Salty Matters

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Brine evolution and origins of potash: primary or secondary? SOP in Quaternary saline lakes: Part 2 of 2

John Warren - Friday, November 30, 2018

 


Introduction

This, the second in this series of articles on potash brine evolution deals with production of sulphate of potash in plants that exploit saline hydrologies hosted in Quaternary saline sumps. There are two settings where significant volumes of sulphate of potash salts are economically produced at the current time; the Ogden salt pans on the northeast shore of the Great Salt Lake in Utah and Lop Nur in China. Although potassium sulphate salts precipitate if modern seawater is evaporated to the bittern stage, as yet there is no operational SOP plant utilising seawater. This is due to concurrent elevated levels of magnesium and chlorine in the bittern, a combination that favours the precipitation of carnallite concurrently with the precipitation of double sulphate salts, such as kainite (Figure 1). Until now, this makes the processing of the multi-mineralogic precipitate for a pure SOP product too expensive when utilising a marine brine feed.

 

Potash in Great Salt Lake, USA (SOP evolution with backreactions)

Today sulphate of potash fertiliser is produced via a combination of solar evaporation and brine processing, using current waters of the Great Salt Lake, Utah, as the brine feed into the Ogden salt pans, which are located at the northeastern end of the Great Salt Lake depression in Utah (Figure 2a). A simpler anthropogenic muriate of potash (MOP) brine evolution occurs in the nearby Wendover salt pans on the Bonneville salt flats. There, MOP precipitates as sylvinite in concentrator pans (after halite). The Bonneville region has a bittern hydrochemistry not unlike like the evolved Na-Cl brines of Salar de Atacama, as documented in the previous article, but it is a brinefield feed without the elevated levels of lithium seen in the Andean playa (Figure 3b).

Great Salt Lake brine contains abundant sulphate with levels sufficiently above calcium that sulphate continues to concentrate after most of the Ca has been used up in the precipitation of both aragonite and gypsum. Thus, as the brines in the anthropogenic pans at Ogden approach the bittern (post halite) stage, a series of sulphate double salts precipitate (Figure 4), along with carnallite and sylvite.


Great Salt Lake brines

The ionic proportions in the primary brine feed that is the endorheic Great Salt Lake water depends on a combination of; 1) the inflow volumes from three major rivers draining the ranges to the east, 2) groundwater inflow, 3) basin evaporation, and 4) precipitation (rainfall/snowfall) directly on the lake (Jones et al., 2009). Major solute inputs can be attributed to calcium bicarbonate-type river waters mixing with sodium chloride-type springs, which are in part hydrothermal and part peripheral recycling agents for NaCl held in the lake sediments. Spencer et al. (1985a) noted that prior to 1930, the lake concentration inversely tracked lake volume, which reflected climatic variation in the drainage. However, since that time, salt precipitation, primarily halite and mirabilite, and dissolution have periodically modified lake brine chemistry and led to density stratification and the formation of brine pockets of different composition.

Complicating these processes is repeated fractional crystallisation and re-solution (backreaction) of lake mineral precipitates. The construction of a railway causeway has restricted circulation, nearly isolating the northern from the southern part of the lake, which receives over 95% of the inflow. Given that Great Salt Lake waters are dominated by Na and Cl, this has led to halite precipitation in the north (Figures 2a, 3a; Gwynn, 2002). Widespread halite precipitation also occurred before 1959, especially in the southern area of the lake, associated with the most severe droughts (Jones et al., 2009; Spencer et al., 1985a). Spencer et al. (1985a) also described the presence of a sublacustrine ridge, which probably separated the lake into two basins at very low lake stands in the Quaternary. Fluctuating conditions emphasise brine differentiation, mixing, and fractional precipitation of salts as significant factors in solute evolution, especially as sinks for CaCO3, Mg, and K in the lake waters and sediments. The evolution of these brine/rock system depends on the concentration gradient and types of suspended and bottom clays, especially in relatively shallow systems.


Brine evolution across the Ogden pans

Figure 3a plots the known hydrochemistry of the inflow waters to the Great Salt Lake and their subsequent concentration. Evolving lake waters are always Na-Cl dominant, with sulphate in excess of magnesium in excess of potassium, throughout. Any post-halite evaporite minerals from this set of chemical proportions will contain post-halite potash bittern salts with elevated proportions of sulphate and magnesium and so will likely produce SOP rather than MOP associations. Contrast these hydrochemical proportions with the inflow and evolution chemistry in the pore brines of the Bonneville salt flat (Figure 3b) the Dead Sea and normal marine waters. Across all examples, sodium and chlorine are dominant and so halite will be the predominant salt deposited after aragonite and gypsum (Figure 4). Specifically, there are changes in sulphate levels with solar concentration (Figure 3b). In brines recovered from feeder wells in the Bonneville saltflat, unlike the nearby bajada well waters, the Bonneville salt flat brines show potassium in excess of sulphate and magnesium. In such a hydrochemical system, sylvite, as well as carnallite, are likely potassium bitterns in post-halite pans. The Wendover brine pans on the Bonneville saltflat produce MOP, not SOP, along with a MgCl2 brine, and have done so for more than 50 years (Bingham, 1980; Warren, 2016).


