Beauheim and Roberts, 2002).


    Evaporites are excellent longterm seals to substantial hydrocarbon columns. This applies to any siliciclastic or carbonate reservoir adjacent to bedded or halokinetic salty seals (Figure 10). The exemplary ability of bedded evaporites to act as seals holding back massive hydrocarbon columns is clearly seen in the Middle East, where Ghawar, the world’s largest oil field, is sealed by bedded anhydrites of the Jurassic Arab D Formation and the overlying Hith Anhydrite seal. This evaporite seal association in Ghawar holds back an estimated remaining reserve of more than 100-200 billion barrels. Bedded platform anhydrites also seal Safaniya, the world’s largest offshore field, also in Saudi Arabia, with estimated reserves of more than 25-30 billion barrels of oil and 5 billion cubic feet of natural gas. Likewise, Permian platform anhydrites are the regional seal to North Field in offshore Qatar, the world’s largest single gas field (non-associated gas) with more than 500 tcf of reserves (Alsharhan and Nairn, 1997). This gas is reservoired and sealed in the evaporitic dolomites of the Permian Khuff Formation and recent announcements by the Qatari Government have postulated more than 900 tcf of certified non-associated gas sealed in the North Field structure.

    In terms of physical processes, buried near-pure halite beds and masses (rock salt) in the sedimentary realm are rheologically unusual compared to other nearby sediments, in that at geological time scales rock salt can set up deformation responses that mimic Newtonian fluid responses. At the same time, intercrystalline textures in the flowing salt mass maintain seal integrity, even as local crystals dissolve and reprecipitate. That is, down to depths of 6-8 km rock salt flows and maintains seal integrity, while adjacent non-salt sediments tend to fault and fracture. The ability of bedded salt to seal (Figure 10a) and of flowing salt to create and seal supra-, intra- and sub-salt reservoirs is well documented (Figure 10b).

    Dense evaporitic brines can pass through adjacent or underlying sediments both at the time the bedded salts are accumulating, or later as a subsurface salt mass dissolves. For example, chemical responses, created within the set-up hydrology of widespread salt deposition and early burial, drive brine reflux, moving magnesium-rich modified seawater brines into and through the underlying carbonate sediments. This hydrology crafts broad reflux dolomite haloes, with local burial anhydrite patches and overall improves intercrystalline connectivity, at least until the dolomite reservoir becomes overdolomitised (Figure 11). Once this happens, all effective polyhedral porosity is lost in the dolomite and the reservoir potential is destroyed. Later dissolution brines can create hydrothermal waters capable of leaching and burial dolomitization, as in the burial-dolomite reservoirs of the Western Canada Sedimentary Basin (Davies and Smith, 2006).


    Organic-hydrocarbon association

    As long ago as the middle of last century, Weeks (1961) emphasised the importance of evaporites as seals to many of the world's major hydrocarbon accumulations. He also pointed out that many of the cycles of deposition that involve organic-rich carbonate marls or muds also end with evaporites. Clearly, in evaporitic settings, there is an association with Type I-II hydrogen-prone kerogens in mesohaline source rocks, and this is related to the ability of halotolerant photosynthetic algae and cyanobacteria to flourish in periodically mesohaline waters. Such kerogens tend to be oil-prone rather than gas prone and typified by long-chain hydrocarbons (Warren, 2011).

    Much of the mesohaline organic matter preserved in evaporitic carbonates, and the resulting source rocks, originated as planktonic blooms (pelagic “rain from heaven”) or from the benthic biomass (“in situ” accumulations). Such organics typically settled out as seriate pulses of organic matter (often pelleted and laminated) that sank to the bottom of a layered brine column. Each pulse was tied to a short period when surface brines were diluted and halotolerant producers (mostly cyanobacteria and algae) flourished in the freshened lit zone. That is, laminated mesohaline mudstones that constitute most evaporitic source rocks reflect biological responses to conditions of “feast or famine” in vacillatingly-layered brine bodies (Figure 12; Warren, 1986, 2011).


    Worldwide, studies of schizohaline evaporitic basins have shown that organic-rich mesohaline sediments can accumulate beneath ephemeral surface brines in salterns, or in basin and slope settings in both marine and continental settings (Figure 12; Kirkland and Evans, 1981; Oehler, 1984; Warren, 1986, 2011; Rouchy, 1988; Busson, 1992). The most prolific accumulations of organics in ancient evaporitic settings tend to be laminated micritic carbonates deposited be-neath intermittently stratified moderately-saline (mesohaline) anoxic water columns of varying brine depth.

    There are three, possibly four, major mesohaline density-stratified settings where organic-rich laminites (source rocks) accumulate in saline environments that are also associated with, or evolve into, evaporite deposits (Warren, 2011):

    1) Basin-centre lows in marine-fed evaporitic drawdown basins (basinwide salts).

    2) Mesohaline intrashelf lows atop epeiric evaporitic platforms.

    3) Saline-bottomed lows in perennial underfilled saline lacustrine basins.

    4) Closed seafloor depressions in halokinetic deepwater marine slope and rise terrains.

    The Metal-evaporite association

    In addition to the evaporite hydrocarbon association, there is an association of evaporite settings with the larger of the known MVT, Sedex, Stratiform Sedimentary Copper and some IOCG associations (Figure 13; see Warren 2016, Chapter 15 and 16 for detailed case histories and models). The role of evaporites in focusing metalliferous ore accumulations is two-fold; 1) In solution (halite-dominant precursor) they can act as chloride-rich metal carriers and 2) Locally, as beds or masses (especially of CaSO4), their dissolution products, especially if trapped, can supply sulphur (mostly as bacteriogenic or thermogenic H2S) and also set up chemical interfaces that act as foci in the setup of brine mixing conditions suitable for precipitation of metal sulphides or native elements. Hence, most evaporite-associated ore systems tend to be epigenetic, rather than syngenetic. Subsurface salt beds and masses are merely the solid part of a large ionic recycling system; dissolved metals are another part, and zones of mixing between the two are typically sites where ores tend to accumulate. Halokinesis steadies the position of a redox interface, tied to a salt dissolution brine halo, and enables an extended phase of focus to a metal precipitative (redox) interface at a stable location in subsurface earth-space (Warren, 2016).


    At the world-scale, evaporite-associated metalliferous 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 so 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 touchdowns can supply sulphur, via bacterial or thermochemical sulphate reduction, while simultaneously focusing 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 within 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.

    If salt is more than a seal, then......

