Salty Matters

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Polyhalite, Geological origins of an alternate low-chloride potash fertiliser

John Warren - Tuesday, July 31, 2018

Introduction

Polyhalite is the hydrated sulfate of potassium, calcium and magnesium, with the formula (K2Ca2Mg(SO4)42H2O). Polyhalite crystallises in the triclinic system, but individual euhedral crystals are very rare in nature where the usual habit is fibrous to massive. It is typically colourless, grey to white, although the natural colour in some evaporite deposits tends to brick red due to iron oxide inclusions. Primary (syndepositional) precipitates can be layered at a mm to cm scale. Its Mohs hardness is 3.5 with a specific gravity of 2.8. It was first described from a Salzburg mine in 1818 and the name comes from a Latin root that refers to the “many salts”evident in its chemical formula.

Polyhalite is relatively easy to distinguish from associated evaporites by simple field tests. Its hardness separates it from most evaporite salts, other than anhydrite. It has a bitter taste and is water soluble (incongruent dissolution), with remnants of gypsum and syngenite (K2Ca(SO4)22H2O), which is also soluble, leaving behind a final residue of gypsum (unlike sylvite and halite which dissolve wholly and congruently in fresh water). Polyhalite is not deliquescent, unlike carnallite, and gives a purple flame result when held in a gas flame due to its potassium content, unlike the non-potassium salts. Table 1 lists the constituents of the various evaporite salts mentioned in this article.


The complete equilibrium evaporation of modern seawater at 25ºC produces the mineral sequence: CaCO3 (calcite and aragonite), CaSO4 (gypsum and anhydrite), halite (NaCl), other sulphates (glauberite, polyhalite, epsomite, hexahydrite, and kieserite) and chlorides (carnallite and bischofite). Halite is the dominant mineral because of the high concentration of Na and Cl in seawater (Figure 1; after Harvie et al., 1980).

This mineral sequence with abundant sulphate bitterns reflects the relatively high concentration of sulphate in modern seawater (molar SO4 > Ca), which, following evaporation and precipitation of CaCO3 and CaSO4, produces a SO4-rich, Ca-depleted brine at halite saturation and beyond, from which Mg-sulphate bitterns precipitate. Polyhalite can form syndepositionally at temperatures below 30 °C, via back-reaction of the evaporating K-Mg-SO4 brine with early-formed gypsum or anhydrite (Hardie 1984). Mg and Cl reach high concentrations during the latest bittern stages of evaporation resulting in the precipitation of carnallite and bischofite (Figure 1).

Utilisation

Today polyhalite is mined as a prime ore target in only one place in the world, the Israel Chemical Limited (ICL)-owned Boulby Mine, located below the North Sea, off the North Yorkshire coast of the UK. Currently, ICL-UK distributes around 500 kt/y from its Boulby mine as a direct application fertiliser or bulk-blend/compound NPK additive and is in the process of expanding its polysulphate output as it downgrades its MOP operations. A second major polyhalite mine, located near the Boulby Mine, is proposed by Sirius Minerals PLC. This renewed interest in polyhalite as an economical source of potash fertiliser is why I am writing this article.

Sulphate of Potash (SOP) fertiliser was initially derived from polyhalite/langbeinite in the US during the first half of the twentieth century. But then during the 1940s, after its discovery in vast quantities by prospectors searching for oil in Saskatchewan, the focus for the world’s potash fertiliser supply moved to Canada and the mining of sylvite (muriate of potash- KCl), (Warren, 2016). The geology of SOP was discussed in an earlier Salty Matters article (May 15, 2015). All aspects of SOP and MOP geology and mining are discussed in detail in Chapter 11, Warren (2016).

ICL currently markets crushed and processed polyhalite from the Boulby mine as polysulphate (Table 2). The purity of the Polysulphate product from the Boulby mine is very high (95% polyhalite) with <5% sodium chloride (NaCl) and traces of boron (B) and iron (Fe) at 300 and 100 ppm, respectively (Yermiyahu et al., 2017). The declared minimum analysis of polyhalite for S, K, Mg and Ca is 48% sulphur trioxide (SO3), 14% potassium oxide (K2O), 6% magnesium oxide (MgO) and 17% calcium oxide (CaO), respectively. This compares to a K2O content of contains 60–63% in MOP and around 50% in SOP.

Polyhalite’s make up in terms of K, Mg, S and Cl proportions is similar to the other major potassium-magnesium-sulphate (SOPM) fertilisers: Langbeinite, Schöenite and Patentkali® (Table 2). All are marketed as low-chloride potash fertilisers with additional magnesium and sulphur components. ICL-UK markets polyhalite as a multi-nutrient, low-chloride fertiliser under the brand name Polysulphate®. Sirius plans to market its polyhalite as POLY4®.

CRU estimates the global consumption of potassium-magnesium-sulphate (SOPM) fertilisers in 2017 at 1.7 Mt total product; a comparatively small total compared to the widely traded 65.5 Mt potassium chloride (MOP) market (https://www.crugroup.com "Will polyhalite disrupt the fertiliser industry?” published online April 2018; last accessed 12 July 2018). Polyhalite accounts for around 450-460 kt of current SOPM fertilisers. ICL-UK is currently ramping-up production to 1 Mt/y (140 kt/y K2O) by 2020, as it simultaneously phases out MOP production at the Boulby mine. At current production levels, this will be equivalent to almost 40% of the current SOPM market in K2O terms.

Sirius Minerals’ planned Phase I mine capacity (10 Mt/y product) is on a different scale altogether; around four times larger than the current SOPM market in K2O terms. This volume is almost the same size as the current global potassium sulphate (SOP) market, which is the most popular low-chloride potash fertiliser, outside China. The success of any future expanded SOPM application in agriculture is contentious; the majority of SOPM consumption has traditionally been concentrated close to production sites, nurtured by the local marketing efforts of producers. The proposed worldwide expansion will be tied to increasing acceptance by the agricultural community of polyhalite as an acceptable cheaper substitute for SOP and perhaps MOP.

Polyhalite as a fertiliser.

In the last few years, the use of polyhalite as an agricultural fertiliser has been tested successfully in a number of studies supported by Sirius minerals (Mello et al., 2018; Pavuluri et al., 2017). Polyhalite supplies four nutrients, is less water soluble than the more conventional potassium sources and may conceivably provide a slower release of nutrients. Studies comparing polyhalite to other K and Mg fertilisers have shown that polyhalite is at least as effective as potassium sulphate (K2SO4) as a slow release source of K, and at least as effective as potassium chloride (KCl) plus magnesium sulphate (MgSO4) as a source of K and Mg (Barbarick, 1991).


The possibility of successful use of polyhalite as a fertiliser is illustrated by the positive effects of its application on the growth of a potato crop at Tapira Brazil (Figure 2; Mello et al., 2017). Its impact on tuber starch and tuber dry matter exceeded that of either MOP or SOP applications. Cultivar Asterix at Tapira is mainly used for frying and chip-making (fries) in the food processing industry. High dry matter and starch content improve texture, and lower sugar contents result in less darkening of fries which is desirable. High dry matter percentage enables lower oil absorption while frying, resulting in lower oil usage per unit product. Tuber firmness is essential to handle mechanical stresses that may occur during tuber harvesting, transport, and storage. Crunchiness and hardness are positively related to starch and dry matter contents and specific gravity.

It seems polyhalite products are probably suitable as a low chloride fertiliser replacement of sylvite in some agricultural applications especially in arid acid, infertile soils as found in parts of Israel and other dry growing areas in the Middle East where salinisation due to fertiliser residues is a known problem. Yermiyahu et al., 2017, found the transport and leaching of Ca, Mg, K and S in soil following polyhalite application is lower than following the application of the equivalent sulphate salt fertilisers. The residual effect of polyhalite fertiliser on the subsequently grown crop is higher than the impact from the equivalent sulphate salts, especially regarding Ca, Mg and S. Irrigation management, as determined by the leaching fraction, has a substantial effect on the efficiency of polyhalite as a source of K, Ca, Mg and S for plant nutrition.