The mineral series in the Ogden pans

Figure 5 illustrates a laboratory-based construction of the idealised evolution of a Great Salt Lake feed brine as it passes through the various concentration pans. Figure 5 is a portion of the theoretical 25°C sulphate-potassium-magnesium phase diagram for the Great Salt Lake brine system and shows precipitates that are in equilibrium with brine at a particular concentration. Figures 4 and 5 represent typical brine concentration paths at summertime temperatures (Butts, 2002). Importantly, these figures do not describe the entire brine concentration story and local variations; mineralogical complexities in the predicted brine stream are related to thermal stratification, retention times and pond leakage. Effects on the chemistry of the brine due to the specific day-by-day and season-by-season variations of concentration and temperature, which arise in any solar ponding operation, require onsite monitoring and rectification. Such ongoing monitoring is of fundamental import when a pilot plant is constructed to test the reality of a future brine plant and its likely products.

 

Figure 6 illustrates the idealised phase evolution of pan brines at Ogden in terms of a K2SO4 phase diagram (no NaCl or KCl co-precipitates are shown; Felton et al., 2010). Great Salt Lake brine is pumped into the first set of solar ponds where evaporation initially proceeds along the line shown as Evap 1 until halite reaches saturation and is precipitated. Liquors discharged from the halite ponds are transferred to the potash precipitation ponds where solar evaporation continues as line Evap 2 on the phase diagram and potassium begins to reach saturation after about 75% of the water is removed. Potassium, sodium levels rise with further evaporation and schoenite precipitates in the schoenite crystalliser. After some schoenite precipitation occurs, the liquor continues to evaporate along the Evap 3 line to the point that schoenite, sylvite, and additional halite precipitate. Evaporation continues as shown by line Evap 3 to the point that kainite, sylvite and halite become saturated and precipitate. From this plot, the importance of the relative levels of extraction/precipitation of sulphate double salts versus chloride double salts is evident, as the evaporation plot point moves right with increasing chloride concentrations. That is, plot point follows the arrows from left to right as concentration of chloride (dominant ion in all pans) increases and moves the plot position right.


Production of SOP in the Great Salt Lake

To recover sulphate of potash commercially from pan bitterns fed from the waters of Great Salt Lake, the double salts kainite and schoenite are first precipitated and recovered in post-halite solar ponds (Figures 6, 7). The first salt to saturate and crystallise in the concentrator pans is halite. This is successively followed by epsomite, schoenite, kainite, carnallite, and finally bischofite. To produce a desirable SOP product requires ongoing in-pan monitoring and an on-site industrial plant whereby kainite is converted to schoenite. The complete salt evolution and processing plant outcome in the Ogden facility is multiproduct and can produce halite, salt cake and sulphate of potash and a MgCl2 brine product. Historically, sodium sulphate was recovered from the Great Salt Lake brines as a byproduct of the halite and potash production process, but ongoing low prices mean Na2SO4 has not been economically harvested for the last decade or so.

The complete production and processing procedure is as follows (Figure 7; Butts, 2002, 2007; Felton et al., 2010): 1) Brine is pumped from the Great Salt Lake into solar evaporation ponds where sodium chloride precipitates in the summer. 2) When winter weather cools the residual (post-halite) brine in the pans to -1 to -4°C, sodium sulphate crystals precipitate as mirabilite in a relatively pure state. Mirabilite crystals can be picked up by large earth-moving machinery and stored outdoors each winter until further processing takes place. 3) The harvested mirabilite can be added to hot water, and anhydrous sodium sulphate precipitated by the addition of sodium chloride to the heated mix to reduce sodium sulphate solubility through the common ion effect. The final salt cake product is 99.5% pure Na2SO4. 4) To produce SOP, Great Salt Lake brines are allowed to evaporate in a set of halite ponds, until approaching saturation with potassium salts. The residual brine is then transferred to a mixing pond, where it mixes with a second brine (from higher up the evaporation series, that contains a higher molar ratio of magnesium to potassium. 5) This adjusted brine is then allowed to evaporate to precipitate sodium chloride once more, until it is again saturated with respect to potassium salts. 6) The saturated brine is then transferred to another pond, is further evaporated and precipitates kainite (Figures 5, 6). Kainite precipitation continues until carnallite begins to form, at which time the brine is moved to another pond and is allowed to evaporate further to precipitate carnallite. 6) Some of the kainite-depleted brine is recycled to the downstream mixing pond to maintain the required molar ratio of magnesium to calcium in this earlier mixing pond (step 4). 7) Once carnallite has precipitated, the residual brine is transferred to deep storage and subjected to winter cooling to precipitate additional carnallite as it is a prograde salt. 8) Cryogenically precipitated carnallite can be processed to precipitate additional kainite by mixing it with a kainite-saturated brine. 9) MgCl2-rich end-brines in the post-carnallite bittern pans are then further processed to produce either MgCl2 flakes or a 32% MgCl2 brine. These end-bitterns are then used as a feedstock to make magnesium metal, bischofite flakes, dust suppressants, freeze preventers, fertiliser sprays, and used to refresh flush in ion exchange resins.

