Geochemistry of Evaporites and Evolution of Seawater

9.17 Geochemistry of Evaporites and Evolution of Seawater M Ba˛bel, University of Warsaw, Warszawa, Poland BC Schreiber, University of Washington, Seattle, WA, USA ã 2014 Elsevier Ltd. All rights reserved. 9.17.1 9.17.2 9.17.3 9.17.4 9.17.4.1 9.17.4.2 9.17.4.3 9.17.4.4 9.17.4.5 9.17.4.6 9.17.4.7 9.17.5 9.17.6 9.17.7 9.17.7.1 9.17.7.2 9.17.8 9.17.9 9.17.10 9.17.11 9.17.11.1 9.17.11.2 9.17.11.3 9.17.12 9.17.12.1 9.17.12.2 9.17.13 9.17.14 9.17.15 9.17.15.1 9.17.15.2 9.17.15.3 9.17.16 9.17.17 9.17.17.1 9.17.17.2 9.17.18 9.17.18.1 9.17.18.2 9.17.19 9.17.19.1 9.17.19.2 9.17.19.3 9.17.20 9.17.20.1 9.17.20.2 9.17.21 9.17.22 9.17.23 9.17.23.1 9.17.23.2 9.17.23.3 9.17.23.4 Introduction Definition of Evaporites Brines and Evaporites Environment of Evaporite Deposition Evaporation Freezing Brine (Evaporating Waters) Salinity Temperature Heliothermal Effect pH Seawater as a Salt Source for Evaporites Evaporite and Saline Minerals Model of Marginal Marine Evaporite Basin Conceptual Model of the Basin Quantitative Model of the Basin Mode of Evaporite Deposition Primary and Secondary Evaporites Evaporation of Seawater – Experimental Approach Crystallization Sequence before K–Mg Salt Precipitation Early Salinity Rise – Calcium Carbonate Precipitation Gypsum Crystallization Field Halite Crystallization Field Crystallization Sequence of K–Mg Salts Natural Crystallization Theoretical Crystallization Paths Isotopic Effects in Evaporating Seawater Brines and Evaporite Salts Usiglio Sequence – A Summary Principles and Record of Chemical Evolution of Evaporating Seawater Principle of the Chemical Divide for Seawater Jänecke Diagrams Spencer Triangle Evaporation of Seawater – Remarks on Theoretical Approaches Sulfate Deficiency in Ancient K–Mg Evaporites Sulfate Deficiency as the Secondary Feature Sulfate Deficiency as a Record of Ancient Seawater Composition Ancient Ocean Chemistry Interpreted from Evaporites Implications from Evaporite Mineralogy and from Usiglio Sequence Implications of Primary Evaporite Minerals (Excluding Implications from Fluid Inclusions) Recognition of Ancient Marine Evaporites Sedimentological Criteria Mineralogical Criteria Geochemical Criteria Fluid Inclusions Reveal the Composition of Ancient Brines Criteria for Seawater Recognition in Halite Fluid Inclusions Reconstruction of Ancient Seawater Composition from Halite Fluid Inclusions Ancient Ocean Chemistry from Halite Fluid Inclusions – Summary and Comments Salinity of Ancient Oceans Evaporite Deposition through Time Late Ediacaran–Phanerozoic Marine Evaporites Precambrian (Pre-Ediacaran) Marine Evaporites Nonmarine Evaporites in Precambrian Pseudomorphs after Evaporite Minerals in Precambrian Treatise on Geochemistry 2nd Edition http://dx.doi.org/10.1016/B978-0-08-095975-7.00718-X 484 484 484 485 486 486 487 487 488 489 489 489 490 491 492 495 496 498 499 499 499 501 501 503 503 504 505 505 505 505 507 508 509 509 511 513 514 514 516 516 517 517 517 518 519 520 523 530 531 532 536 544 544 483 484 Geochemistry of Evaporites and Evolution of Seawater 9.17.24 Significance of Evaporites in the Earth History 9.17.24.1 Paleogeographic Indicators 9.17.24.2 Seals for Hydrocarbons and More (Evaporites and Hydrocarbons) 9.17.24.3 Halotectonics 9.17.24.4 Diagenesis and Metamorphism of Evaporites 9.17.25 Summary Acknowledgments References 9.17.1 Introduction This chapter focuses almost exclusively on marine evaporites and in particular on how the chemistry of seawater is reflected in the mineralogy and facies distribution of deposits in geologic space and time. First, the deposits formed from evaporation of modern seawater are characterized together with their distinctive crystallization paths, and then, we show how the mineralogy and geochemistry of evaporites have been used for the interpretation of the chemical evolution of the ocean through time. In order to fill in the background of the main theme, we attempt to supply more detailed and up-to-date information on the geochemistry of evaporite environments and evaporite deposits important or relevant to the problem of their current geochemical studies. 9.17.2 546 546 546 547 547 547 548 548 Braitsch, 1971; Braitsch and Garrett, 1981). Some authors have suggested other names for salt rocks precipitated by mechanisms other than evaporation (e.g., Berkey, 1922; Debenedetti, 1976; Warren, 1996; Wood et al., 2005); however, these names (‘reactionites,’ ‘precipitates,’ ‘thermalites,’ ‘replacementites’, etc.,) are only rarely used in the geologic literature, or are not generally accepted. Nowadays, the term evaporites appears to be most commonly used in the very broad sense (cf. Twenhofel, 1950). Nevertheless, those evaporites strongly affected by diagenesis with the primary features obliterated, as those occurring in salt diapirs, are more commonly described as salt deposits (e.g., salt diapirs, not evaporite diapirs; cf. Hudec and Jackson, 2007). The name ‘salt deposits’ also can be applied to halite or calcium sulfate deposits precipitated from seawater in the zones of hydrothermal circulation in spreading zones of the oceanic crust (Berndt and Seyfried, 1997; Hansen and Wallmann, 2003; Hovland et al., 2006; Petersen et al., 2000; Talbot, 2008). Definition of Evaporites The Latin word ‘evaporo’ means ‘to change into a vapor,’ and it is used to designate the type of rocks and salts that originate during evaporation of natural solutes on the Earth’s surface. In the nineteenth and the beginning of the twentieth century, these deposits were simply termed ‘salt deposits’ and also, rarely, as evaporates (Goldschmidt, 1937; Grabau, 1920, p. 23). Although both terms, together with the term saline deposits, are in use today, the term evaporites (with modified spelling) introduced by Berkey (1922) became the most popular and it is widely accepted now. Evaporites are difficult to define precisely. The broad definition was suggested by Twenhofel (1950, p. 486) who understood the evaporites as a “group of sedimentary deposits whose origin is largely due to evaporation.” More exactly, he stated that “most evaporites result from evaporation of water of high concentration, but a few are formed by replacement, or freezing of concentrated waters” and added that “if subjected to heat and pressure, the evaporites form new combinations” (Twenhofel, 1950, p. 487). He also included the deposits that “develop through metamorphism of other evaporites” into this group (Twenhofel, 1950, p. 486). Evaporites are similarly defined in the current edition of the Glossary of Geology but include “rocks with saline minerals formed by other mechanisms, e.g., mixing of waters or temperature change” (Neuendorf et al., 2005, p. 221). Evaporite grains “reworked by wind or saline waters as clastic particles” are also considered as evaporites (Neuendorf et al., 2005, p. 221). The latter evaporites are termed allochthonous by Hardie (1984). Some authors restrict the term ‘evaporites’ for sediments formed exclusively by evaporation, and they use the name saline deposits or salt deposits for deposits formed not only by evaporation but also by cooling and salting out (compare 9.17.3 Brines and Evaporites The common feature of all evaporites is that they are composed of salts easily soluble in water (Goldschmidt, 1937). Such soluble salts accumulate in natural water reservoirs and in ocean waters in particular and are removed from these aqueous solutions in significant quantity, only by evaporation of the water. The essential feature of evaporites is that they precipitate from concentrated watery solutions or brines (Sonnenfeld, 1984, p. 1). Other inorganic chemical deposits usually contain minerals that are only slightly soluble in water. These minerals do not form as a result of evaporation of concentrated solutions. The chemical behavior of such substances is commonly relatively easy to predict and to study from solubility products and Eh–pH relations (Berner, 1971; Krauskopf, 1967). By contrast, the solubility of salts and their activity coefficients in brines vary widely and are not readily predictable as they are dependent on concentrations of other ions, among other factors (Karcz and Zak, 1987). In concentrated solutions, “the water structure was shown to be completely destroyed,” and due to ‘water deficiency,’ “the effects of ionic association and competition between oppositely charged ions for water molecules in their hydration shells are intensified” (Figure 1; Krumgalz, 1980, p. 73; Kostenko, 1982). “Formation of ion pairs and triplets apparently is so extensive in highly saline sulfate and carbonate brines that the true ionic strength may be less than half the value calculated from total molalities” (Berner, 1971, p. 48). The average ionic strength of standard seawater is about 0.7, but salt solutions with ionic strength greater than !1 may require more sophisticated models than those applied for seawater (Berner, 1971). Evaporating seawater brines attain ionic strengths nearly Geochemistry of Evaporites and Evolution of Seawater 485 7 All ions in the solution Na+ Mg2+ Start of gypsum Cl− precipitation SO42− Number of ions (n ! 1023) 6 5 Start of halite precipitation Start of epsomite precipitation 4 3 2 1 0 1 1.05 1.1 1.15 1.2 1.25 1.3 1.35 1.4 Density (g cm-3) 100 All H2O molecules Number of ions (n ! 1023) 90 Number of H2O molecules per one ion 80 70 60 Start of halite precipitation Start of gypsum precipitation 50 40 Start of epsomite precipitation 30 20 10 0 1 1.05 1.1 1.15 1.2 1.25 1.3 1.35 1.4 Density (g cm-3) Figure 1 Numbers of molecules in the evaporating Black Sea water, after data by Il’insky (1948, cited by Kostenko, 1982), recalculated by Kostenko (1982). Note ‘water deficiency’ – extreme deficiency in H2O molecules. equal to 1, approximately at the onset of CaCO3 precipitation, and can attain ionic strengths as high as 12.8 (Figure 2; McCaffrey et al., 1987) or even 17.40 close to the end of evaporation (Millero, 2009). For example the waters of the Great Salt Lake and the Dead Sea show ionic strengths of 6.4, and 7.9, respectively (Millero, 2009), and the experimental brines from La Playa lake in Spain – up to 15.52 (Lopez and Mandado, 2007). These brines and their salts thus require a quite different, more complex chemical theoretical approach and an ion-interaction model (Drever, 1997; Pitzer, 1973, 1995), which allows for the calculation of mineral solubility in electrolyte solutions of high ionic strengths. Ion-interaction models provide one of the best approaches applicable for the modeling of salt crystallization from natural solutions (e.g., Brookins, 1988; Christov, 2011; Hamrouni and Dhahbi, 2001; Harvie and Weare, 1980; Krumgalz et al., 1999; Millero, 2009; Ptacek and Blowes, 2000; Song and Yao, 2003; Voigt, 2001). Pitzer’s model (Pitzer, 1973, 1995) is commonly applied to waters with ionic strength, I > 0.72, whereas for waters with ionic strengths, I < 0.72, Debye–Hückel or other theories are best applied (Dargam and Depetris, 1996; Drever, 1997; Langmuir, 1997; Ptacek and Blowes, 2000). The behavior of concentrated brines at saturation or supersaturation with respect to one or more salts, as in an evaporite basin in the ‘productive’ state, may be even more complex. Two or more salts can crystallize simultaneously from the same evaporating brine and can form complicated double salts and many possible hydrates. During evaporation of highly saline brine, some hydrated precipitates can release water to the brine and can become dehydrated salt deposits (at ‘invariant points’). The development of such points of dehydration is a very specific feature of evaporite systems (Gamazo et al., 2011; Ordóñez et al., 1994; Sánchez-Moral et al., 1998). The extremely concentrated brines also show a lot of chemical and physical features and phenomena absent in seawater and freshwater, some of them still poorly understand (e.g., Buch et al., 1993; Karcz and Zak, 1987; Krumgalz, 1980; Sherwood et al., 1991; Sonnenfeld, 1984). Furthermore, the measurements of the chemical and physical parameters in brines are not easy and require special methods in extremely concentrated brines (e.g., Anati, 1999; Tanweer, 1993). 