Geology of polyhalite

Polyhalite is a common constituent of many ancient evaporite sequences, especially in Permian and Neogene deposits, due to evaporation of Na-K-Mg-Cl-SO4 marine brines. These sulphate bittern assemblages correspond to periods of MgSO4-enriched ocean chemistries (Lowenstein et al., 2003; Demicco et al., 2005).


Modern polyhalite occurrences

The presence of syndepositional polyhalite in the supratidal evaporite flats around the Ojo de Liebre lagoon was first discovered by Holser (1966), who attributed its origin to the diagenesis of gypsum by interstitial marine brines. A large part of this area is now occu­pied by artificial salt ponds. However, some remnants of the an­cient evaporite flats are still accessible, for example, on the southeast coast of the evaporitic complex, where sedimentolog­ical, chemical and isotopic investigations were performed on evaporitic sediments and interstitial solutions (Figure 3a, Pierre, 1983). In May 1979, the evaporitic succession was mainly composed of gypsum; a few centimetres below the surface, polyhalite was present in the form of small nodules that were partially replacing former gypsum crystals (Figure 3b,c). In May 1980, this evaporitic succession was drastically modified, since polyhalite replaced gypsum sediments lying below the water table. This gives an exact timing for the mineral transformation from gypsum to halite of one year to replace 10cm thick interval of gypsum with a 10cm interval of polyhalite, which points to a chemical evolution of the solutions permeating the sediments.

During this one-year period, ionic concentrations of interstitial brine increased from thirteen to 18 times with respect to seawater concentration (Pierre, 1983). SO4 levels of interstitial solutions in the sabkha were higher than in normal marine brine and progressively increased in a landward direction, suggesting gypsum dissolution by groundwater crossflows.

Isotope study suggests both water and aqueous sulphate in the mudflat porewater have a mixed marine and continental origin (Figure 4). Thus, it appears that sulphate ions are provided in part by marine brines, in part by continental waters which have dissolved Pleistocene interstitial gypsum present at depth. The replacement of gypsum by polyhalite requires not only high Mg2+ and K+, but also high SO4 concentrations in the crossflowing solu­tions (Braitsch, 1971).


The polyhalite in Ojo de Liebre mudflats is diagenetic but also penecontemporaneous with the crystallisation of gypsum. However, in brines with temperatures >30°C, polyhalite may also be a primary co-precipitate with halite, as is occurring in recent saltworks near Santa Pola, SE Spain (as observed by B.C. Schreiber), and in cool-zone (cryogenic) salt lakes associated with widespread mirabilite-glauberite such as in Karabogazgol (Andriyasova, 1972).

Polyhalite is also a minor but widespread phase associated with glauberite in the Late Pleistocene-early Holocene sediments of Lop Nur China (Ma et al., 2016). There, the natural lake evaporites are nonmarine assemblages of mirabilite-glauberite-polyhalite-bloedite-gypsum-halite. The evaporitic stages of the lake fill contain massive amounts of glauberite and polyhalite compared to the other salts present. Polyhalite in the upper 40 m of the lake column and its predominance, is indicative of pervasive back-reactions, as is the presence of very minor amounts of carnallite and sylvite in the same section (Ma et al., 2010; Dong et al., 2012).

Ancient occurrences

Most ancient occurrences are interpreted as early diagenetic, formed in shallow brine crossflows as replacement of anhydrite or gypsum. Even so, there is a direct association between higher volumes of polyhalite in marine evaporite basins and times of MgSO4 enrichment of ocean waters.

Neogene polyhalite

Polyhalite is not found as a widespread primary precipitate in rocks of this age even though ocean chemistries are MgSO4-enriched. Instead, polyhalite is typically a minor but extensive early burial replacement of anhydrite or gypsum. The better-documented examples of this type of replacement polyhalite are found in association with gypsum and thenardite/glauberite in various Tertiary lacustrine basins of Spain. For example, below the exploited thenardite beds in the Madrid (Tajo) basin, Spain, the succession in the upper part of the lower Miocene unit is characterised by glauberite layers made up of a mixture of glauberite (45.8 %) and halite (41.7 %), with a minor polyhalite (7.8 %), dolomite (2.1 %), and clay minerals (1.8 %) (Herrero et al 2015).

The Madrid Basin is a large Tertiary intra-cratonic depression that contains some of the largest fossil sodium sulphate and sepiolite deposits in the world. Bedded sodium sulphates (glauberite and thenardite) are restricted to the Lower Saline Unit, where they are associated with anhydrite, halite, magnesite, polyhalite and minor clays. Glauberite and thenardite are thought to have been deposited in the most central part of a permanent saline lake. The accumulation of thenardite might have taken place during a stage of contraction of the lake system at the beginning of the middle Aragonian (middle Miocene).


Polyhalite occurs as a diagenetic saline phase related both to calcium and sodium sulphates occurrences. Both sepiolite and bentonite deposits are widely distributed within peripherally in distal fan and marginal lacustrine sequences in the so-called Intermediate Unit of the Miocene (middle to upper Aragonian). Thick beds of nearly pure sepiolite were deposited in ponds extended at the toes of arkosic alluviums. Sepiolite is also found within calcrete profiles in these environments. Minor amounts of sepiolite are commonly recognised along with palygorskite in open lacustrine areas. On the other hand, Mg-bentonites characteristically occur associated with dolostones and fine micaceous sands in sequences that provide evidence of fluctuations in the lake level. Polyhalite typically occurs as felty and spherulitic aggregates that alternate with centimetre-thick halite layers or millimetre-thick glauberite laminae in the Lower Saline Unit(Figure 5). The polyhalite crystals are always associated with micritic magnesite). In its turn, the felty polyhalite may be related to skeletal glauberite crystals. The hal­ite crystals commonly exhibit chevron-type mor­phologies. The thickness of the individual layers of halite ranges from 1 to 6 cm.

Similar polyhalite proportions are entrained in a number of glauberitic mineral assemblage in gypsiferous Neogene continental basins across the Iberian Peninsula, such as those of the Zaragoza (Salvany et al., 2007) or Lerín gypsum units (Salvany and Ortí, 1994), both occurrences are in the Ebro basin. In all cases, the polyhalite tends to be either massive or more typically a fibrous rim on large glauberite crystals.


Polyhalite also occurs as a minor phase in some potash regions the Messinian evaporites of the Mediterranean. In the mined succession exposed in the Realmonte mine,(southern Sicily) the halite unit is approximately 400 m-thick. From the bottom to the top, it consists of irregular anhydrite and marly mudstone breccia layer up to 2 m thick followed by units A to D (Figure 6; Lugli et al., 1999). Unit A, up to 50 m thick, contains evenly laminated halite with anhydrite nodules and laminae passing upward to massive halite beds with irregular mudstone bed some decimeters thick. Unit B (approximately 100m thick) consists of massive even layers of halite inter-bedded with thin kainite laminae, along with millimeter to centimetre-thick layers dominated by polyhalite spherulites and anhydrite laminae Figure 7; Garcia-Veigas et al., 1995). It may well be that along with kainite, the layers of polyhalite spherulites are primary co-precipitates at the potash bittern stage. The upper part of the succession contains several kainite layers up to 12 m thick. The 70–80 m thick unit C, consists of halite 10 to 20 cm thick layers separated by irregular mud laminae and it too contains minor polyhalite and anhydrite. Unit D, up to 60 m thick, begins with a grey anhydrite-rich mudstone passing to an anhydrite laminate sequence, followed by halite millimetre- to centimetre-thick layers intercalated with anhydrite laminae and decimetre-thick halite beds.