Some complexities in the observed mineral precipitation series in the Ogden Pans

Under natural solar pond conditions in the Ogden Pans, the brine temperature fluctuates with the air temperature across day-night and seasonal temperature cycles, and there is a lag time for temperature response in waters any brine pan, especially if the pan is heliothermic. Atmosphere-driven fluctuations in temperature results in changes in ion saturations, which can drive selective precipitation or dissolution of salts in the brine body. Air temperature in the Ogden pans may be 35°C during the day and 15°C at night. Brine at point A in figure 5 may favour the formation of kainite during the daytime and schoenite at night. The result of the diurnal temperature oscillation is a mixture of both salts in a single pond from the same brine. In terms of extracted product, this complicates ore processing as a single pan will contain both minerals, produced at the same curing stage, at the same time, yet one double salt entrains KCl, the other K2SO4, so additional processing is necessary to purify the product stream (Butts, 2002).

The sulphate ion in the pan waters is particularity temperature sensitive, and salts containing it in GSL pans tend to precipitate at cooler temperatures. Surficial cooling during the summer nights can cause prograde salts to precipitate, but the next day's heat generally provides sufficient activation energy to cause total dissolution of those salts precipitated just a few hours before (Butts, 2002). It is not unusual to find a 0.5 cm layer of hexahydrite (MgSO4.6H2O) at the bottom of a solar pond in the morning, but redissolved by late afternoon.

Under controlled laboratory conditions, brine collected from the hypersaline north arm of the Great Salt Lake will not crystallise mirabilite  until the brine temperature reaches 2°C or lower. Yet, in the anthropogenic solar ponds, mirabilite has been observed to crystallise at brine temperatures above 7°C. During the winter, as the surface temperature of the GSL pan brine at night becomes very cold (2°C or lower), especially on clear nights, and mirabilite rafts will form on and just below the brine surface and subsequently sink into the somewhat warmer brine at the floor of the pond. Because there is insufficient activation energy in this brine to completely redissolve the mirabilite, it remains on the pan floor, until warmer day/night temperatures are attained. However, it is also possible for salts precipitated by cooling to be later covered by salts precipitated by evaporation, which effectively prevents dissolution of those more temperature-sensitive salts that would otherwise redissolve (Butts, 2002).

There are also longer terms seasonal influences on mineralogy. Some salts deposited in June, July, and August (summer) will convert to other salts, with a possible total change in chemistry, when they are exposed to colder winter temperatures and rainfall. Kainite, for example, may convert to sylvite and epsomite and become a hardened mass on the pond floor; or if it is in contact with a sulphate-rich brine, it can convert to schoenite. Conversely, mirabilite will precipitate in the winter but redissolve during the hot summer months.

The depth of a solar pond also controls the size of the crystals produced. For example, if halite (NaCl) is precipitated in a GSL pond that is either less than 8 cm or more than 30cm deep, it will have a smaller crystal size than when precipitated in a pond between 8 and 30 cm deep. Smaller crystals of halite are undesirable in a de-icing product since a premium price is paid for larger crystals.

In terms of residence time, some salts require more time than others to crystallise in a pan. Brine that is not given sufficient time for crystallisation before it is moved into another pond, which contains brine at a different concentration, will produce a different suite of salts. For example, if a brine supersaturated in ions that will produce kainite, epsomite, and halite (reaction I), is transferred to another pond, the resulting brine mixture can favour carnallite (reaction 2), while kainite salts are eliminated.

Reaction 1: 9.75H2O + Na+ + 2Cl- + 2Mg2+ + K+ + 2SO42+ —> MgSO4.KCl +2.75H2O + MgSO4.7H2O + NaCl

Reaction 2: 12H2O + Na+ + 4Cl- + 2Mg2+ + K+ + 2SO42+ —> MgCl2.6H2O +MgSO4.6H2O + NaCl

Reaction 1 retains more magnesium as MgCl2 in the brine; reaction 2 retains more sulphate. In reaction 2, it is also interesting to note the effect of waters-of-hydration on crystallization; forcing out salts with high waters of crystallization results in higher rates of crystallization. The hydrated salts remove waters from the brine and further concentrate the brine in much the same way as does evaporation.

Pond leakage and brine capture (entrainment) in and below the pan floor are additional influences on mineralogy, regardless of brine depth or ponding area. As mentioned earlier, to precipitate bischofite and allow for MgCl2 manufacture, around ninety-eight percent of the water from present North Arm brine feed must evaporate. If pond leakage causes the level of the ponding area to drop too quickly, it becomes near impossible to reach saturation for bischofite (due to brine reflux). Control of pond leakage in the planning and construction phases is essential to assure that the precipitated salts contain the optimal quantity of the desired minerals for successful pond operation.