    From a time in the 1950's and 1960's when evaporites were mostly seen as seals, knowledge systems developed over the next five decades now allow us to do much more with respect to predictive industrial geology, centred on the following evaporite facts:

    1) The distribution of evaporite depositional textures maps paleotopography across the underlying carbonates and siliciclastics. So, depositional signatures preserved in the seal can be used to map reservoir quality trends in terms of reflux dolomite intensity, anhydrite cement-patch distributions and zones of subaerial diagenesis

    2) Thick beds of halite tend to accumulate in particular plate tectonic settings. Beds deposited in plate-edge saline sumps tend to flow via sediment loading and extension, without later superimposed tectonic stresses. In contrast, bed deposited in intracratonic settings tend to require later externally-imposed tectonic stress in order to flow. Hydrocarbon trapping geometries are tied to particular styles of salt tectonics

    3) Mesohaline source rock distribution is related to evaporite basin architecture and likely fluid escape pathways, which in turn are related to seal type and timing of salt dissolution or halokinetic withdrawal.

    4) Potash ore quality controls are related to the timing of various brine crossflows generated during deposition, mesogenesis and telogenesis.

    5) Positions of likely base metal and copper accumulations relate to set-ups of dissolution-related of redox interfaces, mesogenetic cross-flows, and in some cases, halokinetic geometries. At the plate tectonic scale, these accumulations occur at particular hydrological interfaces within the basin architecture.

    References

    Alsharhan, A. S., and A. E. M. Nairn, 1997, Sedimentary Basins and Petroleum Geology of the Middle East: Amsterdam, The Netherlands, Elsevier Science B. V., 942 p.

    Bai, G., and Y. Xu, 2014, Giant fields retain dominance in reserves growth: Oil & Gas Journal, v. 112, p. 44-.

    Busson, G., and et al., 1992, Basins paleogenes saliferes de l'Est de la France (Valence, Bresse et Haute-Alsace) Translated Title: Paleogene salt basins of eastern France, Valence, Bresse and Haute-Alsace: Geologie de la France, v. 1, p. 15-64.

    Davies, G. R., and L. B. Smith, 2006, Structurally controlled hydrothermal dolomite reservoir facies: An overview: Bulle-tin American Association Petroleum Geologists, v. 90, p. 1641-1690.

    Ehgartner, B. L., J. T. Neal, and T. E. Hinkebein, 1998, Gas Releases from Salt: SAND98-1354, Sandia National Laborato-ries, Albuquerque, NM, June 1998.

    Gertman, I., and A. Hecht, 2002, The Dead Sea hydrography from 1992 to 2000: Journal of Marine Systems, v. 35, p. 169-181.

    Grunau, H. R., 1987, A worldwide look at the cap-rock problem: Journal of Petroleum Geology, v. 10, p. 245-266.

    Hay, W. W., A. Migdisov, A. N. Balukhovsky, C. N. Wold, S. Flogel, and E. Soding, 2006, Evaporites and the salinity of the ocean during the Phanerozoic: Implications for climate, ocean circulation and life: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 240, p. 3-46.

    Jiang, L., C. F. Cai, R. H. Worden, K. K. Li, and L. Xiang, 2013, Reflux dolomitization of the Upper Permian Changxing Formation and the Lower Triassic Feixianguan Formation, NE Sichuan Basin, China: Geofluids, v. 13, p. 232-245.

    Kirkland, D. W., and R. Evans, 1981, Source-rock potential of an evaporitic environment: Bulletin American Association of Petroleum Geologists, v. 65, p. 181-190.

    Oehler, J. H., 1984, Carbonate source rocks in the Jurassic Smackover trend of Mississippi, Alabama, and Florida, in J. G. Palacas, ed., Petroleum geochemistry and source rock potential of carbonate rocks, v. 18: Tulsa, Oklahoma, American Association of Petroleum Geologists, Studies in Geology, p. 63-69.

    Pilcher, R. S., B. Kilsdonk, and J. Trude, 2011, Primary basins and their boundaries in the deep-water northern Gulf of Mexico: Origin, trap types, and petroleum system implications: Bulletin American Association Petroleum Geologists, v. 95, p. 219-240.

    Ronov, A. B., V. E. Khain, A. N. Balukhovsky, and K. B. Seslavinsky, 1980, Quantitative analysis of Phanerozoic sedimen-tation: Sedimentary Geology, v. 25, p. 311-325.

    Rouchy, J. M., 1988, Relations évaporites-hydrocarbures:l'association laminites-récifes-évaporites dans le Messinien de Mediterranée et ses enseignements, in G. Busson, ed., Evaporites et hydrocarbures, v. 55, Mémoires du Muséum na-tional d'Historie naturelle, (C), p. 15-18.

    Ruch, J., J. K. Warren, F. Risacher, T. R. Walter, and R. Lanari, 2012, Salt lake deformation detected from space: Earth and Planetary Science Letters, v. 331-332, p. 120-127.

    Sneider, R. M., J. S. Sneider, G. W. Bolger, and J. W. Neasham, 1997, Comparison of Seal Capacity Determinations: Con-ventional Cores vs. Cuttings, in R. C. Surdam, ed., AAPG Memoir 67: Seals, Traps, and the Petroleum System.

    Warren, J. K., 1986, Shallow water evaporitic environments and their source rock potential: Journal Sedimentary Pe-trology, v. 56, p. 442-454.

    Warren, J. K., 2010, Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine depos-its: Earth-Science Reviews, v. 98, p. 217-268.

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

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

    Weeks, L. G., 1961, Chapter 5, Origin, migration and occurrence of petroleum, in G. B. Moody, ed., Petroleum Explora-tion Handbook: New York, McGraw-Hill.

     

    Evaporites and climate: Part 2 of 2 - Ancient evaporites

    John Warren - Saturday, February 25, 2017

    Introduction

    Evaporites, along with coal and bauxites, are sediments considered to be climate sensitive. Ancient evaporite distribution and associated paleolatitudes are used to reconstruct the distribution of the world's arid belts across time. As we saw in Part 1 (Salty Matters, Tuesday, January 31, 2017), thick, widespread evaporite deposit are essentially a result of the atmospheric circulation of the Hadley cells. That is, locations of subtropical dry zones and tropical/subtropical deserts of the globe are mostly determined by the positions of subsiding branches of cool, dry descending air in a Hadley cells (aka Trade Wind Belt; Lu et al., 2007; Crowley and North, 1991) within low-lying regions tied a sufficient supply of mother brine. Thus, climate plus ongoing brine supply are the underlying factors controlling locales of significant evaporite deposition (Ziegler et al., 1981). The previous article focused on regional and local climatic controls across the Quaternary (Salty Matters, Tuesday, January 31, 2017). This article extends the time frame for the evaporite/climate association across the Phanerozoic and into the Precambrian.


    As we move back in time, we move out of an icehouse-dominated world climate, with permanent ice caps waxing and waning at the world's poles, into greenhouse-dominant world climates. In greenhouse times there are no permanent polar ice sheet and glaciers occurred only in some high-altitude mountainous belts (Figure 1a, b). The transition into greenhouse climate changes the dominant 4th-order eustatic style from 100m amplitude changes every 100,000 years or so into 4th-order responses with much lower 3-5 m amplitude oscillations every 100,000 years (Figure 2a, b). Lack of polar caps raises world sealevel on the order of 40-50 metres. Thus, even without tectonic considerations, there is less continental freeboard in greenhouse times and increased the potential for significant cratonic coverage by epicontinental and pericontinental seaways (Warren, 2016; Chapter 5).