Lugli et al. (1999) proposed that these lithologies, including the early diagenetic polyhalite, reflect the shallowing and the desiccation of the evaporitic basin resulting from a possible combination of factors: (1) uplift of the basin floor by thrust activity, (2) simple evaporitic drawdown and (3) a basin-wide drop of the Mediterranean sea level.

Polyhalite is also common as a potash contributor along with, in the highly deformed bittern series in the Badenian (Middle Miocene) ores of the Carpathian Foredeep Figure 8a). These beds are highly distorted and host former potash mines extracting a kainite-langbeinite ore target (Figure 8b). These potash-entraining salt deposits occur in western Ukraine within two structural terranes: 1) Carpathian Foredeep (rock and potash salt) and (II) Transcarpathian trough (rock salt) (Figure 8a). Deposits differ in the thickness and lithology, depending on the regional tectonic location (Czapowski et al., 2009). In the Ukrainian part of Carpathian Foredeep, three main tectonic zones are distinguished (Figure 8a): (I) outer zone (Bilche-Volytsya Unit), in which the Miocene molasse deposits overlie the Mesozoic platform basement discordantly at a depth of 10-200 m, and in the foredeep they subsided under the overthrust of the Sambir zone and are at depths of 1.2-2.2 km (Bukowski and Czapowski, 2009); Hryniv et al., 2007); (II) central zone (Sambir Unit), in which the Miocene deposits were overthrust some 8-12 km onto the external part of the Foredeep deposits of the external zone occur at depths of 1.0-2.2 km; (III) internal zone (Boryslav-Pokuttya Unit), where Miocene deposits were overthrust atop the Sambir Nappe zone across a distance of some 25 km (Hryniv et al., 2007).


Potash evaporites of the Carpathian Foredeep host an interesting sulphates group that includes about 20 sulphate evaporite minerals. Exploited potash deposits of the foredeep are composed of kainite, langbeinite, kainite–langbeinite, sylvinite, polyhalite and carnallite rocks with layers of rock salt or interbedded clays and rock salt. In the areas of salt-bearing breccia, a polyhalite–anhydrite layer occurs along the contact with the potash salts bed. Halite, langbeinite and kainite dominated targeted ore levels in these potash deposits. Kieserite, polyhalite, anhydrite, sylvite and carnallite were present in smaller but significant quantities. These deposits, once a source of sulphate of potash, are no longer mined.

A study of sulphur isotopic composition of 10 of the sulphate minerals from the Kalush-Holyh and Stebnyk potash deposits shows that only the basal Ca-sulphates (anhydrite) from the Kalush-Holyn potash deposits has d34SCD values typical of Neogene marine evaporites (+21.0‰; Hryniv et al., 2007). Potash minerals related to the ore-associations in the deposits (polyhalite, anhydrite, kainite, langbeinite and kieserite) show d34SCD values from +15.28 ‰ to +17.54‰, while weathering zone minerals (picromerite, leonite, bloedite, syngenite and gypsum) in the Dombrovo Quarry show values ranging from +14.73‰ to +18.22‰ (Table 3).

According to Hyrniv et al. (2007) the recorded depletion of sulphur isotopic composition of the salt minerals in the Ukranian potash deposits (and their weathering zone) was probably caused by one or more factors as follow: 1) bacterial reduction of sulphate, 2) effect of crystallisation and 3) inflow of surface waters containing sulphates enriched in light sulphur isotopes due to pyrite oxidation. Accordingly, the observed sulphur isotopic composition of minerals from these potash deposits demonstrates the depletion of the original marine brines and continual inflow of new (concentrated) seawater and later meteoric access. The preponderance of lighter sulphur isotopic values recorded in the Stebnyk deposit can be explained by a more intensive inflow of surface waters from the Carpathian nappes or by the oxidation of a part of the pyrite hosted in the sediments. Whatever the case, it seems that once again polyhalite is an early diagenetic mineral.


Permian polyhalite

Permian polyhalite deposits are much more impressive in terms of volume and extent, compared to the Neogene, and are exemplified by massive occurrences in the USA and Europe

Permian polyhalite in West Texas and New Mexico

Polyhalite deposits are by far the most abundant, most numerous, and widespread of all potash mineral occurrences in the Delaware Basin of Texas and New Mexico (Jones 1972; Lowenstein, 1988; Harville and Fritz, 1986). However, langbeinite and sylvite are the economically important potash minerals and have been the focus of many studies, rather than polyhalite documentation (Figure 9a). Permian polyhalite in the Delaware Basin occurs both as massive and disseminated deposits in anhydrite and salt beds and less often in clay beds. Typically, massive deposits and all veins and lenses are composed predominantly of polyhalite, in distinctly compact units that have sharp, clear-cut outlines. Disseminated deposits generally are less defined, shapeless bodiesof spherules as cleavage-parallel growths in a host rock, chiefly in halite. Disseminated occurrences are many times more numerous than the massive deposits, but the amount of polyhalite present is minor in comparison with that present in most massive deposits in anhydrite beds.

Massive polyhalite occurrences outline a crude oval-shaped area in the basin, extending over a region about 325 km long and 220 km wide, covering practically the whole southern half of the area between the Pecos River and the eastern limit of salt in the Ochoa Series (Figure 9a). Occurrences range stratigraphically from low in the Tansill Formation (upper part of Guadalupe Series) in the North-western shelf to near the middle of the Rustler Formation in the north-east corner of the Delaware basin (Figure 9b). Polyhalite beds reach their highest number and size in the Salado Formation (Ochoan), where they have a wide distribution over much of the Delaware and Midland basins and adjacent platform and shelf areas (Figure 9). In the Salado Fm., thick clay seams occur as basal strata that underlie massive polyhalite/anhydrite beds (Harville and Fritz, 1986; Lowenstein, 1988). By virtue of the wide extent and number of massive deposits, polyhalite ranks next to halite and anhydrite among the major constituents of the Salado Formation. Sections with layered halite and polyhalite cover areas of 95,000 km2 and 70,000 km2, respectively (Jones 1972).

Massive polyhalite units are typically compact and fine-grained, exhibiting a variety of colours (grey to red) and textures (irregular to layered to laminated and fibrous to equicrystalline prismatic). Significant volumes are replacements of anhydrite beds, and although they may have almost any shape, most tend to be lenticular to sheet-like masses that spread out along the bedding and replace practically the entire section of anhydrite. Polyhalite units in the McNutt Potash zone, east of Carlsbad, have lateral continuities sufficient to act as marker beds, which separate and define layering in the sylvite-langbeinite ore zones (Figure 10).


As a general rule, sheet-like to crudely tabular polyhalite bodies occur in anhydrite layers where stacked polyhalite units are a few centimetres to a metre thick. Deposits that are more irregular in shape occur mostly in thicker beds of anhydrite (>1m). IN most cases the polyhalite is pseudomorphous after growth-aligned subaqueous and nodular gypsum or nodular anhydrite beds (Figure 11).

Practically all the deposits enclose residual strips and irregular remnants of magnesitic anhydrite, which are mottled and streaked with halitic and anhydritic pseudomorphs after gypsum. Commonly polyhalite crystals and multigrain aggregates project into the magnesitic anhydrite remnants either as elongate crystals and veinlike tongues or as aggregates having scalloped margins convex toward anhydrite.


In many places in the Carlsbad district and nearby parts of the north-western shelf, many of the massive polyhalite deposits grade laterally to an anhydritic hartsalz unit with ore grade levels of sylvite. This is the area known as the McNutt Member or the McNutt potash zone (Figures 9a, 10). The change from polyhalite to hartsalz coincides with a shift from unmineralized to sylvinitic rock peppered with sparse grains and veinlets of carnallite and other magnesium-bearing bittern minerals, such as langbeinite and polyhalite.