9.17.4 Environment of Evaporite Deposition Evaporite environments and brines are characterized by basic physicochemical features and parameters, such as salinity, 486 Geochemistry of Evaporites and Evolution of Seawater 14 Start of gypsum precipitation Ionic strength 12 Start of kainite precipitation 10 8 Start of epsomite precipitation 6 4 Start of carnallite precipitation Start of halite precipitation 2 0 0 10 20 30 40 50 60 70 80 90 100 Degree of evaporation Figure 2 Ionic strength of the evaporating Caribbean seawater brines, based on data by McCaffrey et al. (1987). temperature, and pH, which fluctuate within some limits. The most important of these physical, geochemical, and sedimentological features are reviewed (and some of them more clearly defined) in the following sections, beginning with the process essential for deposition of evaporites – the evaporation. 9.17.4.1 Evaporation Evaporation is the most effective way to separate dissolved salts from the water solute (i.e., promoting their precipitation). This term is widely used in the studies of evaporite deposits – instead of vaporization – “to signify that under natural conditions evaporation occurs just so long as the atmosphere is not saturated with respect to H2O, and so long as no liquid (vapor) phase is present” (Braitsch, 1971, p. 84). Evaporation acts in two ways – first, it is able to bring the undersaturated solution (unable to precipitate the highly soluble salts) into the state of saturation and supersaturation, and then, secondly, it is able to promote more or less continuous precipitation of salts from such a solution in the evaporite basin, which is then in the ‘productive’ state. Evaporation is the most important driving force in evaporite systems. The efficiency (rate) of evaporation is dependent on, or limited by, such factors as temperature (both the brine and the air), humidity, air movement, salinity, and factors such as the appearance of salt crust on the surface of brine, which inhibits the evaporation process (Groeneveld et al., 2010; Sonnenfeld, 1984). Evaporation “is of course greater where the relative humidity is low. Where however the temperature is also low, and therefore the total amount of water which the air can contain is small, saturation is soon reached unless the dry air is constantly replaced. When the temperature is high, evaporation may be much greater even in stagnant air, but where this air is in motion it will be very rapid and extreme. Hence the importance of drying winds, i.e. winds of low relative humidity” (Grabau, 1920, pp. 114–115). The quantitative data and models for rate of brine evaporation are given by many authors, Walton (1978), Laborde (1985), Chen (1992), Steinhorn (1997), Oroud (1999, 2011), Krumgalz et al. (2000), Al-Shammiri (2002), Kampf et al. (2005), and Gamazo et al. (2011), among others. The evaporation of freshwater and normal salinity seawater has been studied quantitatively for a long time and is a well-understood process. The evaporation of concentrated brines is more complex. During evaporation from brines and bitterns, the rate of evaporation slows with the rise of salinity. Evaporation stops at some extremely high-salinity conditions, because above a certain concentration, the brine becomes hygroscopic and adsorbs humidity from the air rather than drying, as shown, for example, by seminatural evaporation of the highly concentrated Dead Sea brine (by Zilberman-Kron 2008, described in Katz and Starinsky, 2009). In the Dead Sea, the condensation of atmospheric humidity into the brine during the summer–fall transition reversed the brine level drop in the experimental containers filled with brine (Yechieli and Wood, 2002). This property makes a seawater bittern a potentially good liquid desiccant (Lychnos et al., 2010). It was documented in the sabkha environment that the brine level in the ground rises during periods of elevated humidity, apparently due to the condensation of water from the air (Yechieli and Wood, 2002). The water from the sabkha surface can, however, evaporate to dryness during the daytime because its surface can attain very high temperatures (up to 60 " C). Here, solar heat contributes sufficient energy to remove water by evaporation, and warm air over the sabkha is able to absorb more water, which escapes from the system carried by the wind (Walton, 1978). In volcanic areas, thermal evaporation of the ground, heated from below, operates in a different way from solar evaporation as it produces slightly different, commonly dehydrated evaporite salt suites, stable in more elevated temperatures, and such evaporation commonly promotes acidity (Pulvirenti et al., 2009). 9.17.4.2 Freezing Freezing acts in a way similar to evaporation in its effect on brines – it removes H2O from the solution in the form of ice and produces a residual, concentrated brine as well as the crystallization of peculiar minerals (Stark et al., 2003). Freezing ends at an eutectic or cryohydric point when all compounds including H2O pass into the solid state (Mullin, 2001). The liquid brines with eutectic temperature below 0 " C are called cryobrines (Möhlmann and Thomsen, 2011), and such brines probably exist on Mars as well as in our Earth’s polar regions (McEwen et al., 2011). The minerals formed by freeze-drying are usually called cryogenic (e.g., Brasier, 2011). Some artificial calcium Geochemistry of Evaporites and Evolution of Seawater chloride cryobrines have extremely low eutectic temperatures, down to 210-215 K (Brass, 1980) and it is thought that some other brines show even lower eutectic temperatures (199 K: Fairén, 2010, his Table 1; 201 K: Möhlmann and Thomsen, 2011, their Table 1). During freezing of seawater, the eutectic temperature is reached at #54 " C when the freezing is according to the Ringer–Nelson–Thompson pathway or at #36 " C when it follows the Gitterman pathway, which is thermodynamically stable pathway for freezing of seawater (Marion et al., 1999; Stark et al., 2003). Freezing of brines is modeled by numerical simulations (e.g., Kargel et al., 2000). The most extreme natural case, on Earth, is noted in the permanently unfrozen Don Juan Pond in Dry Valley, Antarctica, containing one of the saltiest natural brines on Earth (salinity 388.9 g kg#1; Torii and Ossaka, 1965). Those brines can remain unfrozen down to #51 " C (Marion, 1997). 9.17.4.3 Brine (Evaporating Waters) The evaporation of water causes the amount of salt remaining in the water solution to rise – freshwater can become ‘brackish,’ then saline, brine, and finally a bittern. ‘Freshwater’ is defined as sufficiently dilute to be potable, that is, containing less than ! 1000 mg l#1 total dissolved solids (TDS; Drever, 1997), the value that is considered as close to a natural boundary of detection of human taste (Alekin, 1970; Hammer, 1986). The terms brackish, saline, and brine are not defined univocally and can be variously defined depending on the author and the country. According to Drever (1997), brackish waters are too saline to be potable but are significantly less saline than seawater and containing between 1000 and 20 000 mg l#1 TDS. In the older literature, ‘brackish’ concerns the transitional zone between marine and freshwater and refers to those waters of intermediate salinities being a mixture of freshwater and seawater (s. s.) (Hammer, 1986). Saline waters have salinities similar to or greater than seawater (35 000 mg l#1 TDS) (Drever, 1997). In limnology, however, the boundary between freshwater and saline waters is set at 3% (Bayly, 1972; Bayly and Williams, 1966), and the most-mineralized river waters were termed saline by Meybeck (2003) when sum of ion concentrations is only over 24 meq l#1, which is equivalent of 1.4 g l#1 NaCl. Drever (1997) also defined brines as waters significantly more saline than seawater that usually contain much NaCl and are strongly salty in taste. Bittern is the brine stripped of most of its sodium chloride content and with the bitter taste of a magnesium chloride solution (Lychnos et al., 2010; Sonnenfeld, 1984). The residual, most concentrated evaporated bittern is commonly left to soak into the precipitate and is called ‘mother liquor,’ although sometimes this term is also used alternatively with the bittern. Brines formed by evaporation can be called solar brines to distinguish them from the ascending subsurface hot brines called hydrothermal and supercritical brines (Talbot, 2008; note that the proper use of the genetic term hydrothermal requires special caution; see Machel and Lonnee, 2002). 9.17.4.4 Salinity Salinity is the most important parameter characterizing and defining saline water, brine, or bittern. It is a measure of the 487 total amount of salts dissolved in the water/brine, and brine, by definition, contains a lot of salt. It has been measured using many different units and in a number of ways. The absolute salinity, S, as defined by Forschhammer (1865), is the mass of dissolved salts in seawater, brackish water, brine, or other saline solution per mass of that solution and is given in dimensionless units: g per kg, % (per mill), or ppt (parts per thousand) (Anati, 1999; Gamsjäger et al., 2008): S ¼ massðdissolved saltsÞ=massðsaline solutionÞ [1] This salinity is very rarely measured directly. Usually, it is evaluated by measuring some other physical parameter (density, i.e., specific gravity, optical refraction index, electrical conductivity, etc.). It may also be measured by the concentration (contents) of some conservative element (ions that do not participate in any of evaporitic precipitation and hence are conserved in the solution) and accumulates in the brine during its evaporation (such as Cl, Mg, K, Li, and Br; e.g., Brantley et al., 1984) for which a calibrated conversion scale is known. Conversion scales are however different for any particular brine with its own chemical composition (e.g., Jellison et al., 1999; Zinabu et al., 2002). Recently, several standard units for the properties of seawater were introduced (Millero, 2010; Wright et al., 2011), of which the standard seawater composition reflecting the chemical composition of seawater is important for geochemical studies (Table 1; Millero et al., 2008). The recommended measure of salinity for seawater is given as a dimensionless, practical salinity ‘unit’ that is based on conductivity measurements and is commonly designated as ‘psu,’ which is not quite appropriate, because the practical salinity scale has no units (Millero, 1993, 2010; Millero et al., 2008). The seawater of average salinity is 35%, and in the