    Looking back in time, beyond the last few million years of the Quaternary, means our conceptual models must encompass a broader range of tectonic settings as well as changes in the rates of seafloor spreading, supercontinents, and times of significant igneous outpourings (superplumes). All these additional world-scale variables across a longer set of available time explain the greater range of climate potentials on the earth, compared to anything that has occurred in the time-limited base seen in icehouse-dominant Quaternary climate spreads. The last few million years of the current icehouse mode is inhabited by the human species or its primate ancestors. The previous Icehouse dominant mode was in the Carboniferous-Mid Permian when the dominant land animals were amphibians and primitive reptiles.
    One the first questions this broader ancient climate spectrum, tied to evaporites over deep time, offers up is; "How do the positions of Hadley cells vary across geological time frames?" In part 1, we saw how the rise of the Himalayas deflected a belt of cool, dry descending air much further south toward the Equator. Moving back in time creates a broader scaffolding for documenting climate variation, in part driven by the rise and fall of mountain ranges, but also influenced by increases in the rate of seafloor spreading driving ocean basin shallowing and by changes in atmospheric/seawater compositions and temperatures.


    Hadley cells and latitudinal variability over time
    According to Chen Xu et al., 2013 and Boucot et al., 2013, much of the world-scale Phanerozoic distribution of significant bedded evaporite accumulations indicates the ongoing presence of two mid-latitude arid belts, presumably situated beneath Hadley Cells. There are exceptions in locales generated in local rain shadows with orographic control provided by neighbouring mountain ranges. However, in the later Permian, through the Triassic, and much of the Jurassic the two formerly mid-latitude Hadley Cells merged over the more central Pangaeanic regions of Africa, Europe and the adjacent Americas, to form an arid belt that also encompassed low-latitude, equatorial arid regions (Figure 3). Across the continental interior of the Pangean supercontinent, this arid to hyperarid equatorial belt in the supercontinent interior prevented the formation of climate-sensitive sediments that are more typical of humid equatorial conditions that deposit coals, kaolinites, lateritic materials and bauxites. However, in this same time interval, these more humid sediment products are typically present at low latitudes of these time slices adjacent to the Panthalassic ocean.
    That is, in the absence of an equator-spanning supercontinent, low latitudes, typically imply humid and non-seasonal tropical conditions throughout much of the Phanerozoic, as we see today. But the assembly of the Pangaean supercontinent disrupted this latitudinally-zoned atmospheric circulation, replacing it with a progressively more monsoonal (seasonal) circulation and more arid, at times hyperarid, conditions in the equatorial continental interior of Pangaea (Parrish, 1993). The Pangean supercontinent reached its maximum areal extent in the Triassic and was associated with what is known as the Pangaean Megamonsoon. There were immense arid regions across the interior regions of the supercontinent that were nearly uninhabitable, with scorching days and frigid nights. However, Panthalassian coasts still experienced seasonality, transitioning from rainy weather in the summer to dry conditions during the winter and the associated accumulation of humid sediments (Figure 3b' Boucot et al., 2013). Megamonsoon aridity is evidenced not just in the accumulation of low-latitude bedded evaporite deposits. Low latitude continental aridity also drove the accumulation of thick, widespread low-latitude desert redbeds, sourced by eolian, not fluvial, detrital transport (Sweet at al., 2013) and the precipitation of bedded salt crusts in ephemeral saline lakes under exceptionally-high surface temperatures of up to 73°C (Zambito and Benison, 2013).
    Paleolatitude reconstructions Chen-Xu et al. (2013) show these continental interior arid belts in low latitude tropical-subtropical regions persisted from the Permian to the Early Cretaceous. Reunion of the humid regions from both sides of Pangaea by the early Late Cretaceous formed a through-going low latitude humid tropical-subtropical belt. This coincides with the disaggregation of Pangaean supercontinent, as the initial stages of a modern latitudinal climate belt distribution pattern emerged, tied to latitudinally-restricted evaporites that continue to the present (Figure 4).


    Tectonism and eustacy in arid climates drive the formation of mega-evaporite basins
    Within this Phanerozoic climatic framework, there are times when significant volumes of evaporites form what are know as saline giants, or megahalite/megasulphates deposits. These massive accumulations of salts formed beneath arid climates that can span both greenhouse and icehouse climates (Figure 1; Warren, 2010, 2016). Ancient marine saline giants (megahalites and megasulphates) accrued in either of two plate-scale settings, which at times merged into one another, namely; 1) Platform evaporites (Figure 5) and, 2) Basinwide evaporites (Figure 6).

    The first major contrast with nonmarine continental dominance in Quaternary evaporite settings is the fact that platform evaporites require greenhouse eustasy a and a marine feed, the second is that basinwide evaporites require tectonically- and hydrographically-isolated widespread subsealevel depressions, typically found along plate edges with continent-continent proximity in regions with a marine seepage feed and/or periodic marine overflows (Figure 6). Neither platform or basinwide conditions are present on the current earth surface. For basinwides, suitable hydrologic conditions were last present during the Messinian Salinity Crisis in the Mediterranean region, and platform evaporite settings were last present on earth across large parts of the Middle East carbonate platform during the Eocene (Tables 1, 2). There is a third group of ancient evaporite deposits; it encompasses all nonmarine lacustrine beds past and present (Table 3). This group has same-scale modern-ancient counterparts, unlike ancient marine platform and basinwide evaporites (Warren, 2010, 2016). Interestingly, the lacustrine depositional style for bedded salt accumulation dominates in the icehouse climate that is the Quaternary, and so biases a strictly uniformitarian view of the past with respect to the relative proportions of nonmarine versus marine evaporite volumes (see Part 1; Salty Matters, Tuesday, January 31, 2017).


    Platform evaporites
    Are made up of stratiform beds, usually <50 m thick and composed of stacked <1 to 5 m thick parasequences or evaporite cycles, with a variably-present restricted-marine carbonate unit at a cycle base (Table 1). Salts were deposited as mixed evaporitic mudflat and saltern evaporites, sometimes with local accumulations of bittern salts. Typically, platform salts were deposited in laterally extensive (>50-100 km wide), hydrographically-isolated, subsealevel marine-seepage lagoons (salterns) and evaporitic mudflats (sabkhas and salinas). These regions have no same-scale modern counterparts and extended as widespread depositional sheets across large portions of hydrographically isolated marine platform areas that passed seaward across a subaerial seepage barrier into open marine sediments (Figure 5). In marine margin epeiric settings, such as the Jurassic Arab/Hith and Permian Khuff cycles of the Middle East or the Cretaceous Ferry Lake Anhydrite in the Gulf of Mexico, these platform evaporites are intercalated with shoalwater marine-influenced carbonate shelf/ramp sediments, which in turn pass basinward across a subaerial sill into open marine carbonates. Landward they pass into arid zone continental siliciclastics or carbonate mudflats.