In 1988, Lowenstein recognised two types of metre-scale depositional cycles (Type I and Type II) within the McNutt Potash Zone (Figure 11). Both cycles record progressive drawdown and concentration of brine in a shallow, marginal marine drawdown basin. "Type I" cycles have a base of carbonate-siliciclastic mudstone, overlain by anhydrite-polyhalite that is pseudomorphous after primary bedded gypsum. This, in turn, is overlain by bedded halite and capped by muddy halite. Lowenstein (1988) concluded the McNutt Zone of the Salado Formation consists entirely of these two types of metre-scale sequences, variably stacked one upon another (Figure 11).

All units are interpreted as mostly marine-brine dominated units precipitated by evaporation of massive volumes of brines fed by marine seepage or periodic overflows of the Permian ocean water. The upper cap to Type II cycles influenced by inflows of continental groundwater (Figure 11).


A basal mudstone grades upsection into anhydrite-polyhalite that is commonly laminated. Laminae are defined by couplets of anhydrite or polyhalite separated by magnesite-rich mud (Figure 12a-c). The most significant feature of the anhydrite/polyhalite interval is the large number of crystal outlines that occur in the anhydrite-polyhalite laminae. These crystals are now composed of anhydrite, polyhalite, halite, or sylvite but are all interpreted as replacement pseudomorphs after primary gypsum because of their close similarity to typical bottom-nucleated subaqueous gypsum crystal habit,s such as "swallow-tail twins" (Figure 11). In some occurrences, the polyhalite is forming early diagenetic spherules in magnesite layers (Figure 13a). Elsewhere polyhalite directly replaces bottom-nucleated subaqueous gypsum or halite (Figures 12a, c, ), while yet elsewhere it grows as spherular clusters in halite that already has pseudomorphed aligned gypsum crystals Figure 12c). In other places, rippled gypsum beds are replaced by polyhalite and anhydrite. Syndepositional brine reflux likely drove replacement of subaqueous gypsum by anhydrite-polyhalite, in a fashion similar to that described by Hovorka (1992) for halite replacing growth-aligned gypsum.


At the microscopic scale, it is evident that polyhalite forms as a replacement (Figure 13). One of the most common modes of occurrence across the Salado Formation is as coalescing spherules growing in relatively undisturbed magnesite layers (Figure 13a). Elsewhere, coarser mm-scale polyhalite prisms have poikilotopically enclosed anhydrite crystals (Figure 13b). Felted fibrous polyhalite also surrounds euhedral halite (Figure 13c) or forms a replacement rim to halite in the langbeinite-sylvite ore layers (Figure 13d).

"Type II" cycles, lacking the basal mudstone and polyhalite/anhydrite beds, occur between Type I cycles and contain additional halite units (with thinly layered polyhalite) overlain transitionally by muddy halite (also with dispersed polyhalite). Complete brining-upward Type I and Type II cycles record a temporal evolution of depositional environment from a shallow saline lake to an ephemeral salt-pan-saline mudflat complex. The uppermost muddy halite unit interpreted as a continental-dominated sequence, sourced by meteoric inflow from surrounding land areas that mixed with variable amounts of seawater, either from residual pore waters or introduced into the Salado Basin by seepage.

Periodic invasions of seawater best explain the vertical stacking of Type I cycles in the Salado basin, perhaps coincident with eustatic sea-level rises (Lowenstein, 1988). The continental-dominated upper parts of Type I and II cycles formed during intervening periods of eustatic sea-level fall and low stand when nonmarine waters exerted more influence on the brine chemistry. According to Lowenstein (1988), the maximum time interval between major marine incursions averages 100,000 years. The layered nature of the polyhalite replacement implies that this occurred in each eustatic cycle, that is, the replacement was an integral part of the eogenetic hydrology and was not a burial diagenetic (mesogenetic) process.


Permian Polyhalite in Poland and Russia

According to Peryt et al., 2005 (and references therein) there are four polyhalite deposits in the Zechstein of northern Poland, and more than ten polyhalite-bearing areas in the adjacent part of Russia (Figure 14). In addition, K-Mg chlorides are found locally both in Poland and Russia. The K-Mg salts originated during the last stages of chloride accumulation within small, actively subsiding isolated salt basins of the salina type, which were probably tectonically controlled.

The paragenetic sequence in one polyhalite (Zdrada) deposit in the Zechstein of Poland was the result of a very early - penecontemporaneous polyhalitisation of anhydrite that had already pseudomorphed gypsum, much as is seen in the Delaware basin (Peryt et al. 1998). There polyhalite formed by altering anhydrite during crossflows of concentrated brines that were also responsible for potash deposition in local salt basins, while the sulphate-rich brines supplied by the dissolution of emergent parts of the sulfate platform (Peryt et al. 1998).


The timing of the polyhalitisation can be inferred from a S-O isotope crossplot (Figure 15; Peryt et al., 1998). The isotopic compositions of sulphate evaporites indicate that marine solutions were the only source of sulphate ions supplied to the Zechstein basin. The more negative oxygen values associated with the polyhalite compared to its anhydrite precursor indicates somewhat warmer solutions that drove the conversion to polyhalite. These solutions were more saline than those driving the initial shallow anhydritisation that replaced platform gypsum by a reaction with refluxing brines.

 

Polyhalite in the Zechstein of the UK

Polyhalite in the Boulby Mine and the proposed York mine both occur within the Permian Fordon Formation, of the 2nd Zechstein cycle (Z2) in northeast England (Figure 16; Table 4; Stewart 1963; Smith et al., 1986; Kemp et al. 2016). Although initially discovered in 1939, the deeper, polyhalite-bearing Fordon (Evaporite) Formation was largely overlooked until recently. ICL-UK operations at the Boulby Mine have largely depleted the sylvinite target in the Boulby Potash Member, so the mine is now transitioning into polyhalite extraction from the Fordon (Evaporite) Formation (Table 4). The historical output from the Boulby Mine was around 1 Mt/yr of refined KCl product and 0.6 Mt of road salt (Kemp et al., 2016). Polyhalite beds in the proposed York (Whitehall) mine are considered to be so high grade that they can be mined and marketed as SOPM fertiliser with no processing except crushing and sizing (Kemp et al., 2016).

Five evaporite cycles (EZ1-EZ5) are developed in the northwestern corner of the main Permian Zechstein basin where it comes onshore in the UK between Teesside and Lincolnshire (Table 2, Figures 16, 17).


The relationship between the evaporite sequence in the main Zechstein basin and its onshore, lateral gradation into shelf and then semi-continental clastic strata was described by Smith aet al., (1986). Potash salts are known from cycles EZ2, EZ3, and EZ4, and Britain’s only potash producer, the Boulby mine, exploits sylvite from the EZ3 Boulby Potash Member. Sylvite-bearing horizons are also known in the EZ2 cycle, but the key potash resource therein is polyhalite, first discovered in 1939 in the E2 oil exploration hole at Eskdale, Whitby (Stewart, 1949). The only known occurrence of potentially economic volumes of polyhalite in the UK is in the EZ2 Fordon (Evaporite) Formation in this area.

Mineral zonation in the Fordon (Evaporite) Formation was first described in detail by Stewart (1949, 1963) from the Eskdale and Fordon boreholes. Polyhalite was described as partly primary, but mostly a replacement of syndepositional anhydrite. Three subcycles were recognised at Fordon. The Lower subcycle was deposited in a basin that still displayed considerable topographic variation from a shallow-water shelf to a deepwater basin (Figure 17). It contains no known potash occurrences. The Middle subcycle, in which the polyhalite occurs, includes a large volume of basin-fill evaporites, chiefly halite, that filled accommodation space and smoothed out the shelf-basin geometry. Consequently, it shows considerable lateral variation in thickness. The Upper subcycle formed in uniformly shallow-water conditions with no clear distinction between shelf and basin. It hosts a persistent sylvite-bearing horizon known as the Gough Seam. Colter and Reed (1980) showed that Stewart’s mineral zones could be projected far beyond the Fordon borehole and were recognisable throughout much of the British section of the North Sea basin (Doornenbal and Stevenson, 2010).