practical salinity scale, it has a salinity of 35.000. The other commonly used dimensional measure of salinity is TDS (g l#1). It is the unit of total dissolved grams of solids per liter of brine and has the dimension of density, and, as such, it is both temperature- and pressure-dependent, and therefore, without the known temperature (at least), the information about the absolute salinity is always incomplete (Anati, 1999). The unit recommended for monitoring the advance of evaporation of seawater, and the associated processes including any salinity rise or fall, is the ‘evaporation ratio’ defined as the mass (weight) of H2O (not weight of brine) in the original seawater divided by the mass (weight) of H2O in resulting evaporated brine (Garrett, 1980; Holser, 1979a). The other similar unit is ‘volume ratio’ being described as ‘X seawater’ that is the total volume of original seawater to total volume of brine (including dissolved salts; Holser, 1979a). Logan (1987) used ‘volume reduction ratio’ (Ver) defined as Ver ¼ ðVo # Ve Þ=Vo [2] where Vo is volume of original seawater and Ve is volume of evaporative outflow. The other recommended method of monitoring the degree of evaporation (DE) of brine, particularly those trapped in halite fluid inclusions, is by the calculation of the degrees of evaporation for various elements (DEELEMENTS) of brines related to modern seawater composition (Levy, 1977). Any conservative element not removed during the salt precipitation can be used (McCaffrey et al., 1987; Raab and Spiro, 1991; 488 Table 1 Geochemistry of Evaporites and Evolution of Seawater Major and other ions concentrations in seawater after various sources Ca2þ Mg2þ Kþ Naþ SO42# Cl# HCO3# Sr2þ Br# CO32# B(OH)4 F# OH# B(OH)3 CO2 H2O Other Sum of halides Sum of components dissolved in water Sum of components dissolved in water, and water Holland et al. (1986) (from Holland, 1978), average concentration in seawater of 35% salinity (mmol kg#1) Hay et al. (2006) (after Gill, 1989)a; concentration % by weight (¼g kg#1) Lowenstein and Risacher (2009) (after Drever, 1988) (mmol kg#1 H2O) Hay et al. (2006) (after Gill, 1989)a; molal concentration Millero et al. (2008); mi (mol kg –1) 10.2 53.2 10.2 468.0 28.2 545 2.4 nd 0.84 nd nd nd nd nd nd nd 0.41 1.27 0.38 10.59 2.67 19.12 0.12 0.01 0.07 0.02 nd 0.03 nd nd nd 965.28 0.02 9.22 34.72 11 55 11 485 29 565 2.4 nd nd nd nd nd nd nd nd nd nd 0.0102 0.0524 0.0097 0.4608 0.0278 0.5394 0.0020 0.0001 0.0008 0.0003 nd 0.0015 nd nd nd 17,389.8474 nd 0.0106568 0.0547421 0.0105797 0.4860597 0.0292643 0.5657647 0.0017803 0.0000940 0.0008728 0.0002477 0.0001045 0.0000708 0.0000082 0.0003258 0.0000100 55.5084720 nd 1.1605813 1000.00 56.6690534 mi, molality (mol kg#1 of solvent); nd, no data. With modifications to make chlorinity (Cl) of 19.2 equal to a salinity of 34.72% (after Hay et al., 2006) a Vogel et al., 2010; von Borstel et al., 2000). For example, the DE, based on magnesium (DEMg), is calculated following the equation (Zimmermann, 2000, 2001): DEMg ¼ ðmmolMg=kgH2 OÞbrine =ðmmolMg=kgH2 OÞseawater [3] During the K–Mg salt precipitation from evaporating seawater, Mg, K, Br, and Rb are removed from the brine, and only Li and B remain as conservative or relatively conservative elements (Vengosh et al., 1992), which more clearly indicate the true DE (Zimmermann, 2001). The ratios of selected ions, for example, (Na/Cl)eq, Mg/Cl, and Br/Cl, also reflect the DE, with some limitations (Holser, 1963; Levy, 1977). The evaporation of seawater brines can be traced on the diagrams comparing the contents of some conservative components (e.g., Na and Cl, Mg and Cl; Lowenstein et al., 2001) or the ratios of such components (e.g., Mg/Cl vs. Br/Cl; Holser, 1963). Concentration ratios for marine sabkha brine aquifers require special calculations (Wood et al., 2002). The recommended measures of concentration in brines are g/100 g H2O, g/100 g solution (%), and mol/1000 mol H2O (Braitsch, 1971, p. 28). The mutual comparison of various salinity units is complicated, and in particular, such measurement requires the accompanying precise measurements of temperature. For a salinity accuracy of 0.02%, the temperature during salinity measurement must be monitored with an accuracy of at least 0.04 " C (Anati, 1999). The particular problem is within supersaturated brine in state of salt precipitation. It is difficult to measure its properties accurately not only because of the presence of the invisible suspension of salt microcrystals (<4 mm) but also because the continued salt precipitation changes the chemical composition of brine that influences the other physical parameters characterizing the brine and conversion scales (Anati, 1999; Stiller et al., 1997). The measured salinity of seawater brine may reach values of 504.8 (TDS, g l#1) at the beginning of the final bischofite crystallization stage (Fontes and Matray, 1993) and another type of brine in some continental lakes over 500 g l#1 (e.g., 557 g l#1, TDS), in one of the Wadi El Natrun alkaline lakes, Egypt (Taher, 1999). Boiling hot (110 " C) Na–K–Mg–Cl brine from hydrothermal springs in Dallol salt diapir in Ethiopia attains 420 g kg#1 (TDS; Hochstein and Browne, 2000). Subsurface brines can reach extremely high salinities such as 643 g l#1 recorded in Ca–Na–Cl-type brine from the Salina Formation in the Michigan Basin (Case, 1945). This brine showed also one of the highest density, 1.458 g cm#3, so far measured in natural brines. 9.17.4.5 Temperature The temperature variations in various Earth evaporite environments range from ca. minus 50 " C to ca. plus 70 " C, depending Geochemistry of Evaporites and Evolution of Seawater on the climatic zone, and usually oscillate around 20–30 " C in the warm zone of Earth. However, temperatures on the sabkha surface are up to 60 " C (Kinsman, 1969) and in the algal mats up to 60–80 " C (Kinsman, 1966) and air temperature above the sabkha may reach up to 50 " C (Warren, 2000). The temperature of the gypsum surface at Tule Spring, in the Death Valley, United States, reached 190 " F (¼87.78 " C; Hunt et al., 1966). The brine temperature in Salina Ometepec, Baja California, Mexico, attained 70 " C (Casas and Lowenstein, 1989). Normally, however, the brine in shallow pans, including solar saltwork pans at the stage of gypsum or halite crystallization, is no more than 30– 40 " C (Benison and Goldstein, 1999). Many salt lakes and lagoons, for example, Kara Bogaz in Turkmenistan, experience a temperature drop below 0 " C in the winter, and temperatures below #55 " C have been recorded from Antarctica salt lakes (Marion, 1997). Hourly, diurnal, seasonal, and annual temperature fluctuations are particularly important for precipitation of salts from saturated solutions because temperature change in such a solution promotes supersaturation and precipitation (e.g., Ganor and Katz, 1989; Sánchez-Moral et al., 2002). The relatively high temperatures of ancient evaporating brines have been documented by fluid inclusion studies, particularly in ancient halite deposits. For example, the homogenization temperature of fluid inclusions in Permian halite from Kansas, United States, yielded original brine temperatures from 21 to 50 " C (Benison and Goldstein, 1999) and from the Zechstein of Poland from 50 to 62 " C (Vovnyuk and Czapowski, 2007). Similarly, the minimum temperature of halite crystallization in Oligocene Rhine Graben evaporites in France was estimated as 63 " C (Lowenstein and Spencer, 1990), and even higher temperatures (71 " C on average) – for Permian Salado Formation in New Mexico (Lowenstein and Spencer, 1990). Studies of the Devonian sylvite and carnallite of the Pripyat Depression in Belorussia, interpreted as primary deposits, indicated brine temperatures 67–83 " C, although ancient hydrothermal activity is recorded in this area (Hryniv et al., 2007; Petrychenko and Peryt, 2004). Halite inclusions from the Lower Cambrian Angara Suite on Siberian Platform showed brine temperatures from 60 to 86 " C (Petrychenko et al., 2005). On the other hand, the ancient Silurian (Pridolian) halite of the Michigan Basin, United States, showed low temperatures (5–25 " C) suggesting relatively cool climate during deposition of these evaporites (Satterfield et al., 2005). 9.17.4.6 Heliothermal Effect The capacity to store heat increases together with a rise in salinity, and therefore, brine, when heated, cools more slowly than the fresh- or seawater. In calm density-stratified brine pools, heated by sun, the brine below a shallow pycnocline can be heated by the sun’s rays due to a heliothermal effect to extreme values – scalding unsuspecting bathers (Kirkland et al., 1983; Sonnenfeld and Hudec, 1980). The brine in Lake Ursului, Romania, reached 69.5 " C due to this effect, despite the fact it is in the temperate climatic zone (latitude 46" 350 N), (Telegdi von Roth 1899, in Kirkland et al., 1983). Similar heliothermal conditions cause temperatures of 60.5 " C in Solar Lake, Sinai, Egypt (Cohen et al., 1977), and 67 " C in the bottom clays of the Tuzluchnoye Lake, at Iletsk, Forecaspian Depression, Russia (Dzens-Litovskiy, 1968). In some 489 artificial heliothermal solar pans, a temperature above the boiling point of water was reached (109 " C, in New Mexico, United States, Lodhi, 1996; the boiling point of seawater bittern can reach 125 " C in the bittern having a density 38 degrees Baumé; Buch et al., 1993). The heating of brine, to 50 " C and more, both due to input of heliothermal and solar heat (insolation), apparently is limited to shallow water masses (Schreiber and Walker, 1992; Sonnenfeld, 1984). 9.17.4.7 pH pH in saline lakes brine can be as low as 1.7 as recorded in the acid lake Magic (Wave Rock 2), (240–280 ppt, TDS) in Western Australia (Benison et al., 2007; Bowen and Benison, 2009) and attains a pH !12.0, similar to the solar evaporation pans of the Magadi Lake, Kenya (Grant, 2006; Grant et al., 1999). Evaporating seawater brine commonly shows characteristic pH fluctuations (Figure 3). The pH of evaporating seawater may rise from the local values (e.g., 8.3 in Bocana de Virrilá, Peru) to ! 8.6 in the field of carbonate precipitation. It then drops rapidly to 7.3–7.5 at the beginning halite saturation field and, further on, more slowly to 7.0 within this field (Brantley et al., 1984; Des Marais et al., 1992; Dronkert, 1985; Landry and Jaccard, 1984; Levy, 1977; McCaffrey et al., 1987; Nadler and Magaritz 1980; Pierre and Ortlieb, 1981; Pierre et al., 1984a; Rieke and Chilingar, 1961; Vogel et al., 2010). Further drop to pH 5.7 is recorded in the K–Mg salt precipitation field in brine of the density 1.327 g cc#1 at 30 " C (Amdouni, 2000; Buch et al., 1993). The possible reasons for such pH fluctuations were listed by Levy (1977), Nadler and Magaritz (1980), and McCaffrey et al. (1987) and discussed by Bodine (1976) and Krumgalz (1980). The similar pH values of brine were recorded in fluid inclusions from ancient halites (Petrychenko, 1988; Petrychenko et al., 2005; Roedder, 1984), however Benison et al. (1998) found that pH could be as low as 0-1 in some Permian halite lakes. Proper pH measurements of high-salinity brines require special techniques (Bowen and Benison, 2009; Sonnenfeld, 1984). 