The opposite of leakage is brine retention in a precipitated layer; it can also alter brine chemistry and recovery economics. Brine entrained (or trapped) in the voids between salt crystals in the pond floor is effectively removed from salt production and so affects the chemistry of salts that will be precipitated as concentration proceeds and can also drive unwanted backreactions. The time required to evaporate nearly ninety percent of the water from the present north arm Great Salt Lake brine in the Ogden solar pond complex, under natural steady state conditions, is approximately eighteen months.

Summary of SOP production procedures in Great Salt Lake

Sulphate of potash cannot be obtained from the waters of the Great Salt Lake by simple solar evaporation (Behrens, 2002). As the lake water is evaporated, first halite precipitates in a relatively pure form and is harvested. By the time evaporative concentration has proceeded to the point that saturation in a potash-entraining salt occurs, most of the NaCl has precipitated. It does, however, continue to precipitate and becomes the primary contaminant in the potassium-bearing salt beds in the higher-end pans.

Brine phase chemistry from the point of potassium saturation in the evaporation series is complicated, and an array of potassium double salts are possible, depending on brine concentration, temperature and other factors. Among the variety of potash minerals precipitated in the potash harvester pans, the majority are double salts that contain atoms of both potassium and magnesium in the same molecule, They are dominated by kainite, schoenite, and carnallite. All are highly hydrated; that is, they contain high levels of water of crystallisation that must be removed during processing. SOP purification also involves removal of the considerable quantities of sodium chloride that are co-precipitated, after this the salts must be chemically converted into potassium sulphate.

Controlling the exact mineralogy of the precipitated salts and their composition mixtures is not possible in the pans, which are subject to the vagaries of climate and associated temperature variations. Many of the complex double salts precipitating in the pans are stable only under fixed physiochemical conditions, so that transitions of composition may take place in the ponds and even in the stockpile and early processing plant steps.

While weathering, draining, temperature and other factors can be controlled to a degree, it is essential that the Great Salt Lake plant be able to handle and effectively accommodate a widely variable feed mix (Behrens 2002). To do this, the plant operator has developed a basic process comprising a counter-current leach procedure for converting the potassium-bearing minerals through known mineral transition stages to a final potassium sulfate product (Figure 7). This set of processing steps is sensitive to sodium chloride content, so a supplemental flotation circuit is used to handle those harvested salts high in halite. It aims to remove the halite (in solution) and upgrade the feed stream to the point where it can be handled by the basic plant process.

Solids harvested from the potash ponds with elevated halite levels are treated with anionic flotation to remove remaining halite (Felton et al., 2010). To convert kainite into schoenite, it is necessary to mix the upgraded flotation product with a prepared brine. The conversion of schoenite to SOP at the Great Salt Lake plant requires that new MOP is added, over the amount produced from the lake brines. This additional MOP is purchased from the open market. The schoenite solids are mixed with potash in a draft tube baffle reactor to produce SOP and byproduct magnesium chloride.

The potassium sulfate processing stream defining the basic treatment process in the Great Salt Lake plant is summarised as Figure 7, whereby once obtaining the appropriate chemistry the SOP product is ultimately filtered, dried, sized and stored. Final SOP output may then be compacted, graded, and provided with additives as desired, then distributed in bulk or bagged, by rail or truck.


Lop Nur, Tarim Basin, China (SOP operation)

Sulphate of potash (SOP) via brine processing (solution mining) of lake sediments and subsequent solar concentration of brines is currently underway in the fault-bound Luobei Hollow region of the Lop Nur playa, in the southeastern part of Xinjiang Province, Western China (Liu et al., 2006; Sun et al., 2018). The recoverable sulphate of potash resource is estimated to be 36 million tonnes from lake brine (Dong et al., 2012). Lop Nur lies in the eastern part of the Taklimakan Desert (Figure 8a), China’s largest and driest desert, and is in the drainage sump of the basin, some 780 meters above sea level in a BSk climate belt. The Lop Nur depression first formed in the early Quaternary, due to the extensional collapse of the eastern Tarim Platform and is surrounded and typically in fault contact with the Kuruktagh (to the north), Bei Shan (to east) and Altun (to the south) mountains (Figure 8b).

The resulting Lop Nur (Lop Nor) sump is a large groundwater discharge playa that is the terminal point of China’s largest endorheic drainage system, the Tarim Basin, which occupies an area of more than 530,000 km2 (Ma et al., 2010). The Lop Nur sump is the hydrographic base level to local and regional groundwater and surface water flow systems, and thus collectively captures all river and subsurface flow originating in the surrounding mountainous regions. The area has been subject to ongoing Quaternary climate and water supply oscillations, which over the last few hundred years has driven concentric strandzone contractions on the playa surface, to form what is sometimes called the “Great Ear Lake" of the Lop Nur sump (Liu et al., 2016a).

Longer term widespread climate oscillations (thousands of years) drove precipitation of saline glauberite-polyhalite deposits, alternating with more humid lacustrine mudstones especially in fault defined grabens with the sump. For example, Liu et al. (2016b) conducted high-resolution multi-proxy analyses using materials from a well-dated pit section (YKD0301) in the centre of Lop Nur and south of the Luobei depression. They showed that Lop Nur experienced a progression through a brackish lake, saline lake, slightly brackish lake, saline lake, brackish lake, and playa in response to climatic changes over the past 9,000 years.