The description of mineral zones at Eskdale and Fordon by Stewart (1949, 1963) relate to boreholes drilled through the shelf and basin, respectively. The precise correlation of the polyhalite- bearing sulfate deposits between these two environments, or zones, remains ambiguous (Kemp et al., 2016). At present, the polyhalite deposit is referred to as the Shelf seam in the Shelf zone, and the Basin seam in the Basin zone, with a Transition zone across the ramp and in its vicinity where great thicknesses of polyhalite and anhydrite occur with varying amounts of early diagenetic, displacive halite. In borehole SM2 there was solid evidence for overlapping Shelf and Basin seams, separated by 82 m of “sulphatic halite”. Both the shelf and basin polyhalite seams are considered to be of mineable thickness and grade in their relevant sectors, averaging over 12 m in thickness for high-grade sections of >85% polyhalite.

Kemp et al. (2016) argue that the polyhalite is almost entirely secondary, resulting from replacement reactions between freshly deposited anhydrite muds on the seabed, with dense, bottom flowing, K-Mg-rich brines. A sylvite-bearing bittern salt horizon is locally present near the top of the Middle subcycle in both the Basin and the Shelf (though less commonly) and is referred to here as the Pasture Beck seam, after the borehole (also known as SM1) where it was first cored and characterised (Figure 16).

Another sylvite-bearing bittern salt horizon is more commonly present near the top of the Upper subcycle in both Basin and Shelf and is referred to here as the Gough seam; described in the SM4 borehole, where it was first cored and characterised as containing relatively high-grade sylvite. It is not clear why this and the Pasture Beck potash seam are so localised and patchy in distribution, but they may result from bittern brine pools of limited area, cut off from each other as the aggrading basin filled up at the end of each subcycle.


At an even more local scale in the polyhalite ore intervals in the Boulby Mine, there are metre-scale domal= structures interpreted as a form of tepee structure (Figures 18; Abbott, 2017). The height of the domal-shaped structures exposed in the mine workings varies between ~0.4 m and 1.5 m (average = 0.9±0.1 m) and widths ranging from ~2.3 m to 10.5 m (average = 5.3±0.5 m).

Unlike the highly deformed halokinetic flow textures in the overlying sylvite of the Boulby Potash Member, it seems much of the polyhalite ore preserves mostly syndepositional diagenetic alteration structures. Most of the domal features do not show the overthrust brittle ridge crests that define most tepees (Kendall and Warren, 1987). Instead, the domal peak tends to be a fractured and folded local anticlinal culmination. Whether one calls these anticlinal deformations domes in the polyhalite a true tepee, depends on which definition of a tepee one chooses to use. The domal features are thought to be a soft sediment deformation features, formed via polyhalite dewatering, coupled with penecontemporaneous precipitation of halite in opening fractures and below anticline crests in shallow burial. Deformation was driven by fluid crossflows and escapes, as anhydrite converted to polyhalite.

Forming polyhalite?

Nowhere is the present or the past is there evidence of direct primary precipitation of polyhalite. By primary, I mean that to be considered a primary polyhalite, the crystals should drop out of a concentrating at-surface brine either as bottom-nucleated or foundered brine-surface crystals. Such primary textures are widespread in gypsum and halite units but not in polyhalite. Instead, polyhalite textures and isotopic signals indicate polyhalite forms via replacement of gypsum or anhydritised gypsum.

In the modern salt flats of Ojo de Liebre, we see polyhalite replacing gypsum. Likewise, in various Tertiary lacustrine basins in Spain, polyhalite is found in association with gypsum and thenardite/, and it is replacing a CaSO4 phase. In the Badenian marine evaporites of the Carpathian foredeep, the polyhalite is part of the kainite-langbeinite ore sequence. It is in the Permian of the UK and West Texas and New Mexico that the volumes of polyhalite become sufficient for it to become a potential ore target in its own right. Once again all the textural and isotopic evidence indicates polyhalitisation of anhydrite rather than primary precipitation. But this replacement is more likely to be eogenetic (driven by nearsurface hydrologies that were active in the depositional setting) rather than mesogenetic (burial).

The most likely driven mechanism was brine reflux moving highly saline seawater through shallowly buried units of platform or basinal gypsum and anhydrite. This shallow subsurface emplacement occurred while the gypsum anhydrite was still permeable, and so allowed the preservation of pristine texture (pseudomorphs) of the CaSO4 precursors.

Polyhalitisation of basinal and platform gypsum units in the mega-sulphate stages of a saline giant are driven by time separate hydrologies, tied to the changing brine levels in the drawndown basin (Figure 18; Warren, 2016). Marine-derived brine reflux through basinal anhydrites occurs during maximum drawdown in the mega-sulphate basin (blue arrow positions stage b in Figure 18), while reflux through the upper (marginal saltern) parts of a sulphate platform is a response to a relative highstand (blue arrow positions in stage c Figure 18; Warren, 2016 - Chapter 5). The likely loss of permeability as one goes deeper in a sulphate platform and the associated lessening in the volume brine crossflow probably explains why there is an interval of sub-economic polyhalitic sulphate separating the basinal from the shelfal ore zones. The sequence stratigraphic fill model also explains why the patchy potash intervals are located higher in the stratigraphy at the "fill and spill stage" of a hyperarid climate (stage e in Figure 16; Warren 2016).


A drawdown model encompassing two stages of polyhalitisation explains why much of the textures seen in the platform and basin polyhalite units contains evidence of both lamination and subaqueous shallow water deposition (Figure 19). In any megasulphate saline giant, the basin brine level can oscillate between shallow and deep and, depending on the nature of the overlying brine column, and we can deposit primary textures that are mm-cm laminates or upward-aligned gypsum growths, or displacive nodules. Bottom nucleated, upward aligned gypsum crystals indicate relatively stable and saturated bottom chemistries beneath a holomictic brine column (Chapters 1 and 2, Warren 2016). Without holomixis, brine reflux cannot occur. Layered and laminated gypsum sediments interlayered by carbonates indicate subaqueous deposition with fluctuating chemistries in the overlying column Laminites can form via changes in water chemistry in a meromictic deep water mass (as in the modern Dead Sea prior to 1979), or it can indicate a shallow overlying water mass subject to periodic freshening as in the salinas of southern Australia. If the layer and laminate gypsum.anhdrite is interlayered with units of bottom-aligned gypsum or its anhydritised "ghosts," as in west Texas and Poland, then the depositing waters in both units were shallow.


We can now take this reflux model for polyhalitisation and explain why the two polyhalite ore seams in the Forden Evaporite Formation are separated by a low quality "sulphatic halite-anhydrite" unit (Figure 20). At time 1 the basin is at its maximum lowstand and dense reflux brines are sinking into the basinal gypsum units. Water depths below the holomictic brine mass in the basin lows are relatively shallow. At time two the brine levels in the basin are much higher, and a gypsum platform is prograding into the basin. Water depths above the platform are shallow, while they are deep in the basin centre. When the water column is holomictic, brine reflux is occurring across the platform and out into the basin. However descending brines cannot penetrate into all parts of the platform due to compaction and earlier reflux of halite- saturated cements. Brines must pass beyond halite saturation to reach polyhalite (Figure 1). This early loss of permeability created a core of less altered anhydrite below the polyhalite replacement interval.

But we must now ask, why did polyhalitisation of large parts of sulphate platforms reach its zenith in the Permian. A pseudomorphing process with a halite-gypsum focus is seen throughout the rock record (Chapter 5, 7; Warren, 2016). But the volumes of polyhalite we see in the Permian saline giants are different to the much smaller volumes of the Neogen, which is also a time of MgSO4-enrieched waters. Polyhalite is never present in the gypsiferous units of the Messinian or Badenian saline giants in the same volumes we see in the Permian. Then again, the extreme hyperarid hydrologies we see in arid climate belts across the Pangean supercontinent are also unusual. But the seawater chemistry was not too different to that of today (Lowenstein et al., 2005).