9.17.5 Seawater as a Salt Source for Evaporites The mineralogy of evaporites depends on composition of salts dissolved in the evaporating water, and in particular on the proportions of specific dissolved ions. The source of salts for evaporite deposits are easily soluble chemical compounds (salts) dissolved in natural waters on Earth. There are many natural types of salt-containing waters in the Earth environments; however, the crucial source for most evaporites is the ocean – the largest reservoir of saline water that contains the largest amount of the dissolved salts (Table 1). The hydrosphere contains about 1386 million km3 of free (gravitational) water of which 96.5% (1 338 000 km3) is the ocean (Babkin et al., 2003), containing 1.4 x 1021 kg H2O (Lécuyer et al., 1998; Pope et al., 2012, Table S3). The volume of seawater may be up to 10% larger when water stored in bottom ocean silts is added (references in Babkin et al., 2003). Lakes contain only 176 400 km3 of water, that is, merely 0.013% of the hydrosphere. Among them, freshwater lakes account about 91000 km3 and salt lakes only 85 400 km3, and 97% of their volume is found in one lake – the Caspian Sea (Babkin et al., 2003). The ocean contains 490 Geochemistry of Evaporites and Evolution of Seawater 9 Start of gypsum precipitation 8.5 Start of halite precipitation pH 8 7.5 7 6.5 6 0 5 10 15 20 25 30 35 40 Degree of evaporation Figure 3 pH changes during evaporation of Caribbean seawater, based on data by McCaffrey et al. (1987). !47.578( 1018 kg of salts (for an average salinity of 34.72%; Hay et al., 2006). The eight major seawater ions, Cl#, SO42#, HCO3#, Br#, Naþ, Mg2þ, Ca2þ, and Kþ, comprise 99.76% by weight of the dissolved salts, and the most abundant Cl# and Naþ ions make up 85.59% by weight of the salt in the seawater (Hay et al., 2006). The total amount of the salts dissolved in the ocean is enough to form a continent “three times the size of Europe as it appears above the sea level” (Grabau, 1920, p. 50). The ocean was the main source of salts for the largest ancient evaporite deposits popularized by Hsü (1972) under the name the ‘saline giants’ (called also basin-center or basinwide evaporites; Kendall, 1992; Warren, 2006). The components of seawater are therefore the most important for evaporite deposition. From many elements dissolved in seawater and eight major ions listed earlier, only seven of them create stable and volumetrically significant salts being the product of seawater evaporation (proportion of total salts by weight, in %, is given in brackets): Naþ (30.51), Cl# (55.08), SO42# (7.69), Mg2þ (3.67), Ca2þ (1.17), Kþ (1.10), and HCO3# (0.35) (Hay et al., 2006). All these are also common in the other natural water reservoirs on Earth (lakes, rivers, rainwater, and groundwater); however, in different combinations and variable proportions, both are similar and quite different than those of seawater. These seven ions are able to create a dozen of various evaporite minerals, which are more or less common. Not only the composition of soluble salts (major ions) present in seawater and their volume but also the molar proportions of particular ions in this water decide the mineralogy of marine evaporite deposits. The molar proportions of major ions are constant in today’s ocean (Naþ > Mg2þ > Ca2þ ) Kþ and Cl# > SO42# > HCO3#), and this fact is of the crucial importance and is the basis of the global geochemical investigation. The ocean is currently in a nonstratified state and its water masses mix mainly due to continuing thermohaline circulation leading to homogeneity of its composition including semiclosed continental seas. The near constant ratios of seawater constituents (irrespective of the seawater salinity) were first noted by Marcet (1819). It was then documented, in more detail, by Forschhammer (1865) and is variously known as Marcet’s principle, Forchhammer’s principle, the principle of constant proportions (Horne, 1969; Millero et al., 2008; Schopf, 1980), or simply the first law of chemical oceanography (Dana Kester in Millero, 2010). That principle permits a calculation of the value of salinity from the known concentration of one component of given seawater. In such a way from the known concentration of all halogens, which is easy to measure and is defined as chlorinity (Cl in %), the salinity (S in %) is calculated according to the well-known Knudsen’s law (Millero, 2010): Sð%Þ ¼ 0:030 þ 1:805Clð%Þ [4] Due to continued mixing, the isotopic composition of many components is fairly constant in the present-day ocean, including H2O not contaminated by glacial melt or river waters (Holser, 1992; Knauth and Beeunas, 1986; Lécuyer et al., 1998), and also recently recognized the isotopic composition of boron, magnesium, and presumably also chlorine (Argento et al., 2010; Foster et al., 2010; Ling et al., 2011). This is the basis for the construction of many isotopic curves for ancient seawater, crucial for geochemical analysis of Earth sedimentary record as well as widely used for stratigraphic studies (Boschetti et al., 2011a; Holland, 2003; Holser, 1979b; Veizer et al., 1999). 9.17.6 Evaporite and Saline Minerals Braitsch and Garrett (1981) distinguished the evaporite minerals “that have crystallized during the solar evaporation of aqueous solutions, predominantly solutions of strong electrolytes” from saline minerals “consisting of soluble salts, the formation of which includes not only evaporation but also cooling and <salting out >” (Braitsch and Garrett, 1981, p. 451). The main mineral components of evaporites are present in surface and subsurface waters on Earth in the form of Geochemistry of Evaporites and Evolution of Seawater easily soluble salts. They include four cations (Naþ, Kþ, Mg2þ, and Ca2þ) and three anions (Cl#, SO42#, and HCO3#); the last is present in the water in association with CO32# depending on pH. The anion Br# is also common; however, it is not able to create its own solid compounds during evaporation but diadochically replaces Cl# in the crystal lattice of chlorine salts in small quantities (halite, carnallite, sylvite, etc.). The other much less common anions occurring locally in continental environments include B4O72#, NO3#, and I#. In some areas, rare anions like F#, or chromates, also appear. A number of other cations, like Sr2þ, Fe2þ, and Al3þ, are also components of some more or less rare evaporite or saline (soluble) minerals, or minerals associated with evaporites, and some such minerals are less soluble (a list of such recognized rare minerals is growing with time; Łaszkiewicz, 1967; Pueyo, 1991; Sonnenfeld, 1984). Major rock-forming evaporite salts are thus chlorides, sulfates, and carbonates of calcium, sodium, potassium, and magnesium, commonly creating hydrated compounds as well as double, triple, and more complex salts (Table 2). Borates, nitrates, iodates, fluorides, and chromates are encountered in continental environments and originate from specific types of waters. Common evaporite minerals include more than 80 minerals (Braitsch, 1971; Sonnenfeld, 1984; Stewart, 1963), and the majority of them are relatively rarely observed. The list of evaporite minerals is even longer when the minerals formed in extremely cold environments are concerned. Many of these minerals are encountered in other, nonevaporite environments precipitating from fluids of the similar composition to the evaporating brines. Only three evaporite minerals make up the major rockforming and volumetrically most important deposits, and these are gypsum (CaSO4 • 2H2O) and anhydrite (CaSO4), and less commonly halite (NaCl), all together estimated to form more than 90–95% of modern and ancient precipitates (Warren, 2006, p. 564). Very commonly, dolomite (CaCO3 • MgCO3) and magnesite (MgCO3) are associated with evaporites; however, generally, they are not treated as typical evaporite minerals. From these listed three minerals, gypsum is certainly the most common at the surface and in shallow subsurface. K–Mg salts are much rarer and include the following most common minerals: sylvite (KCl), carnallite (KCl • MgCl2 • 6H2O), langbeinite (K2SO4 • 2MgSO4), kainite (4KCl • 4MgSO4 • 11H2O), and polyhalite (K2SO4 • MgSO4 • 2CaSO4 • 2H2O) (Borchert and Muir, 1964; Stewart, 1963). Together with calcite (CaCO3), dolomite (CaMg(CO3)2), and magnesite (MgCO3), they constitute the 12 major minerals encountered in evaporite rocks (Table 3; Stewart, 1963). An important genetic classification of saline minerals may be made according to the source of salts in the brine of evaporite basin, that is, seawater and continental water. Accordingly, the saline minerals can be divided into (1) minerals of the marine or marginal marine evaporites and (2) minerals of continental or nonmarine evaporites (Braitsch and Garrett, 1981). Because many the same minerals occur in both groups, univocal distinction between the marine and nonmarine evaporite minerals is commonly impossible or difficult. Many minerals precipitated from seawater also occur in continental lakes with water very similar in composition to seawater. Some ancient mineral assemblages cannot be crystallized from recent seawater without major modification. Today, those assemblages only occur in saline lake environments: (1) Na 491 carbonate minerals, such as trona, nahcolite, and shortite; (2) Na silicate minerals, such as magadiite and kenyaite; and (3) Na or Ca borate minerals (Smoot and Lowenstein, 1991). The first assemblage of minerals is expected to form from the evaporation of the hypothetical Archean–Proterozoic soda ocean water (Kempe and Degens, 1985). Furthermore, the borates that occur in some Permian evaporite deposits are considered as marine in origin (Helvaci, 2005; Stewart, 1963). Sodium sulfates, such as mirabilite (Na2SO4 • 10H2O), are also encountered in marine evaporites. Mirabilite precipitates from almost any sulfate brine, including seawater brine, during freezing (Garrett, 1970; Valyashko, 1962). For example, it precipitates in winter in the Great Salt Lake that contains brine very similar to seawater brine (Hardie, 1985). Glauberite, which appears to be a typical continental evaporite mineral (Salvany et al., 2007), is theoretically a predicted product of evaporite precipitation from seawater (Holland, 1984). Hardie (1985) included glauberite into his listing of components of marine evaporites as well. The typical marine evaporite minerals, excluding those formed during freezing of seawater and seawater brine, include more than 30 soluble minerals (Sonnenfeld, 1984; Stewart, 1963). High solubility is the essential feature of saline minerals, reflecting the nature of evaporite deposits, that is, formed from the most soluble components (Goldschmidt, 1937), although there are some exceptions (Table 4; Braitsch and Garrett, 1981). 