     

    Platform evaporite deposition occurred in both pericontinental and epicontinental settings, at times of low-amplitude 4th and 5th order sealevel changes, which typify greenhouse eustasy (Figure 5; Warren, 2010). Platform evaporites also typify the saline stages of some intracratonic basins. Platform evaporites cannot form in the high-amplitude, high-frequency sealevel changes of Icehouse eustasy. The 100m+ amplitude oscillations of Icehouse times mean sealevel falls off the shelf edge every 100,000 years, so any evaporite that had formed on the platform is subaerially exposed and leached. Fourth order high-amplitude icehouse eustatic cycles also tend to prevent laterally-continuous carbonate sediment barriers forming at the top of the shelf to slope break, and so icehouse evaporite systems tend not to be hydro-graphically isolated (drawdown) at the platform scale. Rather icehouse eustasy favours nonmarine evaporites as the dominant style, along with small ephemeral marine-margin salt bodies, as seen today in the bedded Holocene halites and gypsums of Lake Macleod in coastal West Australia (Part 1; Salty Matters, Tuesday, January 31, 2017).
    Ancient platform evaporite successions may contain halite beds, especially in intracratonic basinwide settings, but periplatform settings, outside of intracratonic basins, are typically dominated by 5–40 m thick Ca-sulphate beds intercalated with normal-marine platform carbonates (Table 1). The lateral extent of these epeiric platform sulphate bodies, like the Middle Anhydrite Member of the Permian Khuff Fm. of Saudi Arabia and the UAE, with a current area of more than 1,206,700 sq. km., constitute some of the most aerially-extensive evaporite beds ever deposited.


    Basinwide evaporites
    Are made up of thick evaporite units >50–100 m thick made up of varying combinations of deepwater and shallow water evaporites (Figure 1; Table 2). They retain textural evidence of different but synchronous local depositional settings, including mudflat, saltern, slope and basin (Figure 6). When basinwide evaporite deposition occurs, the whole basin hydrology is evaporitic, holomictic, and typically saturated with the same mineral phase across vast areas of the basin floor, as seen on a much smaller scale today in the Dead Sea basin. The Dead Sea currently has halite forming simultaneously as; 1) decimeter-thick chevron-dominated beds on the saline-pan floor of the shallow parts around the basin edge in waters typically less than 1-10 metres deep, and 2) as coarse inclusion-poor crystal meshworks of halite on the deep basin floor that sits below a halite-saturated brine column up to hundreds of metres deep. Ancient basinwide successions are usually dominated by thick massive salt beds, generally more than 100-500 m thick. Deposits are made up of stacked thick halite beds, but can also contain substantial volumes of thick-bedded Ca-sulphate and evaporitic carbonate, as in the intracratonic basinwide accumulations of the Delaware and Otto Fiord Ba-sins (Table 2).
    Owing to inherent purity and thickness of the deposited halite, many halite-dominant basinwide beds are also remobilized, via loading or tectonics, into various halokinetic geometries (Hudec and Jackson, 2007). Some basinwide systems (mostly marine-fed intracratonic settings) entrain significant accumulations of marine-fed potash salts, as in the Devonian Prairie Evaporite of western Canada. In contrast, all Quaternary examples of commercial potash deposits are accumulating in continental lacustrine systems (Warren 2016; Chapter 11).

     

    Basinwide evaporite deposits are the result of a combination of tectonic and hydrological circumstances that are not currently active on the world’s surface (Figure 1). They were last active in the Late Miocene (Messinian), in association with soft-suture collision basins tied to the Alpine-Himalaya orogenic belt, and in Middle Miocene (Badenian) basins developed in the early rift stages of the Red Sea. Basinwide systems will be active again in the future at sites and times of appropriate plate-plate interaction, when two continental plate edges are nearby, and the intervening seafloor is in or near a plate-edge rift or suture and is both subsealevel and hydrographically isolated (Figure 6). Unlike most platform evaporites, basinwides do not require greenhouse eustacy, only the appropriate association of arid climate and tectonics. The latter sets up a deep hydrographically-isolated subsealevel tectonic depression with a geohydrology that can draw on a huge reserve of marine mother brine in the nearby ocean. For this reason, saline giants tend to form at times of plate-scale continent-continent proximity and so occur mostly in craton-margin settings.

     

    Lacustrine (nonmarine) evaporites
    Quaternary continental playa/lacustrine are constructed of stratiform salt units, with the greater volume of saline sediment accumulating in lower, more-saline portions of the lacustrine landscape. Beds are usually dominated by nodular gypsum and displacive halite, deposited in extensive evaporitic mudflats and saltpans with textures heavily overprinted by capillary wicking, rather than as bedded bottom-nucleated layers on the subaqueous floors of perennial brine lakes (Ruch et al., 2012). In ancient counterparts, the total saline lacustrine thickness ranged from meters to hundreds of meters, with lateral extents measured in tens to hundreds of kilometres (Table 3). Lacustrine salt beds are separated vertically, and usually surrounded by, deposits of lacustrine muds, alluvial fans, ephemeral streams, sheet floods, eolian sands, and redbeds. As today, ancient lacustrine salts accumulated in endorheic or highly restricted discharge basins, with perennial saline water masses tending to occur in the drainage sumps of steep-sided drainage basins (Warren, 2010, 2016). Saline lake basins accumulating gypsum, or more saline salts like halite or glauberite, typically have a shallow water table in peripheral saline mudflat areas and so are dominated by continental sabkha textures. Nearby is the lowermost part of the lacustrine depression or sump where deposition is typified by ephemeral ponded brine pan deposits, rather than permanent saline waters.
    Saline lacustrine mineralogies depend on compositions of inflowing waters, so depositional sumps in regions with non-marine ionic proportions in the feeder inflow, accumulate thick sequences of nonmarine bedded salts dominated by trona, glauberite, and thenardite. In contrast, nonmarine areas with thalassic (seawater-like) inflows tend to accumulate more typical sequences of halite, gypsum, and anhydrite.
    Across the Pliocene-Quaternary icehouse, less-saline perennial saline-lake beds tend to occur during more humid climate periods in the same continental-lacustrine depressions where saline-pan beds form (e.g., Lake Magadi, Great Salt Lake, Lake Urmia). On a smaller scale, in some modern saline lake basins, parts of the lake floor can be permanently located below the water surface (Northern Basin in the Dead Sea or Lake Asal). In some modern saline sumps dominated by mudflats, a perennial saline lake water mass is located toward the edge of a more central salt-flat zone, forming a perennial water filled “moat” facies surrounding a seasonally desiccated saline pan (as in Salar, de Atacama, Salar de Uyuni, Lake Magadi, Lake Natron). These permanent to near-permanent saline water “moat” regions are typically created where fresher inflows encounter saltier beds of the lake centre, dissolve them, and so form water-filled peripheral depressions. Bottom sediment in the moats tends to be mesohaline carbonate laminites, which can contain TOC levels as high as 12%.
    High-water stage perennial saline lacustrine sediments tend to be carbonate-rich or silica-rich (diatomaceous) laminites. Ancient examples of large saline lacustrine deposits made up of alternating humid and desiccated lacustrine units include the Eocene Green River Formation of Wyoming and the Permian Pingdiquan Formation of the Junggar Basin, China (Table 4). Evaporites deposited in a suprasealevel lacustrine basin (especially Neogene deposits) have numerous same-scale Quaternary analogues, unlike the more voluminous ancient marine platform and basinwide evaporites (Figure 7). Clearly, across the Quaternary, saline continent lacustrine settings possess areas of bedded salt accumulation that are far greater than those of any contemporaneous marine-fed salt sumps (Part 1; Salty Matters, Tuesday, January 31, 2017). But in ancient climes, especially during in the continental interior of the Pangean supercontinent (mid Permian to Triassic), regions of continental interior sabkhas and saline pans had areas far greater than any seen in Quaternary continental saline sumps (Zambia and Benison, 2013).