References

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Stable isotopes in evaporite systems: Part III - 18O (Oxygen)

John Warren - Sunday, July 01, 2018

 

Introduction

Oxygen isotope determinations in evaporitic sediments are typically based on: 1) using oxygen held in the water molecule itself (H2O); 2) oxygen in the carbonate anion held in evaporitic dolomites or limestones or; 3) in sulphate from evaporitically precipitated gypsum or anhydrite. Oxygen measures on the water molecule can be co-associated with deuterium (D - heavy hydrogen) determinations. So isotopic sampling of evaporitic limestone and dolomites means carbon isotope values can be co-determined from the same mineral phase (CO3 source). Likewise, with the calcium sulphates, the sulphur isotope is always available for co-study (SO4 source in gypsum or anhydrite).

We have already discussed sulphur and carbon isotope variations in evaporitic settings in the previous two articles (30 April 2018 and 31 May 2018, respectively). So, in this article, we shall look at how oxygen isotope values vary with the co-associated deuterium, carbon and sulphur isotope phases. We focus on three sources for isotope samples (water molecules in a brine, evaporitic carbonate minerals, calcium sulphate minerals) and show that when oxygen values are co-plotted against deuterium, carbon or sulphur isotope values, it becomes a handy tool in defining depositional and diagenetic evolution in a range of evaporitic settings.


Oxygen isotope fractionation in water molecules in evaporating brines

The stable isotope community has long known of the potentially extreme effects of evaporation on the isotopic composition of liquids and the residual enrichment of the heavier isotope in the remaining brine. After all, Urey himself applied this knowledge when he demonstrated the existence of deuterium through evaporative enrichment of liquid hydrogen (Urey et al., 1932). Enrichment in heavier isotopes in the residual brine is documented in settings as diverse as evaporating Dead Sea brines (Gat, 1984) and degassing epithermal systems (Zheng, 1990).

As any water (brine) evaporates there is a commensurate preferential escape of the lighter 16O water molecules, this leaves behind an increasing proportion of heavier water molecules containing 18O. Hence, with increasing degrees of evaporation the δ18O signature in the remaining water mass becomes increasingly positive (Figure 1). Co-variance of deuterium with increasing oxygen isotope values in a concentring brine is a long-established observation (Figure 2; Cappa et al., 2003), and defines a type of Raleigh fractionation or distillation.


There is another factor involved in the degree of enrichment of the heavier isotopes of oxygen or deuterium, and that is the humidity of the air above the evaporating brine. Humidity controls the extent of evaporative concentration, and there is a differential level of isotope enrichment in the residual brine tied to changing humidity (Figure 2). It is a response to the lowering of the evaporation rate with increasing humidity. The humidity effect in evaporative settings is documented both experimentally and in natural settings such as modern sabkhas and salinas (Chapter 2 in Warren 2016, for a summary of literature). As a general rule, the lower the humidity, the greater the degree of enrichment of the heavier isotope. Horton et al., (2016) show that δ18OSMOW values of saline lake waters from are often shifted by >+10‰ relative to source waters discharging into the lake (Figure 3, especially 3c).


Up until February 1979, the Dead Sea was a permanently stratified hypersaline water body (see Warren 2016, Chapter 4 for hydrological and sedimentological details). Both the upper and lower water masses were moderately enriched in δ18OSMOW (Figure 4: Gat 1984). After the overturn and mixing the surface waters, the degree of enrichment in δ18O in the surface waters constitutes a balance between the dilution by freshwater influx and the isotope fractionation (enrichment) which accompanies evaporative water loss and vapour exchange with the atmospheric moisture. Gat's modelling of the seasonal cycle and long-term trends of δ18OSMOW in response to the changes in the environmental parameters, shows that the dominant control on isotope enrichment in the surface waters, post overturn, is exercised by the salinity of the surface waters, through its effect on the vapour pressure gradient between the lake's surface and the atmosphere. Interestingly, before the overturn event the upper water mass was more homogenous in terms of salinity and isotope enrichment and its enriched isotope values mostly tracked those of the much more stable and somewhat more saline lower water mass.


Deuterium-oxygen isotope plots of water molecules can also be useful in studying the origin of hydrated salts such as gypsum, but only if there has been minimal postdepositional alternation of the primary precipitate. A classic paper focusing on the composition of structural water held in the gypsum lattice of Messinian (Late Miocene) evaporites of Sicily was published by Bellanca et al., 1986. In Sicily, there are two main types of texture in gypsum-dominated outcrops in the Messinian sub-basins of Sicily (laminated and massive). The laminar gypsum, locally known as balatino, is a shallow-water saltern deposit, the other is a massive form of gypsum typically interpreted as a diagenetic replacement of either primary gypsum of anhydrite.

The different isotopic compositions of hydration water in the two gypsum lithotypes are shown in Figure 6. Laminar gypsum shows a predominance of positive values for both oxygen (range-1.59‰ < δ18OSMOW < +6.02‰) and deuterium (range -7.3‰ < δD < +22.7‰), while both oxygen and deuterium ranges in the massive gypsum are negative (-4.21‰ < δ18OSMOW > -2.23‰; -40.9‰ < δD < -34.4‰).


In Figure 5 the majority of points representative of the laminar gypsum mother waters fall to the right of the meteoric water line of Craig ( 1961) and lie on a path characterised by a positive slope (δD = 3.97δ18OSMOW - 0.59) and includes the SMOW point. Such a distribution is consistent with an origin of the gypsum by direct pre­cipitation from an evaporating solution saturated with respect to gypsum and is close to those of mother waters in recent gypsum samples precipitated in Mediterranean salinas and, there­fore, suggest that the solutions from which the laminar gypsum precipitated were marine waters concentrated by evaporation. A few other examples show δ18O and δD values shifted towards negative values, which indicate stages of dilution with large masses of continental waters poured into the deposition basin during the crystallisation of gyp­sum (Bellanca et al., 1986).

In contrast, the waters from Massive Gypsum plot along a line with a negative slope (δD = -2.66 δ18OSMOW -46.73). Clearly, these structural waters have a different origin. Bellanca (op. cit) argues these distinctive signatures are indicative of rehydration from anhydrite; others argue massive gypsum is a result of subsurface recrystallisation of primary gypsum without an intervening anhydrite stage (see Testa and Lugli (2000) for the detailed discussion of this topic)

Carbon and oxygen isotope co-variations in evaporitic carbonates

The isotopic makeup of residual water molecules evolving into a brine is not the only phase affected by the chemical consequences of evaporation (Horton et al., 2016; Warren 2016). As any natural water evaporates, its chemistry changes, as concentrating dissolved phases and increasing alkalinity force changes in equilibrium conditions. One of the most obvious consequences of evaporation is the formation of sedimentary evaporites, including brine pool carbonates (e.g. calcite, aragonite, dolomite, trona). The coupled δ18O and δ13C enrichment during evaporation, and the precipitation of endogenic Holocene carbonates is documented and discussed at some length in a number of review papers (Horton et al., 2016; Pierre, 1988).


Horton et al. (2016) document a general tendency for calcites precipitated in lakes located in somewhat less humid climates to show enrichment in the heavier isotope. The observed average lake carbonate δ18OPDB values from the 57 lakes plotted in Figure 6 are more positive than the modelled summer month meteoric water derived calcite δ18O values (Horton et al., 2016). Lake calcites precipitating in humid environments generally plot closer to the 1:1 line, suggesting lakes in these environments are less impacted by evaporative modification. Yet, 46 of the 57 lake records analysed (i.e. 81%) plot to the right of the 1:1 line consistent with evaporative modification of lake water δ18O. Forty-two percent of the lake carbonate δ18O records are >5‰ shifted towards more positive δ18O values than would be expected for summer-month carbonate precipitates derived from unmodified local meteoric water. Although many lakes with vastly different modern aridity index values show similar offsets between modelled and observed δ18O, lakes from currently arid and semi-arid environments have a much larger average δ18O offset (5.4‰) than sub-humid and humid environment lakes (2.0‰).