9.17.7 Model of Marginal Marine Evaporite Basin Chemical models for an evaporite basin are critical in understanding the sedimentology of evaporites, and a working model for a marine evaporite basin is particularly important. There is no functioning marine-sourced basin on the Earth today able to produce evaporites on the scale of ancient saline giants. This is perhaps due to the effect of sea-level rise and flooding of the coastlines related to ice-caps melting during decline of the last Pleistocene glaciation (Glennie, 1987). Additionally, there is no active K–Mg salt-forming basin of marine origin today. Small, short-lived halite basins can serve only as a partial analog of the ancient evaporite sedimentation that took place in the past, over areas comparable to continent sizes. Hence, without a substantial working model, there are difficulties in establishing a fully operative estimate of the hydrological, sedimentological, and geochemical processes that may operate in such a basin, particularly during the deposition of K–Mg salts, and many attempts have been made to build a reasonable model of the function of such a basin (Ba˛bel, 2007; Dronkert, 1985; Sonnenfeld, 1984). Some large saline lake basins, like the Dead Sea, can help in creating a match but cannot give answers to all the questions. As pointed out by Hardie (1984), what is understood by the term ‘marine’ evaporite basin is ‘at best’ a ‘marginal marine’ basin, surrounded by land and thus being always under some influence of nonmarine sources of water. Such a basin, depending on the degree of isolation from the sea, can evolve into a ‘nonmarine’ basin with the brine being the mixture of many types of waters inflowing into the basin from the land and/or from subsurface, with the water of the ancient ocean. The model of such a basin is of the crucial importance for the understanding of the geochemical evolution of the ancient 492 Geochemistry of Evaporites and Evolution of Seawater Table 2 Significant evaporite and salt minerals (after various sources, abbreviations for some minerals after Eugster et al., 1980; Usdowski and Dietzel, 1998) Chloride Simple salts ha sy hh bi Ant Double salts ca Tc Triple salt Sulphatochlorides Double salt ka Triple salt da Carbonate Simple salts A Table 2 (Continued) lg le pc Halite Sylvite Hydrohalite Bischofite Antarcticite NaCl KCl NaCl * 2H2O MgCl2 * 6H2O CaCl2 * 6H2O vc c r, S r vr Carnallite Tachyhydrite KCl * MgCl2 * 6H2O CaCl2 * 2MgCl2 * 12H2O c r Rhinneite FeCl2 * 3KCl * NaCl r G ks Ki hx ep th mi Double salts gs Ap vh bl lw (Continued) Gl Syn Triple salt Po Chlorocarbonate Triple salt K2SO4 * 2MgSO4 K2SO4 * MgSO4 * 4H2O c r K2SO4 * MgSO4 * 6H2O r Na2SO4 * CaSO4 K2SO4 * CaSO4 * H2O c r Polyhalite K2SO4 * MgSO4 * 2CaSO4 * 2H2O c Northupite Na2CO3 * MgCO3 * NaCl r Burkeite Na2CO3 * MgSO4 r Tychite 2Na2CO3 * 2MgCO3 * Na2SO4 r Hanksite 9Na2SO4 * 2Na2CO3 * KCl r Lautarite Ca(IO3)2 vr Niter Nitratine Darapskite KNO3 NaNO3 Na2SO4 * NaNO3 * H2O r r r Kernite Borax Colemanite Na2B4O7 * 4H2O Na2B4O7 * 10H2O Ca2B6O11 * 5H2O r r r Ulexite NaCaB5O9 * 8H2O r Tarapacaite Lopezite K2CrO4 K2Cr2O7 vr vr Sulphocarbonate Double salt Kainite 4KCl * 4MgSO4 * 11H2O c Triple salt D’ansite Aragonite, calcite Thermonatrite Natron (natural soda) Nahcolite MgSO4 * 3NaCl * 9Na2SO4 r CaCO3 Na2CO3 * H2O Na2CO3 * 10H2O vc c c NaHCO3 c Double salts Sulfate Simple salts A Langbeinite Leonite Picromerite (¼schoenite, schönite) Glauberite Syngenite (¼kalushite) vca r c c c c Dolomite Huntite Trona Shortite Pierssonite Gaylussite (¼natrocalcite) CaCO3 * MgCO3 3MgCO3 * CaCO3 NaHCO3 * Na2CO3 * 2H2O 2CaCO3 * Na2CO3 CaCO3 * Na2CO3 * 2H2O CaCO3 * Na2CO3 * 5H2O Anhydrite Bassanite (hemihydrate) Gypsum Kieserite Sanderite Leonhardtite Pentahydrite (¼pentahydrate, allenite) Hexahydrite (¼sakiite) Epsomite (¼reichardtite, bitter salt) Thenardite Mirabilite (¼Glauber’s salt) Celestite CaSO4 CaSO4 * 0.5H2O vc r CaSO4 * 2H2O MgSO4 * H2O MgSO4 * 2H2O MgSO4 * 4H2O vc c r r MgSO4 * 5H2O vr MgSO4 * 6H2O r MgSO4 * 7H2O c Na2SO4 Na2SO4 * 10H2O c c, S SrSO4 c Na2SO4 * 3K2SO4 r 3Na2SO4 * MgSO4 Na2SO4 * MgSO4 * 4H2O vr c 2Na2SO4 * 2MgSO4 * 5H2O r Glaserite (¼aphthitalite) Vanthoffite Bloedite, blödite (¼astrakhanite) Loeweite, löweite Sulpho-chlorocarbonate Triple salt Iodates (exemplary) Nitrates and sulphatonitrates (exemplary) Borates (exemplary) Simple salt Double salt Chromates Components of evaporite rocks: vc, very common (typically rock-forming); c, common and relatively common; r, rare; vr, very rare; S, seasonal mineral, precipitates from cooled or frozen brine, and dissolves in warmer brine. a Ideal stoichiometric composition, the chemical formulae of natural dolomite varies from calcian to magnesian dolomite, Ca(1þx)Mg(1#x)(CO3)2, and documented composition ranges from Ca1.16Mg0.94(CO3)2 to Ca0.96Mg1.04(CO3)2 (Warren, 2000). evaporite deposits and for quantitative geochemical studies. The ‘universal’ conceptual and quantitative model of the evolving marginal marine evaporite basin developed during the studies of evaporites is briefly outlined in the succeeding text. 9.17.7.1 Conceptual Model of the Basin The marginal marine evaporite basin is commonly considered as a depression separated from the sea by a topographic barrier, which can drown or emerge, and with sporadically restricted open water connections between the basin area and the open Geochemistry of Evaporites and Evolution of Seawater ocean. The former basin type was termed a salina, the latter is considered a salt lagoon (Figure 4(a) and 4(b)), and one basin type can pass into the other during the course of geologic evolution, such as has been observed in the Kara Bogaz (DzensLitovskij and Vasil’ev, 1962; Grabau, 1920; Kosarev et al., 2009). Table 3 Major minerals encountered in evaporite rocks, excluding siliciclastic components (Borchert and Muir, 1964; Garret, 1970; Stewart, 1963) Anhydrite Calcite Carnallite Dolomite CaSO4 CaCO3 KCl * MgCl2 * 6H2O CaCO3 * MgCO3 Gypsum Halite Kainite Kieserite Langbeinite Magnesite Polyhalite Sylvite CaSO4 * 2H2O NaCl 4KCl * 4MgSO4 * 11H2O MgSO4 * H2O K2SO4 * 2MgSO4 MgCO3 K2SO4 * MgSO4 * 2CaSO4 * 2H2O KCl Table 4 (ideal stoichiometric composition) 493 Both basin types can eventually pass into a saline lake when the influx of seawater to the basin is entirely arrested (Figure 4(c)). The salina model of the basin has many well-recognized modern and subfossil analogs (Ba˛bel, 2007; Logan, 1987; Nunn and Harris, 2007) and until now was considered as the most important or even the only one reasonable model for ancient saline giants (e.g., Rouchy and Caruso, 2006; Warren, 2010). The salt lagoon model (Dronkert, 1985; Sonnenfeld, 1984) was criticized as hydrologically unsound and unrealistic (Kendall, 1988, 2010; Kendall and Harwood, 1996; Shaw, 1977; Warren, 2000); however, the recent evaporite deposition and refluxing bottom brine currently recorded in the Kara Bogaz (Kosarev et al., 2009) suggest that this model (Figure 4(a)) is functional in some instances. The crucial element in creating an evaporite basin is the requirement of a negative water balance. The outflow of water (brine) from the system (via evaporation, seepage, or a return current) should be greater than inflow of the waters of any kind (seawater plus meteoric water). This implies the presence of a number of climatic factors that accelerate the rate of evaporation, which should exceed precipitation at least during some part of the year (Schmalz, 1971). Solubility in water of evaporite salts and some other minerals Mineral Chemical formula Solubility (g l#1) Temperature (" C) References Barite Calcite BaSO4 CaCO3 Aragonite Dolomite Celestite Gypsum CaCO3 CaMg(CO3)2 SrSO4 CaSO4 · 2H2O Anhydrite CaSO4 Glaserite (aphtithalite) Epsomite (reichardtite, bitter salt) Na2SO4 · 3K2SO4 MgSO4 · 7H2O Hexahydrite (sakiite) Sylvite MgSO4 · 6H2O KCl Halite NaCl Thenardite Na2SO4 Mirabilite (Glauber’s salt) Na2SO4 · 10H2O Bischofite MgCl2 · 6H2O Antarcticite CaCl2 · 6H2O 0.0025 !0.012a 0.06b, 0.4c !0.014a 0.05b, 0.3c 0.114 0.207e 2.0 2.4 0.20f 0.275g 2.98 145 38.5e 262 308 32.95f 35.5g 340 35.86f 36.12g 360 16.83f 388 5.0e 28.0e 448 56.7e 2635 5360 25 + 1 25 25d 25 25d 25 + 1 25 20 25d 18 25 20 20 25 20 20 18 25 20 18 25 25d 18 40 0 25 20 25 20 20 Davis and Collins (1971) Pia (1933) in Hutchinson (1975, p. 661) Freeze and Cherry (1979) Pia (1933) in Hutchinson (1975, p. 661) Freeze and Cherry (1979) Davis and Collins (1971) Bock (1961) Borchert and Muir (1964) Freeze and Cherry (1979) Smith (1918); Grabau (1920, p. 21) Bock (1961) Sonnenfeld (1984) Borchert and Muir (1964) Grabau (1920, p. 22); Seidell (1940) Borchert and Muir (1964) Borchert and Muir (1964) Smith (1918); Grabau (1920, p. 21) Grabau (1920, p. 22); Seidell (1940) Borchert and Muir (1964) Smith (1918); Grabau (1920, p. 21) Grabau (1920, p. 22); Seidell (1940) Freeze and Cherry (1979) Smith (1918); Grabau (1920, p. 21) Borchert and Muir (1964) Grabau (1920, p. 22); Seidell (1940) Grabau (1920, p. 22); Seidell (1940) Borchert and Muir (1964) Grabau (1920, p. 22); Seidell (1940) Borchert and Muir (1964) Sonnenfeld (1984) a Solubility in supposedly CO2-free water, solubility is higher in water containing much dissolved CO2. At partial pressure PCO2 = 10#3 bar. c At partial pressure PCO2 ¼ 10#1 bar. d And at 1 bar (105 Pa) pressure. e Amount of pure compound without water of crystallization, in g/100 g H2O. f In g/100 cm3 H2O. g In g/100 g H2O. b 494 Geochemistry of Evaporites and Evolution of Seawater 3. If a salt is formed, evaporation is the driving force that leads to the precipitation of this salt (Mullin, 2001). The removal of water to the atmosphere by evaporation raises the concentration of particular ionic components of this salt which is necessary for its crystallization. Lagoon basin Sea level p rm e sc oc sr (a) sc - Surface inflow current oc - Bottom outflow current sr - Seepage reflux rm - Run-off meteoric water p - Precipitation e - Evaporation Salina basin rm Sea level p e Basin water level d se sr (b) su d - Range of evaporite se - Seepage influx drawdown su - Surface inflow p - Precipitation sr - Seepage reflux rm - Run-off meteoric water e - Evaporation Saline lake No any seawater inflow through the barrier rm Sea level p e Basin water level X sr (c) rm - Surface run-off sr - Seepage reflux p - Precipitation e - Evaporation Figure 4 Principal models of evaporite basins: the marginal marine evaporite basin of the lagoon (a) and the salina type (b) and the marginal saline lake basin (c). Evaporation is a crucial factor in such a basin playing three fundamental roles: 1. It lowers the basin water level (depressed water level defines the range of the evaporite ‘drawdown’ during the emergence of the barrier separating the basin from the sea; Figure 4(b)) and builds the hydraulic head that forces the marine water to flow or seep into the evaporite basin utilizing a permeable barrier, thus promoting the transport of dissolved seawater salts to a depositional site. This mechanism explains the great thicknesses of ancient evaporite deposits. 