    Is the present-day climate the key to evaporite understanding?
    This short answer is yes, Hadley Cells across the Phanerozoic are mostly tied to climate belts that maintain sub-tropical positions, but to this notion, we must add a geological context. Today we are living in an icehouse climate mode and have been for the last 12 Ma. It is tied to the presence of polar ice-sheets that wax and wane over 100,000-year time frames, so moving the position of the Hadley cells and changing the intensity of atmospheric circulation. In this icehouse climate, large eustatically-controlled marine-fed evaporite deposits are not preserved, as sea level falls off the continental shelf every 100,000 years or so. The world's largest bedded salt deposits formed sometime in the last 2 million years, are found in continental interiors, typically in endorheic tectonic sumps in either hot arid or steppe climate settings, often with salt diapirs outcropping or subcropping in the drainage basin and the basin floor can be located at elevations well above sealevel (Part 1; Salty Matters, Tuesday, January 31, 2017).
    As we move back into much of Phanerozoic time, we see world climates dominated by greenhouse modes, with shorter episodes of polar ice-sheets and icehouse climates in the Carboniferous-Early Permian ≈ 50 million years long, and the Ordovician, some 15 million years long (Figure 1). Greenhouse climate lacks permanent polar ice sheets, so sea-levels are higher, and 4th-order eustatic amplitudes in sea level are much less (a few meters versus hundred meters plus changes). Greenhouse sets up epeiric and intercontinental seaways that when hydrographically isolated, but still marine-fed, can deposit huge areas of platform evaporites centred in isolated seepage-fed sub-sealevel sumps. These platform deposits can also form outside of Greenhouse times in marine-fed tectonically-induced intracratonic sumps.
    Basinwide evaporite deposits span icehouse and greenhouse mode arid belts, whenever a marine-fed subsealevel tectonic sump forms at positions of continent-continental proximity in an arid belt. Across much of the Phanerozoic, basinwide deposits typically accumulated beneath subtropical belts of cool, dry descending air set up in a Hadley cell, and so are located north and south of a tropical equatorial belt. But the accretion of the Pangaean supercontinent (Carboniferous to Jurassic) set up conditions of continentality and orographic shadowing that allowed an arid saline belt to span the hyperarid interior of the supercontinent.
     
    References

    Boucot, A., Chen-Xu, and C. Scotese, 2013, Phanerozoic Paleoclimate: An Atlas of Lithologic Indicators of Climate: Concepts in Sedimentology and Paleontology, v. 11: Tulsa, OK, SEPM, 32 p.

    Chen-Xu, A. J. Boucot, C. R. Scotese, F. Junxuan, W. Yuan, and Z. Xiujuan, 2012, Pangaean aggregation and disaggregation with evidence from global climate belts: Journal of Palaeogeography, v. 1, p. 5-13.

    Crowley, T. J., and G. R. North, 1991, Paleoclimatology: New York, Oxford University Press, 339 p.

    Hudec, M. R., and M. P. A. Jackson, 2007, Terra infirma: Understanding salt tectonics: Earth-Science Reviews, v. 82, p. 1-28.

    Lu, J., G. A. Vecchi, and T. Reichler, 2007, Expansion of the Hadley cell under global warming: Geophysical Research Letters, v. 34.

    Parrish, J. T., 1993, Climate of the Supercontinent Pangea: Journal of Geology, v. 10.

    Ruch, J., J. K. Warren, F. Risacher, T. R. Walter, and R. Lanari, 2012, Salt lake deformation detected from space: Earth and Planetary Science Letters, v. 331-332, p. 120-127.

    Sweet, A. C., G. S. Soreghan, D. E. Sweet, M. J. Soreghan, and A. S. Madden, 2013, Permian dust in Oklahoma: Source and origin for Middle Permian (Flowerpot-Blaine) redbeds in Western Tropical Pangaea: Sedimentary Geology, v. 284–285, p. 181-196.

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

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

    Zambito, J. J., and K. C. Benison, 2013, Extremely high temperatures and paleoclimate trends recorded in Permian ephemeral lake halite: Geology, v. 41, p. 587-590.

    Ziegler, A. M., S. F. Barrett, C. R. Scotese, and B. W. Sellwood, 1981, Palaeoclimate, Sedimentation and Continental Accretion [and Discussion]: Philosophical Transactions of the Royal Society of London. Series A, Mathematical and Physical Sciences, v. 301, p. 253-264.


     

     

    Evaporites and climate: Part 1 of 2 - Are modern deserts the key?

    John Warren - Tuesday, January 31, 2017

    Salt deposits and deserts

    Much of the geological literature presumes that thick sequences of bedded Phanerozoic evaporites accumulated in hot arid zones tied to the distribution of the world’s deserts beneath regions of descending air within Hadley Cells in a latitudinal belt that is typically located 15 to 45 degrees north or south of the equator (Figure 1a: Gordon, 1975). As this sinking cool air mass approaches the landsurface beneath the descending arm of a Hadley Cell it warms, and so its moisture-carrying capacity increases. The next two articles will discuss the validity of this assumption of evaporites tying to hot arid desert belts in the trade wind belts, first, by a consideration of actual Quaternary evaporite distributions as plotted in a GIS database with modern climate overlays, then in the second article via a look at ancient salt/climate distributions.