Presently, the Lop Nur playa lacks perennial long-term surface inflow and so is characterised by desiccated saline mudflats and polygonal salt crusts. The upward capillary flux from the shallow groundwater helps to maintain a high rate of evaporation in the depression and drives the formation of a metre-thick ephemeral halite crust that covers much of the depression (Liu et al., 2016a).

Historically, before construction of extensive irrigation systems in the upstream portion of the various riverine feeds to the depression and the diversion of water into the Tarim-Kongqi-Qargan canal, brackish floodwaters periodically accumulated in the Lop Nur depression. After the diversion of inflows, terminal desiccation led to the formation of the concentric shrinkage shorelines, that today outline the “Great Ear Lake” region of the Tarim Basin (Figure 8b; Huntington, 1907; Chao et al., 2009; Liu et al., 2016a).

The current climate is cool and extremely arid (Koeppen BSk); average annual rainfall is less than 20 mm and the average potential evaporation rates ≈3500 mm/yr (Ma et al., 2008, 2010). The mean annual air temperature is 11.6°C; higher temperatures occur during July (>40°C), and the lower temperatures occur during January (<20°C). Primary wind direction is northeast. The Lop Nor Basin experiences severe and frequent sandstorms; the region is well known for its wind-eroded features, including many layered yardangs along the northern, western and eastern margins of the Lop Nur salt plain (Lin et al., 2018).


Salinity and chemical composition of modern groundwater brine varies little in the ‘‘Great Ear” area and appears not to have changed significantly over the last decade (Ma et al., 2010). Dominant river inflows to the Lop Nor Basin are Na-Mg-Ca-SO4-Cl-HCO3 waters (Figure 9). In contrast, the sump region is characterised by highly concentrated groundwater brines (≈350 mg/l) that are rich in Na and Cl, poor in Ca and HCO3+CO3, and contain considerable amounts of Mg, SO4 and K, with pH ranging from 6.6 to 7.2 (Figure 9). When concentrated, the Luobei/Lop Nur pore brines is saturated with respect to halite, glauberite, thenardite, polyhalite and bloedite (Ma et al., 2010; Sun et al., 2018).

Groundwater brines, pooled in the northern sub-depression, mostly in the Luobei depression, are pumped into a series of pans to the immediate south, where sulphate of potash is produced via a set of solar concentrator pans. Brines in the Luobei depression and adjacent Xingqing and Tenglong platforms are similar in chemistry and salinity to the Great Ear Lake area but with a concentrated saline reserve due to the presence of a series of buried glauberite-rich beds (Figure 9; Hu and Wang, 2001; Ma et al., 2010; Sun et al., 2018).

K-rich mother brines in the Luobei hollow also contain significant MgSO4 levels and fill open phreatic pores in a widespread subsurface glauberite bed, with a potassium content of 1.4% (Liu et al., 2008; Sun et al., 2018). Feed brines are pumped from these evaporitic sediment hosts in the Luobei sump into a large field of concentrator pans to ultimately produce sulphate of potash (Figure 8a).

Brine chemical models, using current inflow water and groundwater brine chemistries and assuming open-system hydrology, show good agreement between theoretically predicted and observed minerals in upper parts of the Lop Nor Basin succession (Ma et al., 2010). However, such shallow sediment modelling does not explain the massive amounts of glauberite (Na2SO4.CaSO4) and polyhalite (K2SO4MgSO4.2CaSO4.2H2O) recovered in a 230 m deep core (ZK1200B well) from the Lop Nor Basin (Figure 9a).


Hydrochemical simulations assuming a closed system at depth and allowing brine reactions with previously formed minerals imply that widespread glauberite in the basin formed via back reactions between brine, gypsum and anhydrite and that polyhalite formed via a diagenetic reaction between brine and glauberite. Diagenetic textures related to recrystallisation and secondary replacement are seen in the ZK1200B core; they include gypsum-cored glauberite crystals and gypsum replacing glauberite. Such textures indicate significant mineral-brine interaction and backreaction during crystallisation of glauberite and polyhalite (Liu et al., 2008). Much of the glauberite dissolves to create characteristic mouldic porosity throughout the glauberite reservoir intervals (Figure 10, 11b)


Mineral assemblages predicted from the evaporation of Tarim river water match closely with natural assemblages and abundances and, in combination with a model that allows widespread backreactions, can explain the extensive glauberite deposits in the Lop Nor basin (Ma et al., 2008, 2010). It seems that the Tarim river inflows, not fault-controlled upwelling hydrothermal brines, were the dominant ion source throughout the lake history. The layered distribution of minerals in the more deeply cored sediments documents the evolving history of inflow water response to wet and dry periods in the Lop Nor basin. The occurrence of abundant glauberite and gypsum below 40 m depth, and the absence of halite, polyhalite and bloedite in the same sediment suggests that the brine underwent incomplete concentration in the wetter periods 10b).