The dolomite forming lakes of the Coorong region show a similar set of enrichment in both oxygen and carbon isotopes within that type of Holocene dolomite precipitating directly from evaporating surface brines (dolomite Type-A; Rosen et al., 1989; Warren 1990, 2000). The other type of Holocene dolomite in the Coorong lakes (dolomite-B) shows no noticeable C-O covariant trend related to Raleigh distillation (Figure 7a). Type-A dolomite has a heavier oxygen isotope signature than type-B and is 3 - 6‰ heavier in 13C (Figure 7a). Type-A dolomite also has distinct unit cell dimensions (Rosen et al., 1989).

Type A tends to be magnesium-rich with up to 3-mole percent excess MgCO3, while type-B is near stoichiometric or calcian-rich. Type-A dolomite typically occurs in association with magnesite and hydromagnesite, Type B with Mg-calcite. Transmission electron microscopy (TEM) shows that Type A dolomites have a heterogeneous microstructure due to closely spaced random defects, while type B dolomites exhibit a more homogeneous microstructure implying excess calcium ions are more evenly distributed throughout the lattice. TEM studies show that the two types of Coorong dolomite are distinct and are not intermixed with other mineral phases; they are primary precipitates, and not replacements and are not transitional (Miser et al., 1987).

Within the lake stratigraphy the dolomites occupy two distinct positions, Type A dolomites occur as surficial 'yoghurt' textured gels that in each water-filled winter season are washed and blown across the lake surface. By late spring and through summer these surface waters have dried up (summer salinities ≈ 120‰), and the lake sediment surface is a mud-cracked interval of massive carbonate (Warren, 1990; 2016). Type B dolomites occur in the laminated unit that underlies the laminated with signatures implying precipitation from waters with bicarbonates, perhaps showing a stronger strong input from organic materials and are especially prevalent in the more marginward part of the laminated fille where meteoric groundwaters are continually flowing into the edges of the lakes and mixing with lake pore brines.

Figure 7b places these two Coorong dolomites in the context of other areas of primary dolomite accumulations within Holocene carbonate depositional settings. Today sulphate-reducing bacteria or archeal methanogens have been called upon to explain the primary precipitation of dolomite in bacterial biofilms in almost all these other settings. It is not my intention to question the importance of bacterial metabolism in these other dolomite-accumulating settings, only to point out the bicarbonate from which the Coorong type A dolomites have precipitated show a positive and co-variant enrichment in both carbon and oxygen valued that are more typical of evaporative concentration. Evaporative enrichment in carbon values tied CO2 degassing in highly saline waters was documented in the Dead Sea by Stiller et al., 1985 and discussed in last month's article (31 May 2018).

Evaporitic carbonates especially when interbedded with calcium sulphate beds can also dissolve and alter (Warren, 2016; Chapter 7). Evaporite-derived dedolomites are often associated with evaporite dissolution breccias, which indicates the stratigraphic position of the now dissolved calcium sulphate bed that supplied the excess calcium needed to dedolomitise (Lee, 1994; Fu et al., 2008). Dedolomite under this scenario forms via the reaction of calcium sulphate-rich solutions with pre-existing dolomite to produce calcite with magnesium sulphate as a possible byproduct. The latter is rarely preserved, as it is highly soluble, and either remains as dissolved ions in the escaping waters or is quickly redissolved and flushed by through-flowing groundwaters (Shearman et al., 1961). The CaSO4 dissolution process is often driven by meteoric flushing of nearsurface oxidising waters and former ferroan dolomites are preferentially replaced. The resulting calcitised dolomites are outlined by intervals stained red with iron oxides and hydroxides.


With uplift-related (telogenetic) dedolomites the distribution and isotopic composition of dedolomite can reflect variations in the regional hydrology. This can be seen in the dedolomites of the Lower Cretaceous Edwards Group in the Balcones fault zone area of south-central Texas (Ellis, 1985, 1986). The Edwards Group consists of 120-180 metres of porous limestone and dolomite that accumulated on the Comanche shelf in shallow-water subtidal, intertidal, and supratidal marine environments. During early burial diagenesis, carbonate mud neomorphosed to calcitic micrite, aragonite and Mg-calcitic allochems were altered to calcite or were leached, and evaporites formed in tidal-flat sediments. Each of these phases had a characteristic stable isotope signature (Figure 8). Dolomite is widespread and formed in environments ranging from hypersaline to fresh-water as shown by the two isotope clusters in the Edwards dolomite (meteoric versus evaporitic reflux).

Late Tertiary faulting along the Balcones fault zone, tied to Jurassic salt withdrawal, initiated a circulating, fresh-water aquifer system to the west and north of a fairly distinct “bad-water line,” which roughly parallels the Balcones fault zone. To the south of the bad-water line, interstitial fluids remained relatively stagnant and contain over 1000 mg/l dissolved solids. Because of the differences in the chemistry of the interstitial fluids, post-faulting diagenesis in the two zones has been very different.

Water in the bad-water zone can be saturated with respect to calcite, dolomite, gypsum, celestite, strontianite, and fluorite, whereas water in the fresh-water zone is saturated only with respect to calcite. Due to the change in water chemistry, rocks in the fresh-water zone have been extensively recrystallised to coarse microspar and pseudospar, extensive dedolomitization has occurred, and late sparry calcite cements have precipitated. This creates a suite of covariant isotope trends and clusters with the dedolomite showing a distinctive set of carbon and oxygen values relate to soil water influences indicated by calcites with more negative carbon values (Figure 8 indicated by brown shading). In contrast, rocks in the bad-water zone retain fabrics associated with pre-Miocene diagenesis, and there is little or no evidence of widespread dedolomite, indicated by pink shading in Figure 8.

The importance of meteoric diagenesis in the formation of dedolomite in shallow, subsurface telogenetic environments is illustrated by the fact that the Edwards Group had a stable mineralogy of calcite and dolomite before the circulation of fresh water began and drove the precipitation of meteoric spar, microspar and dedolomite. Isotopic values for the dedolomites follow a similar trend to those of the microspars and pseudospars. As with the microspars and pseudospars formed by the entry of telogenetic water, it can be shown that dedolomites are in isotopic equilibrium with Edwards water on a regional scale, which supports the contention that the dedolomites are still forming from crossflows of present-day formation-water (Ellis, 1985).

Sulphur and oxygen relationships in calcium sulphate

Modern seawater sulphate has a homogeneous and well-defined isotopic composition for both sulphur and oxygen:

34SSO4 = +20 ± 0.5‰ CDT

18OSO4 = +9.5 ± 0.5‰ SMOW

Likewise, the fractionation of sulphur and oxygen, which occurs during the transition from aqueous to the solid state of sulphate is also near constant at earth surface temperatures. For gypsum, the mean values of the isotope enrichment factor are (Pierre, 1988):

δ34Sgypsum—SO4 = 1.65‰

and,

δ18Ogypsum—SO4 = 3.5‰

Thus the δ34S and δ18O values of sulphate evaporites are directly related to the state of the aqueous sulphate reservoir wherever precipitation occurred. A plot of ancient marine CaSO4 evaporites shows the sulphur curve for seawater has varied across time from +30‰ in the Cambrian, to around +10‰ in the Permian and that it increased irregularly into the Mesozoic to its present value of +20‰. Oceanic oxygen isotope values show much less variability. Sulphur is largely resistant to isotopic fractionation during the increasing temperatures associated with burial alteration and transformation (Worden et al., 1997). All of these aspects are discussed in detail in the April 30, 2018 article.

With this knowledge of the relative lack of fractional in the subsurface compared to the much greater susceptibility of oxygen isotopes in the mesogenetic and telogenetic realms let us now look in more detail at the significance of oxygen variation in a variety of sulphate entraining settings.