2. It raises the salinity of the basinal water and produces a brine; the concentrations of particular ions in the brine increase together with the increase in salinity in a process known as evaporitic concentration. In the zones of high evaporation, waters in hydrologically closed and semiclosed basins, due to a negative water balance, evolve slowly toward a state of saturation and supersaturation of soluble salts. According to Valyashko (1962), the life of such basins can be divided into earlier ‘preparatory’ and later a ‘selfprecipitating’ stage, beginning from the time when the state of saturation and supersaturation of the first soluble salt is reached in the water body, and this salt is precipitating in the basin. The dominating mechanisms of deposition in ‘self-precipitating’ basins are the processes of crystallization, dissolution, and transformation (early diagenesis) of salts. Therefore, these are crucial concepts for understanding the sedimentary record of such basins and principles of their development. Evaporation, however, is not the only driving force of the salt precipitation in such a productive basin. The salts can also crystallize due to temperature changes of the saturated solution (Sloss, 1969), mixing of brines known as salting out or salination (Raup, 1970, 1982; Sonnenfeld, 1984), and also brine freezing or freeze-drying may occur in areas of very cold climates (Marion et al., 1999; Sonnenfeld, 1984; Strakhov, 1962). These cold climate salts are the most ephemeral deposits on Earth. Water is the main carrier of salts. The water flowing into the evaporite basin transports dissolved salts into the areas of deposition. The largest mass and volume of such salts normally come from the sea because the seawater contains much more salt than any kind of meteoric water inflowing into a basin. Most meteoric waters observed on continents (waters derived from rain or snow) contain much less than 1 g kg#1 of dissolved solids. The ‘salinity’ or content of dissolved salts (Ca2þ, Mg2þ, Naþ, Kþ, CI#, SO42#, HCO3#, NO3#, and dissolved SiO2) in large rivers all over the world ranges from a few milligrams per liter to 1000 mg l#1 (Meybeck, 1976), and the average total concentration of dissolved solids in the large world rivers was estimated as 89.2 mg l#1 (Meybeck, 1979). The river waters flowing into an evaporite basin would be thus at least 35 times less saline than seawater but generally are much less saline. This means that the influx of such freshwater to the saline water, even in large volume, is not able to substantially influence the ionic proportions (ratios) of the dissolved salts within it. Furthermore, fresh nonmarine, meteoric, or river waters would be able to modify the ionic proportions in the basinal brines only if the ionic proportions in that nonmarine water were significantly different than that in the seawater or the seawater brines in the basin. The influence of nonmarine waters would be negligible when these differences are small. Therefore, even though the basin is supplied with extremely large amounts of meteoric water, which otherwise is difficult to expect in arid evaporite environments, marine salt composition very likely dominates the mixture of these nonmarine meteoric water and seawater. Rigaudier et al. (2011) calculated, from the isotopic composition of water trapped in fluid inclusions in Messinian halite, that the crystals grew in the mixture of seawater and meteoric waters dominated by the latter Geochemistry of Evaporites and Evolution of Seawater (50–75%). The proportion of the ionic components in basinal waters should not be significantly different than in seawater or ‘pure’ seawater brines especially in the early stages of evaporite basin evolution, unless some significant influx of highly saline waters, like hydrothermal brines, enters the basin (Hardie, 1990; Lowenstein and Risacher, 2009). Brine partially escapes from the system by seepage reflux, as in the McLeod salina, Australia (Logan, 1987), and by outflow bottom current, when the barrier separating the basin from the sea is below the sea level – as in the Kara Bogaz (Kosarev et al., 2009; Sonnenfeld, 1984). The presence of a brine reflux by underground seepage in case of a salina basin, and additionally also by return bottom current – in a lagoon basin, is the important feature of the model basin. Together with this brine, some of the dissolved salts are removed from the basin. The assumed brine reflux is the best and the simplest explanation for the ‘escape of salts’ from the system, which is a necessary condition for the deposition of ancient evaporite sequences that always seem to have different proportions of salts from those expected by complete evaporation of seawater in the hydrologically closed system (Hite, 1970; King, 1947; Klein-BenDavid et al., 2004). Long-term chemical evolution of the brine in some marginal marine basins can lead to brines having nonmarine characteristics (Klein-BenDavid et al., 2004). When inflowing continental or other nonmarine waters show ionic composition and/or ionic proportions extremely different from those of marine water for a sufficiently long time, or show relatively high salinity, mixing can change the initial marine proportions of ions and produce a mixed brine (Hardie, 1984, 1990). The ionic composition of a basinal brine changes with time principally because of the precipitation of successive evaporite minerals and various back and early diagenetic reactions with chemical sediments. During advancing evaporation, all these processes selectively remove particular ions from the parent brine (Valyashko, 1962). The other causes of chemical changes are dissolution and reprecipitation (recycling) of earlier or older salts (Holser, 1979a). In the case of variations in bottom relief, the brine can evolve in different ways in each subbasin or parts of the basin. The refreshment of brine is possible due to increased influx of both marine water and meteoric water (Holser, 1979a). In relatively wetter climates, brackish subbasins can develop on the landward side of a large evaporite basin, perhaps showing the peculiar chemical composition of the waters. They also can appear in the final stages of evolution of the basin, when the water level in the basin is at sea level and the influx of marine water is minimal (Kendall and Harwood, 1996). In summary, a marginal marine basin can show various types of brines depending on place and stage of basin evolution – from strictly marine to nonmarine, mixed, and even brackish, in the case of an interval of wetter climate. Many ancient evaporite basins commonly show sediments having both marine and nonmarine physical features, contain both rare marine and nonmarine fossils, and reveal geochemical characteristics of salts pointing to both marine and nonmarine derivation of the brine. As shown by Kirkland et al. (1995, 2000) and Denison et al. (1998), these contradictions are relatively simple to resolve, assuming that the basin was of a marginal marine type and the brine was neither exclusively marine nor exclusively nonmarine but was a mixture of marine and various nonmarine waters. 9.17.7.2 495 Quantitative Model of the Basin The marginal marine evaporite basin model (Figure 4) was developed quantitatively and subjected to numerical simulations (e.g., Logan, 1987; Nunn and Harris, 2007). The crucial idea, developed by Valiaev (1970), that the salinity evolution in the basin depends on the ratio of inflow to outflow was supplemented by Kopnin (1977) and Holser (1979a). It was later explored by Sonnenfeld (1980, 1984) and then fully developed by Sanford and Wood (1991) and Ayora et al. (1994) who integrated it with the model of evaporite precipitation of salts from seawater and related brines by Harvie et al. (1980). For the purpose of numerical modeling, the basin was assumed to evolve with a constant (conservative) water volume (this volume can also vary in some alternate models; Ayora et al., 1995). Thus, the evaporated and refluxing volume of water may be replaced by an equal volume of influx water that was made up of seawater, other water, or a mixture of the two (Figure 5). The inflowing water is presumed to be completely mixed with the basinal water. Atmospheric precipitation and evaporation were omitted in salt balance calculations because they usually do not carry any substantial amount of salts. The salts ‘escape’ from the system carried by refluxing water or being precipitated on the basin floor. The steady state of the basin – its constant volume – and the evolving salinity of basinal brines were considered as the basic condition for the accumulation of the particular sequences of evaporite minerals. The changes from one evaporite mineral sequence to another were controlled by the particular parameter called ‘degree of restriction’ or ‘restriction index’ (Cendón et al., CW MP E SW OC SP SR MP - Meteoric precipitation E - Evaporation SW - Seawater inflow current OC - Basinal water outflow current SR - Seepage outflow CW - Continental water inflow SP - Salt precipitates Atmospheric water, negligible % of salts Seawater, 3.5% of salts Continental water <<0.1% of salts Basinal water or brine, mixture of seawater, continental and atmospheric water Figure 5 Qualitative model of marginal marine evaporite basin used for numerical modeling of evaporite salt sequences, dependent on restriction index of the basin; after Sanford and Wood (1991) and Ayora et al. (1994). 496 Geochemistry of Evaporites and Evolution of Seawater 2003; Fanlo and Ayora, 1998) formerly described as the leakage ratio (Sanford and Wood, 1991). It is the value of water outflow relative to the total inflow or, in short, the ‘leakage’ to ‘inflow’ QL/QI, where QL is the outflow due to direct reflux to the sea by bottom current and leakage (seepage) to aquifers and QI is the total inflow (Ayora et al., 1994, 1995; Cendón et al., 2008; Sanford and Wood, 1991). The restriction index QL/QI ranges from zero in a completely closed basin to one for the open ocean. The predicted sequence and thickness of the evaporite mineral facies were related to the number of ‘evaporated basins’ (see Ayora et al., 1994, 1995; Sanford and Wood, 1991; Sonnenfeld, 1980, 1984 for detailed equations and explanations). In the basin model, the accumulation of a great thickness of an evaporite salt is possible when the basin is under a steady state regime, that is, preserving both a constant basin volume and solute concentration of the basinal brine (Ayora et al., 2001; Sanford and Wood, 1991). A steady state develops after some time for a given QL/QI ratio (the degree of restriction). Numerical models showed that the QL/QI ratio influences the type of salts precipitated (paragenesis) and the thickness of the particular mineral formed, while the chemical composition and proportion of the inflow waters influence the relative amount of solutes, and to a lesser extent, the paragenesis of salts (Ayora et al., 1994). Highly limited outflow or a completely restricted basin (QL/QI ¼ 0) causes an unrealistically low amount of the each salt to precipitate previous to the next soluble salt (Ayora et al., 1995). The numerical model thus developed has been used successfully to interpret the sequences of ancient halite and K–Mg evaporites in terms of the evolving basinal brines in marginal marine basins, being a mixture of marine and some nonmarine waters. In the marginal marine basin, the main source of salts is seawater, and therefore, both brine and evaporite salts should generally show the marine geochemical characters that are observed in recent coastal salinas (Logan, 1987). Salinity rise should lead to the deposition of evaporite salts following the Usiglio sequence (Stewart, 1963). Modeling, supported by field observations from the MacLeod basin, however, revealed that this model pathway depends on the inflow/outflow rate (seepage reflux in a salina basin; Figure 4(b)), which controls the maximum salinity level in the basin and thus limits the possibility of the precipitation of higher salts (Logan, 1987; Sonnenfeld, 1984; Valiaev, 1970; Wood and Sanford, 1990). Changes in the inflow/outflow rate influence the order of the precipitated salts and, for example, gypsum can be deposited after halite, as documented in the MacLeod salina (Kendall and Harwood, 1996; Kendall and Warren, 1988, Figure 2.35; Logan, 1987). The inflow/outflow rate also controls the thickness of deposited evaporite