In contrast, the increasing abundance of halite, polyhalite and bloedite in the top 40 m of core from the ZK1200B well indicate relatively dry periods (Figure 10a), where halite precipitated at lower evaporative concentrations (log Concentration factor = 3.15), while polyhalite and bloedite precipitated at higher evaporative concentrations (log = 3.31 and 3.48 respectively). Following deposition of the more saline minerals, the lake system once again became more humid in the later Holocene, until the anthropogenically-induced changes in the hydrology over the last few decades, driven by upstream water damming and extraction for agriculture (Ma et al., 2008). These changes have returned the sump hydrology to the more saline character that it had earlier in the Pleistocene.

The Lop Nur potash recovery plant/factory and pan system, located adjacent to the LuoBei depression (Figures 8, 11a), utilises a brine-well source aquifer where the potash brine is reservoired in intercrystalline and vuggy porosity in a thick stacked series of porous glauberite beds/aquifers.

Currently, 200 boreholes have been drilled in the Lop Nor brine field area showing the Late-Middle Pleistocene to Late Pleistocene strata are distributed as massive, continuous, thick layers of glauberite with well-developed intercrystal and mouldic porosity, forming storage space for potassium-rich brine (Figure 11b; Sun et al., 2018). However, buried faults and different rates of creation of fault-bound accommodation space, means there are differences in the brine storage capacity among the three brinefield extraction areas; termed the Luobei depression, the Xingqing platform and the Tenglong platform areas (Figures 9a, 11a).

In total, there are seven glauberitic brine beds defined by drill holes in the Luobei depression, including a phreatic aquifer, W1L, and six artesian aquifers, W2L, W3L, W4L, W5L, W6L, and W7 (Figure 10b; Sun et al., 2018). At present, only W1L, W2L, W3L, and W4L glauberite seams are used as brine sources. There are two artesian brine aquifers, W2X and W3X, exposed by drill holes in the Xinqing platform and there are three beds in the Tenglong extraction area, including a phreatic aquifer, W1T, and two artesian aquifers, W2T and W3T (Figure 10b).

W1L is a phreatic aquifer with layered distribution across the whole Luobei depression, with an average thickness of 17.54 m, water table depths of 1.7 to 2.3 m, porosities of 6.98% to 38.45%, and specific yields of 4.57% to 25.89%. Water yield is the highest in the central and northeast of the depression, with unit brine overflows of more than 5000 cubic meters per day per meter of water table depth (m3/dm). In the rest of the aquifer, the unit brine overflows range from 1000 to 5000 m3/dm (Sun et al., 2018). The W2L artesian aquifer is confined, nearly horizontal with a stratified distribution, and has an average thickness of 10.18 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 20 to 40 m, porosities of 4.34% to 37.8%, and specific yields of 1.08% to 21.04%. The W3L artesian aquifer is confined, with stratified distribution and an average thickness of 8.50 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 40 to 70 m, porosities of 2.85% to 19.97%, and specific yields of 1.10% to 13.37%. The W3L aquifer is also confined with stratified distribution, with an average thickness of 7.28 m, unit brine overflows of 10 to 100 m3/dm, water table depths of 70 to 100 m, porosities of 5.22% to 24.72%, and specific yields of 1.03% to 9.91%. The lithologies of the four brine storage layers are dominated by glauberite, and occasional lacustrine sedimentary clastic rocks, such as gypsum (Figure 10a).

The Xinqing platform consists of two confined potassium-bearing brine aquifers (Figure 10b). Confined brines have layered or stratified distributions. The average thicknesses of the aquifers are 4.38 to 7.52 m. Due to the F1 fault, there is no phreatic aquifer in the Xinqing platform, but this does not affect the continuity of the brine storage layer between the extraction areas. The W2X aquifer is confined, stratified, and distributed in the eastern part of this ore district with a north-south length of 77.78 km, east-west width of 16.82 km, and total area of 1100 km2. Unit brine overflows are 2.25 to 541.51 m3/dm, water table depths are 10 to 20 m, porosities are 3.89% to 40.69%, and specific yields are 2.01% to 21.15%. The W3X aquifer is also confined and stratified, with a north-south length of 76.10 km, east-west width of 18.81 km, and total area of 1444 km2. Unit brine overflows are 1.67 to 293.99 m3/d m, water table depths are 11.3 to 38 m, porosities are 4.16% to 26.43%, and specific yields are 2.11% to 14.19%23.

The Tenglong platform consists of a phreatic aquifer and two confined aquifers. W1T is a phreatic, stratified aquifer and is the main ore body, and is bound by the F3 fault (Figure 10b). It is distributed across the northern part of the Tenglong extraction area, with a north-south length of about 33 km, east-west width of about 20 km, and total area of 610 km2. Water table depths are 3.26 to 4.6 m, porosities are 2.03% to 38.81%, and specific yields are 22.48% to 1.22%. On the other side of the F3 fault, in the southern part of the mining area, is the W2T confined aquifer (Figure 10b). Water table depths are 16.91 to 22 m, porosities are 3.58% to 37.64%, and specific yields are 1.35% to 18.69%. W3T is also a confined aquifer, with a stratified orebody distributed in the southern part of the mining area, with a north-south length of about 29 km, east-west width of about 21 km, and total area of 546 km2. Water table depths are 17.13 to 47 m, porosities are 2.69% to 38.71%, and specific yields are 1.26% to 17.64%.