    Significant volumes of Quaternary evaporite salts are normally interpreted as being allied to the distribution of the world’s hot arid deserts (Figure 1b). In a general way this is true, but, as Warren (2010) shows, the correlation is an oversimplification. A hot arid desert does not necessarily equate to occurrences of laterally extensive bedded evaporites; there must also be a significant long-term brine inflow to the evaporative sump, incoming waters may be meteoric, marine, a hybrid and perhaps the sump is fed brines coming from dissolution of earlier formed salts in the drainage basin, including diapiric salt (Table 1; Warren, 2016).


    Actually, there are different ways of defining a desert and by implication its associated evaporites. One accepted approach is to define desert as a terrestrial area receiving less that 250 mm (10 inches) of annual precipitation. Using this definition some 26.2% of the world’s landsurface is desert (Figure 1b). But, in terms of evaporite distribution and the economics of the associated salts, this climatic generalization related to annual rainfall conceals three significant hydrological truisms. All three need to be met in order to accumulate thick sequences of bedded salts (Warren, 2010): 1) For any substantial volume of evaporite to precipitate and be preserved, there must be a sufficient volume of cations and anions in the inflow waters to allow thick sequences of salts to form; 2) The depositional setting and its climate must be located within a longer term basin hydrology that favours preservation of the bedded salt, so the accumulating salt mass can pass into the burial realm; 3) There must be a negative water balance in the basin with the potential for more water to leave the local hydrological sump than enter. When using a rainfall (precipitation) based definition of desert, the significance of these three simple hydrological axioms and the consequences, as to where bedded evaporites accumulate, is lost in the generalization that “evaporites form in the world’s deserts.”


    Continental-interior evaporites

    In an evaporite context it is better to define and plot saline hydrologies within a climatic framework where deserts are given the same hydrological consideration that is required for evaporite salts to form. That is, a desert is an area of land where annual precipitation (inflow) is less than potential evapotranspiration (outflow). This definition of a desert is the one used by Köppen (1900). With slight modification, his climatological scheme is still in widespread use to breakout the various climatic zones across the world’s landsurface (Kottek et al., 2006). Using a Köppen climate base, figure 2 plots worldwide occurrences of modern saline depositional systems with areas greater than 250 km2. Table 1 compares characteristics of some of the larger Quaternary bedded evaporite settings in marine edge and continental interiors. These regions contain evaporite salts accumulating in saline soils, sabkhas, salinas, saline lakes, playas and salt flats, with textural forms ranging from; isolated crystals and nodules in a terrigenous matrix, to salt crusts, to stacked beds of salts that can be more than 10 metres thick. The majority of the plotted saline areas are in arid regions, as defined by Köppen (Zone B), but not all such areas of widespread salts are in deserts (as defined by Köppen) and not all are in hot climates.


    The range of large (>250 km2) saline systems in the world’s arid landscape is more climatically diverse than just evaporite occurrences within a hot arid desert (BWh), although such associations do constitute some 38% of saline occurrences (Figure 3a; Warren, 2010; 2016). Cold arid deserts (BWk) host 23% of large saline occurrences, making a combined total of 61% for large evaporite accumulation (area >250 km2) occurrences in modern arid deserts (BW group), while the arid steppes (BSh and BSk) host another 22%. In total the world’s arid climatic zones host 83% of today’s larger evaporite occurrences (Figure 3b).


    This leaves another substantial, but not widely recognized, climate zone where significant volumes of Quaternary evaporites can accumulate, this is the polar tundra (ET); an environment where some 11% of large evaporite areas occur. In terms of evaporite volumes, the polar tundra (ET) is typically an arid high altitude belt, mostly in the Horse Latitude (Trade Wind) belts, and not located in polar or near-polar higher latitudes. The lakes and saline pans of the high plateaus of the Andes (Altiplano) and the Himalayas (Tibetan Plateau) typify this style of tundra (ET) evaporite. Water may be commonplace in the ET zone, but is there mostly as ice, and cryogenic salts are commonplace (see Salty Matters, Feb 24, 2015). The remaining region where significant evaporite volumes are found, some 6% of the total of large saline occurrences is a group of deposits defined by continental interior snow climates (group D), some with hot dry summers with solar evaporites alternating with dry winters favouring the possible accumulation of cryogenic salts (e.g. Great Salt Lake, USA).


    In the Northern Hemisphere the occurrence of large evaporite systems within arid deserts and steppe climates (BW and BS settings) extends much further south toward the equator and much further poleward (from 5-55°N) than the narrower range of large evaporite occurrences and associated climates in the southern hemisphere (Figure 4). This hemispheric asymmetry in evaporite occurrence is mostly a response to world-scale adiabatic effects associated with the collision of India with Eurasia and growth of the Himalayas. Today, a Cainozoic mountain range, centred on the Himalyas, diverts world-scale atmospheric air flows from the more north-south trajectory, usually associated with Hadley Cell circulation. For example, the Kunlun Mountains, first formed some 5.3 Ma, prevents moisture from the Indian Monsoon reaching much of the adjacent Tibet Plateau. Its adiabatic rain shadow creates the Taklamakan desert, the second largest active sand desert in the world (BWk).

    The uplift of the Himalayas also creates a dry easterly jet stream, moving arid cool air across the Tibet Plateau, around the northern side of the Himalayas, and then equatorward across the Arabian Peninsula toward Somalia where it descends and gains heat. That is, this stream of cool southwesterly-flowing dry air warms as it moves across the Eastern Mediterranean land areas and so heightens existing aridity. This helps create an adiabatic desert zone that today ranges across Arabia and northern Africa almost to the Equator (Figures 5).


    In the southern hemisphere, the uplift of the Andes has formed high intermontane depressions and the allied adiabatic aridity that are cooler with lower evaporation rates and higher relied in the immediate basin compared to groundwater depressions in flatter lower-elevation continental interior deserts like the Sahara. This higher stability hydrology favours salars over dry mudflats, as typified by Salar de Atacama and Salar de Uyuni. Atacama has a Quaternary saline sediment fill made up of a more than 900 m thickness of interlayered salt and clay, while Uyuni holds a more than 120 m thick interval of interbedded salt and clay infill, with areas of 3,064 km2 and 9,654 km2 and elevations of 2250 m and 3650 m, respectively (Figure 7a). These salars are the two largest known examples of Quaternary bedded halite accumulation, worldwide. Yet neither resides in hot arid desert settings (BWh); both are located in cold arid deserts (BWk) and in actively subsiding, high altitude (>2500m) intermontane (high relief) endorheic depressions.