Isotopically, the effects of dissolution and brine recycling in fracture-filling fibrous gypsum cements of various ages emplaced in a formation's burial evolution can be used define the sequential development of the superimposed diagenetic textures in the original gypsum unit (Figures 9, 10; Moragas et al., 2013). The upper Burdigalian Vilobí Gypsum Unit, located in the Vallès Penedès half-graben (NE Spain) and consists of a 60-m thick succession of laminated-to-banded primary and secondary gypsum. The unit is variably affected by Neogene extension in the western part of Mediterranean Sea. Tertiary extensional events are recorded in the evaporitic gypsum unit as six fracture sets and fills (faults and joints - S1 - S5), which can be linked with basin-scale deformation stages.

Combined structural, petrological and isotopic study of the unit by Moragas et al. (2013) established a chronology of fracture formation and infilling, from oldest to youngest as: (i) S1 and S2 normal faults sets with formation and precipitation of sigmoidal gypsum fibres; (ii) S3 joint sets with perpendicular fibres; (iii) S4 inverse fault sets, infilled by oblique gypsum fibres and associated with thrust-driven deformation of the previous fillings; and (iv) S5 and S6 joint sets tied to later dissolution processes and infilled by macrocrystalline gypsum cements likely related to the telogenetic realm. The fractures provided ongoing pathways for focused fluid circulation within the Vilobí Unit. The oxygen, sulphur and strontium isotope compositions of the original host rock and the various precipitates in the fractures imply ongoing convective recycling processes across the host-sulphates to the fracture infillings, as recorded by a general enrichment trend toward heavier S–O isotopes, from the oldest precipitates (sigmoidal fibres) to the youngest (macrocrystalline cements). The marine strontium signal is mostly preserved in the various postdepositional infillings, unlike the oxygen and to a lesser extent the sulphur isotope signals, which are evolving with the origin and temperature of the waters flowing in the fracture sets (Figure 10).


In any ancient silicified anhydrite nodule or bedded silicified succession, not all silica-replacing anhydrite in a particular region need come from the same source or be emplaced by the same set of processes. Silicified nodules within middle-upper Campanian (Cretaceous) carbonate sediments from the Laño and Tubilla del Agua sections of the Basque-Cantabrian Basin, northern Spain preserve cauliflower morphologies, together with anhydrite laths enclosed in megaquartz crystals and spherulitic fibrous quartz (quartzine-lutecite). All this shows that they formed by ongoing silica replacement of nodular anhydrite (Figures 10, 11; Gómez-Alday et al., 2002).

Anhydrite nodules at Laño were produced by the percolation of saline marine brines, during a period corresponding to a depositional hiatus. They have δ34S and δ180 mean values of +18.8‰ and +13.6‰, respectively, both consistent with Upper Cretaceous seawater sulphate values. Higher δ34S and δ180 (mean values of + 21.2‰ and 21.8‰ characterise nodules in the Tubilla del Agua section and are interpreted as indicating a partial bacterial sulphate reduction process in a more restricted marine environment (Figure 11a). Later calcite replacement and precipitation of geode-filling calcite in the siliceous nodules occurred in both sections, with δ13C and δ180 values indicating the participation of meteoric waters in both regions (Figure 11b). Synsedimentary activity of the Penacerrada diapir (Kueper salt - Triassic), which lies close to the Laño section, played a significant role in driving the local shallowing of the basin and in the formation of the silica in the nodules. In contrast, eustatic shallowing of the inner marine series in the Tubilla del Agua section led to the generation of morphologically similar quartz geodes, but from waters not influenced by brines derived from the groundwater halo of a diapir.


Conclusion

This and the previous two articles have underlined the utility of stable isotope samples of brine or precipitates in better understanding the origin of a range of brines and their associated precipitates. But other than the sampling of water molecules in modern brines, the interpretation of all isotope values is equivocal without a petrographic understanding of how and when the sampled textures formed. Stable isotopes of evaporitic minerals with sulphur, carbon and oxygen are the mainstays of isotope work in the study of most evaporite basins, both modern and ancient. Other isotopes that may be useful are 11B and 37Cl, and we shall look at their application to evaporitic sediments in a later blog.

References

Bellanca, A., and R. Neri, 1986, Evaporite carbonate cycles of the Messinian, Sicily; stable isotopes, mineralogy, textural features, and environmental implications: Journal of Sedimentary Petrology, v. 56, p. 614-621.

Cappa Christopher, D., B. Hendricks Melissa, J. DePaolo Donald, and C. Cohen Ronald, 2003, Isotopic fractionation of water during evaporation: Journal of Geophysical Research: Atmospheres, v. 108.

Ellis, P. M., 1986, Post-Miocene carbonate diagenesis of the Lower Cretaceous Edwards Group in the Balcones fault zone area, south-central Texa, in P. L. Abbott, and C. M. Woodruff, eds., The Balcones escarpment, geology, hydrology, ecology and social development in central Texas, Geological Society of America, p. 101-114.

Fu, Q. L., H. R. Qing, K. M. Bergman, and C. Yang, 2008, Dedolomitization and calcite cementation in the Middle Devonian Winnipegosis Formation in Central Saskatchewan, Canada: Sedimentology, v. 55, p. 1623-1642.

Gat, J. R., 1984, The stable isotope composition of Dead Sea waters: Earth and Planetary Science Letters, v. 71, p. 361-376.

Gómez-Alday, J. J., F. Garcia-Garmilla, and J. Elorza, 2002, Origin of quartz geodes from Lano and Tubilla del Agua sections (middle-upper Campanian, Basque-Cantabrian Basin, northern Spain): isotopic differences during diagenetic processes: Geological Journal, v. 37, p. 117-134.

Horton, T. W., W. F. Defliese, A. K. Tripati, and C. Oze, 2016, Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects: Quaternary Science Reviews, v. 131, p. 365-379.

Lee, M. R., 1994, Emplacement and diagenesis of gypsum and anhydrite in the late Permian Raisby Formation, north-east England: Proceedings - Yorkshire Geological Society, v. 50, p. 143-155.

Miser, D. E., J. S. Swinnea, and H. Steinfink, 1987, TEM observations and X-ray structure refinement of a twiined dolomite microstructure: American Mineralogist, v. 72, p. 188-193.

Moragas, M., C. Martínez, V. Baqués, E. Playà, A. Travé, G. Alías, and I. Cantarero, 2013, Diagenetic evolution of a fractured evaporite deposit (Vilobí Gypsum Unit, Miocene, NE Spain): Geofluids, v. 13, p. 180-193.

Pierre, C., 1988, Application of stable isotope geochemistry to the study of evaporites, in B. C. Schreiber, ed., Evaporites and hydrocarbons: New York, Columbia University Press, p. 300-344.

Rosen, M. R., D. E. Miser, M. A. Starcher, and J. K. Warren, 1989, Formation of dolomite in the Coorong region, South Australia: Geochimica et Cosmochimica Acta, v. 53, p. 661-669.

Shearman, D. J., J. Khouri, and S. Taha, 1961, On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura: Proceedings Geological Association of London, v. 72, p. 1-12.

Shearman, D. J., J. Khouri, and S. Taha, 1961, On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura: Proceedings Geological Association of London, v. 72, p. 1-12.

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

Testa, G., and S. Lugli, 2000, Gypsum-anhydrite transformations in Messinian evaporites of central Tuscany (Italy): Sedimentary Geology, v. 130, p. 249-268.

Urey, H. C., F. G. Brickwedde, and G. M. Murphy, 1932, A Hydrogen Isotope of Mass 2: Phys. Rev., v. 39, p. 164.

Warren, J. K., 1990, Sedimentology and mineralogy of dolomitic Coorong lakes, South Australia: Journal of Sedimentary Petrology, v. 60, p. 843-858.

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

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

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

 


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