Lop Nur is an unusual potash source

The glauberite-hosted brinefield in the Luobei depression and the adjacent platforms makes the Lop Nur SOP system unique in that it is the world's first large-scale example of brine commercialisation for potash recovery in a Quaternary continental playa aquifer system with a non-MOP brinefield target. Elsewhere, such as in the Dead Sea and the Qarhan sump, Salar de Atacama and the Bonneville salt flats, the brines derived from Quaternary lacustrine beds and water bodies are concentrated via solar evaporation in semi-arid desert scenarios. Potash plants utilising these Quaternary evaporite-hosted lacustrine brine systems do not target potassium sulphate, but process either carnallitite or sylvinite into a commercial MOP product

 Glauberite is found in a range of other continental Quaternary evaporite deposits around the world but, as yet,                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                                       outside of Lop Nur is not economically exploited to produce sulphate of potash. For example, glauberite is a significant component in Quaternary cryogenic beds in Karabogazgol on the eastern shore of the Caspian Sea, in Quaternary evaporite beds in Laguna del Rey in Mexico, in saline lacustrine beds in the Miocene of Spain and Turkey, and in pedogenic beds in hyperarid nitrate-rich soils of the Atacama Desert of South America (Warren, 2016; Chapter 12).

In most cases, the deposits are commercially exploited as a source of sodium sulphate (salt cake). In a saline Quaternary lake in Canada, SOP is produced by processing saline lake waters. This takes place in Quill Lake, where small volumes of SOP are produced via mixing a sylvite feed (trucked into the site) with a cryogenic NaSO4 lake brine.

The Lop Nur deposit is mined by the SDIC Xinjiang Luobupo Hoevellite Co. Ltd, and the main product is potassium sulfate, with a current annual production capacity of 1.3 million tons. Pan construction began in 2000, and the plant moved in full -cale operation in 2004 when it produced ≈50,000 tons. The parent company, State Development and Investment Corporation (SDIC), is China’s largest state-owned investment holding company. The company estimates a potash reserve ≈ 12.2 billion tons in the sump. This makes Lop Nur deposit the largest SOP facility in the world, and it is now a significant supplier of high premium fertiliser to the Chinese domestic market.

Implications

A study of a few of the Quaternary pans worldwide manufacturing economic levels of potash via solar evaporation shows tha,t independent of whether SOP or MOP salts are the main product, all retain abundant evidence that salt precipitates continue to evolve as the temperature and the encasing brine chemistry change. As we shall see in many ancient examples discussed in the next article, ongoing postdepositional mineralogical alteration dominates the textural and mineralogical story in most ancient potash deposits.

As we saw in the previous article, which focused on MOP in solar concentrator plants with brine feeds from Quaternary saline lakes, SOP production from brine feeds in Quaternary saline lakes is also related strongly to cooler desert climates (Figure 12). The Koeppen climate at Lop Nur is cool arid desert (BWk), while the Great Salt Lake straddles cool arid steppe desert and a temperate climate zone, with hot dry summer zones (BSk and Csa)


Outside of these two examples, there are a number of other Quaternary potash mineral occurrences with the potential for SOP production, if a suitable brine processing stream can be devised (Warren, 2010, 2016). These sites include intermontane depressions in the high Andes in what is a high altitude polar tundra setting (Koeppen ET), none of which are commercial (Figure 12b).

Similarly, there a number of non-commercial potash (SOP) mineral and brine occurrences in various hot arid desert regions in Australia, northern Africa and the Middle East (Koeppen BWh). Today, SOP in Salar de Atacama is currently produced as a byproduct of lithium carbonate production, along with MOP, as discussed in the previous article in this series.

As for MOP, climatically, commercial potash brine SOP systems are hosted in Quaternary-age lacustrine sediments are located in cooler endorheic intermontane depressions (BWk, BSk). The association with somewhat cooler desert and less arid cool steppe climates underlines the need for greater volumes of brine to reside in the landscape in order to facilitate the production of significant volumes of potash bittern.

Put simply, in the case of both MOP and SOP production in Quaternary settings, hot arid continental deserts simply do not have enough flowable water to produce economic volumes of a chemically-suitable mother brine. That is, currently economic Quaternary MOP and SOP operations produce by pumping nonmarine pore or saline lake brines into a set of concentrator pans. Mother waters reside in hypersaline perennial lakes in steep-sided valleys or in pores in salt-entraining aquifers with dissolving salt compositions supplying  a suitable ionic proportions in the mother brine. In terms of annual volume of product sold into the world market, Quaternary brine systems supply less than 15% The remainder comes from the mining of a variety of ancient solid-state potash sources. In the third and final article in this series, we shall discuss how and why the chemistry and hydrogeology of these ancient potash sources is mostly marine-fed and somewhat different from the continental hydrologies addressed so far.