salts, and sometimes, there is only a small accumulation of gypsum before halite (Sanford and Wood, 1991). In the case of limited outflow, the Naþ, Kþ, Mg2þ, and Cl# ions are not involved in the initial Ca carbonate and Ca sulfate precipitation, and the SO42#, not fully used for Ca sulfate precipitation, can accumulate in the brine, and the basin then has a great potential for the deposition of NaCl and K–Mg salts (see Hite, 1970). Ayora and coworkers (Ayora et al., 2001) considered the influence of the chemistry of meteoric (nonmarine) water influx on the course of evaporitic precipitation in great detail. They assumed that in marine basins, QSW (seawater inflow in liters per time) is higher than QRW (continental water inflow, i.e., rivers and groundwater) and that cSW (the concentration, in moles per liter of solution, of the particular solutes in seawater) is one order of magnitude higher than cRW (the concentration of the solutes in continental waters), and therefore, the mass of the solutes precipitated from continental waters can be neglected in calculations. For example, Ayora et al. (1995) noted that the concentration of sulfate ions in river water is one order of magnitude lower than that of seawater and therefore its influence on isotopic composition (O, S) of basinal ‘marine’ water is not significant and can be neglected (see Claypool et al., 1980; Kirkland et al., 1995; Denison et al., 1998, for similar calculations). The influence of CaCl-rich hydrothermal brine influx on deposition in basins leading to precipitation of KCl-type evaporites was modeled by Cendón et al. (2003) and Garcı́a-Veigas et al. (2009). The influence of recycling processes on the sequence of crystallization in this basin type was further modeled by Cendόn et al. (2004). Cendón et al. (1998) explored the same model of the basin to predict the salt deposition in two interconnected subbasins containing the halite and potash deposits. The loss of salt components from the basin via the atmosphere in the form of salt-bearing aerosols or via aeolian transport of salts (by salt storms) is usually neglected in the models, although in nature they are well documented and may be quantitatively significant (Risacher et al., 2006; Wood and Sanford, 2007). The basin model has some unrealistic features and one of them is the assumption of the complete mixture of waters inflowing into it. In reality, the separation of brine bodies within the same basin is a typical feature of many evaporite basins. Shallow basins, such as Kara Bogaz or Bocana de Virrilá, show horizontal salinity gradients; deep basins, such as the Dead Sea, are commonly permanently stratified. The influx of seawater, flood, or river water into deep basins commonly leads to the establishment of a pycnocline and to specific depositional processes occurring in between the brine bodies (Ba˛bel and Bogucki, 2007; Holser, 1979a; Torfstein et al., 2005). As with any model that is an extreme simplification of reality – the numerical models thus far introduced and described earlier do not cover all the environmental processes observed acting in real basins, such as sulfate reduction, or interaction of the precipitated minerals with the solution, such as dehydration. Some of these are covered by models of closed (lake) basins (Sánchez-Moral et al., 2002; Yan et al., 2002). 9.17.8 Mode of Evaporite Deposition When an evaporating brine becomes supersaturated with respect to a particular salt, it can precipitate in any portion of the brine body. Some areas, however, are most favorable, particularly the bottom of the basin and the brine–air interface. Shoals and basin margins (bays and evaporite flats) are particularly advantageous. The shallow water has a smaller volume and thus warms more easily than the greater volume of deep water. Because of the Geochemistry of Evaporites and Evolution of Seawater higher temperature, the evaporation rate is more rapid on the shoals. Secondly, there is more surface area per unit of water volume in the shallows in comparison with the deep parts of the basin, and hence, even the same evaporation rate leads to faster salinity rise in these shallow areas of the basin (in small bays, evaporite flats, etc.). Thus, the salinity and concentration of ions are higher in shallow zones, leading to more rapid precipitation of salts (Cornée et al., 1992). Salt may precipitate (1) at the brine/air interface; (2) within the brine column, particularly at the pycnocline; (3) directly on the floor of the evaporite basin; and (4) in brine-soaked sediments or brine-soaked organic mats as displacive crystals or pore-filling cements (Logan, 1987; Schreiber, 1978). The accumulations of such crystals can create several more or less distinct genetic groups of deposits: (1) subaqueous crystal cumulates, (2) subaqueous bottom precipitates (bottomgrown crystals and crystal crusts), (3) intrasediment precipitates (incorporative, displacive, and replacive crystals and nodular aggregates), and (4) clastic accumulations (Figure 6; see Hanford, 1991; Kendall, 1992, 2010; Logan, 1987; Lowenstein, 1982; Schreiber et al., 1976; Smoot and Lowenstein, 1991). These genetic groups are best known from gypsum and halite precipitates. Bottom precipitates are commonly formed as firmly cemented, interlocking, orderly crystal crusts. The specific feature of gypsum deposits is the growth of extremely large selenite crystals in such layers. The examples include 4.5-m-long gypsum Clastic halite twinned crystals from the Messinian of the Cyprus and Sicily and 3.5-m-long gypsum crystals from the Badenian of Poland (Ba˛bel et al., 2010, with references; see a comment in Lugli et al., 2010, p. 94). Similar sizes can also be reached in secondary (diagenetic) salt crystals (halite and gypsum) growing in some synsedimentary karst cavities within evaporite sediments (Dı́az et al., 1999). Gypsum crystals commonly form twins and are able to create complicated crystalline structures, showing different morphology within each sedimentary subfacies, and are specific for each environment within any one basin (Ayllón-Quevedo et al., 2007; Ba˛bel et al., 2010; Lugli et al., 2010; Ortı́, 2011; Rodrı́guez-Aranda et al., 1995). Many morphologies and sedimentary sequences, however, seem to be repeated from basin to basin and are recognizable even when moving from primary gypsum morphologies to subsurface anhydrite. One of the most significant factors, which control gypsum morphology and facies distribution, appears to be the presence of microbial communities and the organic compounds they produce in the basinal brine (Oren, 2010). Additionally, some evidence indicates that the bottom-grown selenite crystals may grow in a regionally oriented manner under the influence of brine currents (Ba˛bel and Becker, 2006; Ba˛bel and Bogucki, 2007; Lugli et al., 2010). Such region-wide currents may even shape the selenite–gypsum microbialite domes in some areas (Ba˛bel et al., 2011) similar to the elongated elliptical halite ridges growing in more saline brine currents, as in the Dead Sea Floating rafts Bottom-grown halite Gypsum microbialites HALITE Microbialites Sabkha intrasediment precipitates 497 Crystal rafts Clastic Cumulates Selenite domes Bottom grown crystals GYPSUM Selenite crust Figure 6 Modes of evaporite deposition, the idea for the diagram borrowed from Kendall (2010). 498 Geochemistry of Evaporites and Evolution of Seawater (Karcz and Zak, 1987). Such domal structures may range in size from a few centimeter–decimeter to several meters and are common in selenite crusts (Ortı́ et al., 1984a; Warren, 1982), but oriented growth forms also are noted in some halite deposits (Talbot et al., 1996). Halite crystals growing in bottom crusts commonly show millimeter-scale zoning arranged into chevron-like pattern reflecting the upward cyclic accretion of the faces of cube. Some of the growth zones are diurnal and several growth zones can form in 1 day (Roberts and Spencer, 1995). These rapidly growing cube faces typically trap the brine inclusions that commonly are the target of geochemical studies (Figure 15). The precipitation of salts at the brine–air interface results in the development of floating crystals or crystal rafts. The most characteristic and common are halite ‘boat-like’ or ‘hopper’ crystals (Arthurton, 1973; Hanford, 1991; Valyashko, 1951). Because of their rapid growth rate, they are able to trap copious amounts of fluid inclusions (Roberts and Spencer, 1995). Similar rafts are known from carnallite as well as gypsum crystals (Chivas, 2007; Ortı́ et al., 1984a; Talbot et al., 1996). Floating single crystals coalesce and form rafts. Sunken crystals and rafts create specific deposits known as cumulates (Lowenstein, 1982; Lowenstein and Hardie, 1985; Shearman, 1978). Halite commonly grows faster in the near surface zone of brine mainly due to night cooling effect of the NaClsaturated brine heated during solar evaporite concentration in the daytime (but sometimes also due to heating; Karcz and Zak, 1987). Because of accelerated growth in this zone, the specific crystalline, mounded structures may form (salt mushrooms, e.g., Ganor and Katz, 1989). Similarly, salt umbrellas form when the crystal growth is associated with water level fall (Müller, 1969). Sometimes, pillar- or atoll-like structures several meters in size are formed (Talbot et al., 1996). The other spectacular but rare form of salt mineral crystallization is in high-energy environments as growth of accretionary (coated) grains that are termed pisoids or ooids when they are more rounded in shape. Mirabilite, halite, and gypsum ooids and pisoids are known from both modern and ancient deposits (Ba˛bel and Kasprzyk, 1990; Castanier et al., 1999; Tekin et al., 2007, 2008; Weiler et al., 1974). Evaporite deposits, and particularly gypsum, commonly form in the presence of microbial (cyanobacterial) mats. They create microbialite domal structures (analogous to carbonate microbialites) that, however, are complex hybrid structures presumably more inorganic (chemical) than organic (microbial) in origin (Ba˛bel et al., 2011; Petrash et al., 2012; Riding, 2008; Rouchy and Monty, 1981, 2000; Vogel et al., 2010). 