    Worldwide, distribution of most of the larger (>250 km2) Quaternary evaporite settings located in hot arid (BWh) desert settings, tie either to; 1) endorheic river terminations along desert margins, especially if adjacent to mountain belts (e.g. the various circum Saharan chotts, playas and sabkhas adjacent to Atlas Mountains), or 2) to ancient inherited paleodrainage depressions (as in the majority of the interior salt lakes of Australia or the southern Africa pans). Another hot arid desert (BWh) evaporite association is defined by termination outflow rims of deep artesian systems, as in Lake Eyre North, Australia (8,528 km2, with an ephemeral halite crust up to 2 m thick in its southern portion). Similar, deeply-circulating, meteoric artesian hydrologies help explain the distribution of chotts in the BWh zone of NE Africa. Unlike Atacama and Uyuni in the Andes, none of these modern BWh artesian systems preserve stacked decametre-thick salt beds, nor did they do so at any time in the Quaternary. Rather, the most extensive style of BWh salt in meteoric-fed artesian outflow zones is as dispersed crystals of gypsum and halite in a terrigenous redbed matrix (sabkha) or as visually impressive large ephemeral saline flats and pans covered by metre-scale salt crusts that dissolve and reform with the occasional decadal freshwater flood (Warren 2016; Chapter 3). Sediments of such continental groundwater outflow zones are typically reworked by eolian processes and, due to a lack of long term watertable stability, the longterm sediment fill is matrix-rich and evaporite-poor, with the Quaternary sediment column typified by episodes of deflation, driven by 10,000-100,000 year cycles of glacial-interglacial climate changes.

    It seems that to form and preserve laterally-extensive decametre-thick stacked beds of halite in a Quaternary time-framework requires an actively subsiding tectonic depression in a cooler high-altitude continental desert, where temperatures and evaporation rates are somewhat lower than in BWh settings, allowing brine to pond and remain at or near the surface for longer periods (Figures 6, 7a). But perhaps more importantly, all of the larger regions of the Quaternary world, where thick stacked bedded (not dispersed) evaporites are accumulating, are located in continental regions with drainage hinterlands where dissolution of older halokinetic marine-fed salt masses are actively supplying substantial volumes of brine to the near surface hydrology. This halokinetic-supplied set of deposits includes, Salar de Uyuni and Salar de Atacama in the Andean Altiplano, the Kavir salt lakes of Iran, the Qaidam depression of China and the Dead Sea (Table 1).


    Marine-edge evaporites

    Thick sequences of stacked Quaternary evaporite beds with a marine-brine feed are far less common and far smaller than meteoric/halokinetic Quaternary continental evaporite occurrences. Relatively few marine-fed evaporite regions exist today with areas in excess of 250 km2 (Table 1; Figures 6, 7b). By definition, in order to be able to draw on significant volumes of seawater, these basins must operate with a subsealevel hydrology. This allows large volumes of seawater to seep into the depression and evaporate. It also means most of these deposits are located near the continental edge where a freestanding mass of seawater is not too distant

    The largest known deposit of this group is Lake Macleod on the west coast of Australia, with an area of 2,067 km2 and containing a 10m-thick Holocene gypsum/halite bed (Figure 7b). It hosts a saltworks producing some 1,500,000 tonnes/year of halite from lake brines in a BWh setting (Warren, 2016). A smaller marine seepage example, with a similar Quaternary coastal carbonate dune-hosted seepage hydrology, is Lake MacDonnell near the head of the Great Australia Bight. It has an area of 451 km2, a 10m-thick fill of Holocene bedded gypsum and is located in a milder BSk setting, compared to Lake Macleod. Even so, annually, the Lake MacDonnell operation is quarrying more than 1.4 million tonnes of Holocene coarsely-crystalline near-pure gypsum and producing more that 35,000 tonnes of salt via by pan evaporation of lake brines.

    Interestingly, when the climatic settings of Holocene coastal salinas of southern and western Australia are compared, all show similar interdunal sump seepage hydrologies with unconfined calcarenite aquifers, yet it is clear that gypsum evaporite beds dominate in BSk and lower precipitation levels, with more carbonate, in Csb coastal settings typified by hot dry summers. Halite dominates the marine-fed bedded fills in BWh coastal settings, while Coorong-style meteoric-fed carbonates dominate in similar interdunal coastal seepage depressions in the more the humid and somewhat cooler Csb settings of the Coorong region.


    One of the most visually impressive marine seep systems in the world is Lake Asal, immediately inland of the coast of Djibouti, with an area of only 54 km2 it is much smaller than the 250 km2 cutoff used for this discussion. It is located in a BWh climate, similar to that of the Danakil depression, contains subaqueous textured gypsum and pan halites, and lies at the bottom of a hydrographically-isolated basalt-floored depression with a brine lake surface some 115m below sealevel (Figure 8; see Warren, 2016 Chapter 4 for geological detail). This difference in elevation between the nearby Red Sea and the lake floor drives a marine seep hydrology, so that seawater-derived groundwater escapes as springs along the basaltic lake margin. Most of the cations and anions in the seawater feed, that ultimately accumulate as salts in the subsealevel Asal sump, move lakeward via gravitational seepage along fractures in a basaltic ridge aquifer separating the Red Sea from the brine lake, in what is an actively extending continental rift that will shortly become and arm of the Red Sea.

    That there are very few large marine-fed bedded Quaternary evaporite systems is illustrated by summing the total area of large active continental-fed evaporite areas (>250km2) listed in Figure 2, which gives a total surface area worldwide in excess of 360,000 km2. In contrast, the total area of large modern marine-fed bedded salt systems, at slightly less 10,000 km2, is more than an order of magnitude smaller. This low value for large coastal marine-fed evaporite occurrences is in part because many classic coastal sabkhas, with characteristic dispersed evaporites in their supratidal sediments and long considered to be archetypal marine-fed groundwater evaporite system, are now seen as mostly continental brine-fed hydrologies.

    For example, the modern Abu Dhabi sabkha system (1,658 km2; BWh) has been shown by Wood (2010) to be a continental groundwater outflow area, where most of ions precipitating as salts in the supratidal zone are supplied by upwelling of deeply-circulated meteoric waters, not seawater flooding (Warren, 2016; Chapter 3). Similar continental hydrologies, resurging into the coastal zone. supply much of the salt accumulating in the suprasealevel sabkhas near Khobar in Saudi Arabia (BWh), Rann of Kutch in India (28,000 km2, BWh), Sabkha Matti (2,955km2, BWh) in the Emirates, and Sabkha de Ndrhamcha (634 km2, BWh) in Mauritania. And yet today all of these large sabkha occurrences are located just inland of modern coastal zones and so, without hydrological knowledge, are easily interpreted as marine. The distinction between a water budget and a salt supply budget emphasizes a general observation in Holocene evaporite systems that, in order to define the main fluid carrier for the salt volume found in any supratidal part of an evaporitic depression, it is important to understand and quantify the nature of the groundwater feed, be it by marine or nonmarine, in both continental-interior and marine-margin settings. It also means that surface runoff and seawater washovers, while at times visually impressive, typically do not supply the greater volume of preserved salts in most modern coastal sabkha systems.