References

Behrens, P., 2002, Industrial processing of Great Salt lake Brines by Great Salt Lake Minerals and Chemical Corporation, in D. T. Gywnne, ed., Great Salt Lake: A scientific, historical and economic overview, Utah Geological and Mineral Survey, Bulletin 116, p. 223-228.

Bingham, C. P., 1980, Solar production of potash from brines of the Bonneville Salt Flats, in J. W. Gwynn, ed., Great Salt Lake; a scientific, history and economic overview. , v. 116, Bulletin Utah Geological and Mineral Survey, p. 229-242.

Butts, D., 2002, Chemistry of Great Salt Lake Brines in Solar Ponds, in D. T. Gywnne, ed., Great Salt Lake: A scientific, historical and economic overview, Utah Geological and Mineral Survey, Bulletin 116, p. 170-174.

Butts, D., 2007, Chemicals from Brines, Kirk-Othmer Encyclopedia of Chemical Technology, John Wiley & Sons, Inc., p. 784-803.

Chao, L., P. Zicheng, Y. Dong, L. Weiguo, Z. Zhaofeng, H. Jianfeng, and C. Chenlin, 2009, A lacustrine record from Lop Nur, Xinjiang, China: Implications for paleoclimate change during Late Pleistocene: Journal of Asian Earth Sciences, v. 34, p. 38-45.

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

Felton, D., J. Waters, R. Moritz, D., and T. A. Lane, 2010, Producing Sulfate of Potash from Polyhalite with Cost Estimates, Gustavson Associates, p. 19.

Hu, G., and N.-a. Wang, 2001, The sand wedge and mirabilite of the last ice age and their paleoclimatic significance in Hexi Corridor: Chinese Geographical Science, v. 11, p. 80-86.

Huntington, E., 1907, Lop-Nor. A Chinese Lake. Part 1. The Unexplored Salt Desert of Lop: Bulletin of the American Geographical Society, v. 39, p. 65-77.

Jones, B., D. Naftz, R. Spencer, and C. Oviatt, 2009, Geochemical Evolution of Great Salt Lake, Utah, USA: Aquatic Geochemistry, v. 15, p. 95-121.

Lin, Y., L. Xu, and G. Mu, 2018, Differential erosion and the formation of layered yardangs in the Loulan region (Lop Nur), eastern Tarim Basin: Aeolian Research, v. 30, p. 41-47.

Liu, C., W. Mili, J. Pengcheng, L. I. Shude, and C. Yongzhi, 2006, Features and Formation Mechanism of Faults and Potash-forming Effect in the Lop Nur Salt Lake, Xinjiang, China: Acta Geologica Sinica - English Edition, v. 80, p. 936-943.

Liu, C.-A., H. Gong, Y. Shao, Z. Yang, L. Liu, and Y. Geng, 2016a, Recognition of salt crust types by means of PolSAR to reflect the fluctuation processes of an ancient lake in Lop Nur: Remote Sensing of Environment, v. 175, p. 148-157.

Liu, C. L., M. L. Wang, P. C. Jiao, W. D. Fan, Y. Z. Chen, Z. C. Yang, and J. G. Wang, 2008, Sedimentary characteristics and origin of polyhalite in Lop Nur Salt Lake,Xinjiang: Mineral Deposits.

Liu, C. L., J. F. Zhang, P. C. Jiao, and S. Mischke, 2016b, The Holocene history of Lop Nur and its palaeoclimate implications: Quaternary Science Reviews, v. 148, p. 163-175.

Ma, C., F. Wang, Q. Cao, X. Xia, S. Li, and X. Li, 2008, Climate and environment reconstruction during the Medieval Warm Period in Lop Nur of Xinjiang, China: Chinese Science Bulletin, v. 53, p. 3016-3027.

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

Spencer, R. J., H. P. Eugster, and B. F. Jones, 1985b, Geochemistry of Great Salt Lake, Utah II: Pleistocene-Holocene evolution: Geochimica et Cosmochimica Acta, v. 49, p. 739-747.

Spencer, R. J., H. P. Eugster, B. F. Jones, and S. L. Rettig, 1985a, Geochemistry of Great Salt Lake, Utah I: Hydrochemistry since 1850: Geochimica et Cosmochimica Acta, v. 49, p. 727-737.

Sun, M.-g., and L.-c. Ma, 2018, Potassium-rich brine deposit in Lop Nor basin, Xinjiang, China: Scientific Reports, v. 8, p. 7676.

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.

 

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

John Warren - Thursday, May 31, 2018

 

Introduction

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

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

Interpreting 13C

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


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


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

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

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

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

δin = fcarbδ13Ccarb + forgδ13Corg

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


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


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

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

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


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


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

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


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

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

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

Implications for some types of 13C anomaly

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

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

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

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

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

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

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

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

 

Stable isotopes in evaporite systems: Part I: Sulphur

John Warren - Monday, April 30, 2018

 

Introduction

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

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


Sulphur isotopes across the bittern series

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

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

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

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

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

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

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

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

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

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


Ancient oceanic sulphate

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


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

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

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

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

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

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

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

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


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

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

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

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

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


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


References

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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


 


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