    Due to the highly impervious nature of muddy sabkha sediment, an occasional seawater storm flood into a coastal margin mudflat does not mean the pooled seawater ever penetrates the underlying sabkha. Most of its solute load is deposited as an ephemeral surface crust of halite that is a few centimetres thick atop the sabkha muds (Wood et al., 2005). Based on Wood’s and other hydrological studies of modern coastal sabkhas in BWh settings, centred on the Middle East, it seems that the salt supply to most modern coastal sabkha depressions is still not at hydrological equilibrium with the present sealevel, which was reached at the beginning of the Holocene some 6,000-8,000 years ago. That is, modern zones of continental groundwater outflow along an arid coast can have watertables that are up to a metre of two above sealevel and are strongly influenced by the resurgence of seaward-flowing deeply circulating, continental groundwaters (it is axiomatic that surface runoff and local meteoric recharge tend to be limited in BWh settings). In this situation resurging waters in the sabkha tend to have a dominantly continental ionic supply for their preserved salts.

    Lastly, in terms of the climatic and hydrological pre-requisites for the accumulations of thick bedded Holocene salt deposits, one must recognize that marine-margin sabkha mudflats do not have the same geohydrology as a near-coastal, hydrographically-isolated seepage-fed sub-sealevel salina depression. A BWh or BSk near-coastal salina is slowly but continually resupplied ions via springs fed by the effusion of seawater, moving under a gravity drive (evaporative drawdown), through an aquifer that is the barrier separating the salina depression from the adjacent ocean. Until the depositional surface reaches hydrological equilibrium, it draws continually on the near limitless supply of ions in the adjacent salty reservoir that is the world’s ocean. Salts growing displacively in the supratidal parts of a sabkha cannot call on a drawdown hydrology and can only be supplied marine salts from salt spray as it is blown inland or from waters washed over the sabkha by the occasional storm driven wash-over and breakout (Figure 9; Warren and Kendall, 1985).


    So what?

    So, if we based our ideas of where ancient bedded evaporites formed on a strictly uniformitarian approach using Quaternary analogues, then the study of where bedded salts have formed best in the world over the last two million years leads to a conclusion that large bedded salt deposits are not marine fed. Rather, in the Quaternary, thick bedded salts form best in tectonically active, subsiding hypersaline groundwater sumps located in high-altitude high-relief cold-arid deserts. These depressions are hydrographically closed, with the purer, most voluminous, examples of 100m-thick evaporite successions located in tectonically active piggy-back depressions, with lake floors and hydrologies that are more than 2000 m above sea level. Their hydrologies are endorheic and strongly influenced by deeply circulating meteoric waters. In addition, the better examples of thick stacked bedded salts are supplied ions via the dissolution of older marine halokinetic salts in the surrounding drainage basin (Table 1).

    Yet, anyone working in ancient (pre-Quaternary) evaporites knows from simple arguments of salt volumetrics versus brine sources, and the nature of the enclosing sediments, that they are part of a dominantly marine basin fill. This will be the focus

    References

    Gordon, W. A., 1975, Distribution by latitude of Phanerozoic evaporite deposits: Journal of Geology, v. 83, p. 671-684.

    Köppen, W., 1900, Versuch einer Klassification der Klimate, vorzsugsweise nach ihren Beziehungen zur Pflanzenwelt: Geographraphische Zeitschrift, v. 6, p. 593-611.

    Kottek, M., J. Grieser, C. Beck, B. Rudolf, and F. Rubel, 2006, World Map of the Köppen-Geiger climate classification updated: Meteorologische Zeitschrift, v. 15, p. 259-263.

    Stieljes, L., 1973, Evolution tectonique récente du rift d'Asal: Review Geographie Physical Geologie Dynamique, v. 15, p. 425-436.

    Warren, J. K., 2010, Evaporitic source rocks: mesohaline responses to cycles of “famine or feast” in layered brines, Quaternary carbonate and evaporite sedimentary facies and their ancient analogues, John Wiley & Sons Ltd., p. 315-392.

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

    Warren, J. K., and C. G. S. C. Kendall, 1985, Comparison of sequences formed in marine sabkha (subaerial) and salina (subaqueous) settings; modern and ancient: Bulletin American Association of Petroleum Geologists, v. 69, p. 1013-1023.

    Wood, W. W., 2010, An historical odyssey: the origin of solutes in the coastal sabkha of Abu Dhabi, United Arab Emirates, Quaternary carbonate and evaporite sedimentary facies and their ancient analogues, John Wiley & Sons Ltd., p. 243-254.

    Wood, W. W., W. E. Sanford, and S. K. Frape, 2005, Chemical openness and potential for misinterpretation of the solute environment of coastal sabkhat: Chemical Geology, v. 215, p. 361-372.



    [1] The word desert comes from the Latin dēsertum (originally "an abandoned place"), a participle of dēserere, "to abandon."

     

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

    John Warren - Saturday, December 31, 2016

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


    Extrasalt gas and brine intersections

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

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

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

     

    Caprock and other salt periphery-held gases

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

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

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

    Intrasalt gas

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

    Gas types and sources at the local and basin scale

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

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

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

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

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

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

    Summary

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

    References

    Baar, C. A., 1977, Applied salt-rock mechanics; 1, The in-situ behavior of salt rocks: Developments in geotechnical engineering. 16a.

    Clayton, C. J., S. J. Hay, S. A. Baylis, and B. Dipper, 1997, Alteration of natural gas during leakage from a North Sea salt diapir field: Marine Geology, v. 137, p. 69-80.

    Dusseault, M. B., V. Maury, F. Sanfilippo, and F. J. Santarelli, 2004, Drilling around salt: Stresses, Risks, Uncertainties: Gulf Rocks 2004, In 6th North America Rock Mechanics Syposium (NARMS), Houston Texas, 5-9 June 2004, ARMA/NARMS 04-647.

    Fischer, M., R. Botz, M. Schmidt, K. Rockenbauch, D. Garbe-Schönberg, J. Glodny, P. Gerling, and R. Littke, 2006, Origins of CO2 in Permian carbonate reservoir rocks (Zechstein, Ca2) of the NW-German Basin (Lower Saxony): Chemical Geology, v. 227, p. 184-213.

    Gropp, 1919, Gas deposits in potash mines in the years 1907-1917 (in German): Kali and Steinsalz, v. 13, p. 33-42, 70-76.

    Löffler, J., 1962, Die Kali- und Steinsalzlagerstätten des Zechsteins in der Dueutschen Deomokratischen Republik, Sachsen: Anhalt. Freiberg. Forschungsh C, v. 97, p. 347p.

    Mikhail, S., and D. A. Sverjensky, 2014, Nitrogen speciation in upper mantle fluids and the origin of Earth's nitrogen-rich atmosphere: Nature Geoscience, v. 7, p. 816-819.

    Schofield, N., I. Alsop, J. Warren, J. R. Underhill, R. Lehné, W. Beer, and V. Lukas, 2014, Mobilizing salt: Magma-salt interactions: Geology, v. 42, p. 599-602.

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

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

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

     

     


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