9.17.9 Primary and Secondary Evaporites The high solubility of evaporite salts and their halokinetic properties made them chemically and physically very mobile material both in the sedimentary environment and particularly during diagenesis and burial. They are easily dissolved, and rather easily ‘recrystallized,’ as well as able to easily replace one another forming new minerals during burial history. The hydrous salts can be commonly dehydrated and rehydrated in the sedimentary environment, during burial, exhumation, and weathering (Jowett et al., 1993; Testa and Lugli, 2000). The most common mode of cementation of an evaporite sediment is via the formation of syntaxial overgrowths that may blur the distinction of the grain from the cement. The reconstruction of the diagenetic evolution of most evaporite rocks, which have undergone the burial–exhumation cycle, is therefore difficult. The basic problem is the proper pathway for the reconstruction of the sequence and time(s) of the petrological, mineralogical, and textural–structural changes and to find the criteria for recognition of the primary feature of the evaporite rock. Dronkert (1985, p. 94) following Ingerson (1968), defined that the primary evaporite minerals are those that “precipitated directly from the solution” whereas secondary minerals “formed later than the primary ones and at least in large part from them” and suggested 17 geochemical and structural criteria for their distinction. It is well known that many evaporite sediments can be transformed very early (replaced by more stable mineral association), just after they are formed in the basin, and therefore the term primary should be understood in the broader way – it should include depositional as well as postdepositional but preburial processes (Hardie et al., 1985). The synsedimentary alteration of epsomite to bloedite in the Quero Lake (Spain) is one clear example (Sánchez-Moral et al., 1998). Indeed, Braitsch (1971) defined the term ‘primary precipitation’ in an extremely extended way including “the (early) diagenetic alterations of metastable to stable precipitates” (Braitsch, 1971, p. 92). He justified that “from the standpoint of the conditions of formation,” “no changes are necessary in the parameters such as temperature, concentration etc. for the onset of stable equilibria” except of “the adjustment of the activation energy necessary for the onset of stable equilibria” (Braitsch, 1971, p. 92). Thus, depending on the time of formation, the minerals and fabric (including fluid inclusions) of evaporite deposits can be subdivided into three main groupings (Hardie et al., 1985, p. 12): 1. “Depositional, i.e., formed at the time of deposition of a sedimentation unit or deposited in its existing form.” 2. “Post-depositional but pre-burial, i.e., formed diagenetically soon after deposition by processes controlled by the existing depositional environment.” This second stage of formation is equivalent of the ‘eogenetic stage’ distinguished and defined by Choquette and Pray (1970, p. 219) as “the time interval between final deposition and burial of the newly deposited sediment or rock below the depth of significant influence by processes that either operate from the surface or depend for their effectiveness on proximity to the surface.” The lower limit of the eogenetic zone was defined at “that point at which surface recharged meteoric waters, or normal (or evaporated) marine waters, cease to actively circulate by gravity or convection” (Moore, 1989, p. 25). 3. Post-burial, i.e., formed by late diagenetic or metamorphicmetasomatic processes controlled by the subsurface burial environment” (Hardie et al., 1985, p. 12). Geochemistry of Evaporites and Evolution of Seawater The criteria for syndepositional features include “(1) mechanical sedimentary structures and detrital textures and fabrics produced during traction and suspension load deposition of chemical sediment particles; (2) crystalline textures and fabrics produced as chemically precipitated minerals grew in place on and within bottom sediment; and (3) features indicative of contemporaneous cementation, dissolution and reprecipitation of salts. Additional criteria come from fluid inclusions and mineral stability ranges” (Hardie et al., 1985, p. 13). These and other criteria for distinguishing the primary and secondary crystals in evaporites are listed elsewhere (Ba˛bel and Becker, 2006; Hardie, 1984; Holser, 1966; Spencer, 2000). Some criteria for the primary nature of crystals growing on the substrate in nonevaporite environments are also useful for evaporites (Dejonghe, 1990; Kendall and Iannace, 2001; Sumner and Grotzinger, 2000). Extraordinary preservation of layering may indicate the primary origin of salts (Braitsch, 1971); however, secondary structures present in deformed evaporites commonly mimic primary features, making distinction of primary from secondary evaporites very difficult (Schreiber and Helman, 2005; Warren, 2006). Hardie (1984, p. 201) further subdivided the evaporites into the following types: 1. The ‘primary’ evaporites or ‘modified primary evaporites’ – which are “not sufficiently altered by burial metamorphism or metasomatism to hide the identity of the primary (¼syndepositional) mineral assemblages” 2. The ‘secondary’ evaporites – “so altered after burial that the primary minerals cannot be unambiguously identified” For geochemical studies, the crucial theme of this chapter, only the primary autochthonous evaporites that are precipitated in place can give reliable information about the chemistry of the evaporite basin water (Hardie, 1984). Evaporites can occur both as primary and secondary minerals in ancient deposits, although some of them are more typical as primary, the others more common as secondary. Of the common potash minerals, sylvite (KCl), the most common component of the marine K–Mg salts (Dean, 1978; Garrett, 1970; Holser, 1979a; Stewart, 1963), was commonly interpreted as the product of replacement or incongruent dissolution of primary carnallite (KCl • MgCl2 • 6H2O), which could take place syndepositionary (e.g., El Tabakh et al., 1999a; Richter-Bernburg, 1972) or during exhumation (retrograde diagenesis; Harville and Fritz, 1986). The following reaction of incongruent dissolution of carnallite (after Hardie, 1984) is considered as the most important in diagenesis of potash evaporites • (Braitsch, 1971): KCl , MgCl2 , 6H2 OðsÞ ! KClðsÞ þ Mg2þðaqÞ þ 2Cl# ðaqÞ þ 6H2 OðlÞ [5] (where s ¼ solid, aq ¼ aqueous, or soluble in water, and l ¼ liquid). By utilizing a similar reaction, sylvite is produced commercially from carnallite by dissolution in water (Fokker et al., 2000). Recently, an increasing number of reports revealed the primary, synsedimentary nature of ancient sylvite and carnallite deposits (e.g., Cendón et al., 1998; Lowenstein and Spencer, 499 1990; Rahimpour-Bonab et al., 2007). Both ancient sylvite and carnallite crystals show structures typical of the growth on the bottom of evaporite basin (Kendall, 2010; Wardlaw, 1972a,b). 9.17.10 Evaporation of Seawater – Experimental Approach The crucial concept for an understanding of the geochemistry of marine evaporites is an understanding of the process of evaporation of seawater leading to salinity increase and evaporitic concentration of particular ions, followed by the ordered precipitation of particular salts. The first reported experiment of complete evaporation of seawater using a marine water sample, taken on the French coast of the Mediterranean, was by Usiglio (1849a,b). This water sample had a 38.45% salinity, with an associated air temperature of 40 " C. Usiglio described the process and order of salts precipitated in a quantitative way and this order is now called the Usiglio sequence (see, e.g., Logan, 1987). The experiment was repeated by Bassegio (1974) and McCaffrey et al. (1987), among others. The early stages of evaporation leading to gypsum and halite precipitation are particularly well known and are observed in many marine solar saltworks (Busson et al., 1982; Geisler-Cussey, 1986; Herrmann et al., 1973; Ortı́ and Busson, 1984). All these empirical observations were made in slightly different and fluctuating temperature (air and brine), but all showed coincident results, particularly during the early stages of evaporation (Table 5). Some differences and inconsistencies, mainly related to the influence of temperature differences, appear in the later stages of the precipitation of K–Mg salts from highly concentrated brine (Garrett, 1970). The final stages of evaporation and precipitation of K–Mg salts are known mainly from experiments and theoretical calculations (for a review of the experiments, see Braitsch, 1971). 9.17.11 Crystallization Sequence before K–Mg Salt Precipitation The sequence of crystallization up to saturation with K–Mg salts is well known from solar saltworks, where seawater passes through three stages or fields characterized by precipitation of calcium carbonate, calcium sulfate, and sodium chloride (Figure 7). 9.17.11.1 Early Salinity Rise – Calcium Carbonate Precipitation The calcium carbonate field is the first field of elevated salinity (>35%, seawater density: 1.0258 g cm#3 at 12 " C) up to the salinity characterizing the first precipitation of gypsum (i.e., ! 140–200%, seawater brine density: 1.11–1.13). Logan (1987), similar to Usiglio (1849a), established that aragonite started to precipitate at the volume reduction ratio Ver ¼ 0.5. Usiglio recorded the end of CaCO3 precipitation within the field of gypsum. A common phenomenon within the Ca carbonate field is the appearance of thick microbial (cyanobacterial) mats at salinities greater than 110% (S ¼ 110 g kg#1 500 Stages and brine types Volumetric mass (density, kg m#3 TDS (g l#1) Cl (mg l#1) (mmol kg#1) SO4 (mg l#1) (mmol kg#1) Na (mg l#1) (mmol kg#1) Mg (mg l#1) (mmol kg#1) Ca (mg l#1) (mmol kg#1) K (mg l#1) (mmol kg#1) Mg (mg l#1) (mmol kg#1) Br (mg l#1) (mmol kg#1) 0.0 Seawater 1.0 Gypsum beginning 2.0 Halite beginning 2.1 Halite 2.2 Halite 2.3 Halite 3.0 Epsomite beginning 4.0 Sylvite beginning 5.0 Carnallite beginning 5.1 Carnallite 6.0 Bischofite beginning 1.022 35.8 1.084 124.7 1.204 307.9 1.220 334.4 1.247 332.0 1.238 383.8 1.286 400.2 410.3 1.305 418.2 1.325 462.6 1.364 504.8 2770 29.2 10 100 110 19 100 222 28 900 339 36 400 414 65 400 797 82 200 966 56 100 664 35 400 416 27 100 327 34 900 423 11 000 485.2 37 800 1714 95 100 4616 89 000 4371 65 600 3119 63 000 3200 48 200 2367 22 100 1093 15 000 723 8150 412 1680 85 1320 55.1 4530 194 13 400 615 20 900 971 35 500 1596 50 500 2432 56 120 2607 72 900 3410 85 700 3976 108 800 5186 122 000 5841 420 10.6 1540 40.1 450 12.5 237 6.68 170 4.64 96 2.81 tr 1.290 19 780 565.7 69 000 2029 175 600 5527 188 200 5994 185 200 5709 189 900 6271 190 500 6066 223 900 7179 257 600 8194 304 600 9953 337 300 11 074 408 10.6 1470 39.2 3600 103 5300 153 7730 216 12 900 386 17 680 510 25 900 753 17 000 490 860 25.5 860 25.6 1320 55.1 4530 194 13 400 615 20 900 971 35 500 1596 50 500 2432 56 120 2607 72 900 3410 85 700 3976 108 800 5186 122 000 5841 68 0.86 234 3.05 578 8.07 950 13.4 1327 18.2 1830 26.8 2970 41.9 4770 67.9 5300 74.8 7380 107 7530 110 tr, traces; ‘-‘, no data. tr – – 60 1.74 Geochemistry of Evaporites and Evolution of Seawater Table 5 Evolution of evaporating modern seawater brines, based on data from semi-natural (saltworks), natural, and experimental evaporation, compiled and averaged data from many sources repeated after Fontes and Matray (1993) Geochemistry of Evaporites and Evolution of Seawater 35‰–(140-200)‰ 1.03–(1.10-1.13) (140-200)‰–(290-325)‰ g cm−3 501 (290-325)‰–375‰ (1.10-1.13)–(1.20-1.26) g cm −3 (1.20-1.26)–(1.32) g cm−3 Salinity, density Initial evaporation pans Seawater Gypsum pans Halite pans 35‰ Seawater brine >370‰ CaCO3 CaSO4·2H2O NaCl Remaining SO42- -rich brine 35‰ 1.0258 g cm−3 Sea (at temp. 12 °C) Figure 7 Scheme of the marine saltwork pan after Ortı́ et al., 1984a,b, modified. Remaining SO42#-rich brine flows back to the ocean or, in some saltworks, back to the concentration pans to promote more precipitation of gypsum, which crystallizes due to a mixture of brines (Raup, 1982). and density 1.087 g cm#3; Segal et al., 2006) that dominate until 150% or even higher, where they cease with the onset of the precipitation of gypsum (references in Ba˛bel, 2004a). The photosynthetic activity of cyanobacteria raises the content of dissolved oxygen that shows daily fluctuations (up to 7.8 mg l#1, 131% supersaturation), particularly in the zones where accumulations of O2-rich bubbles are seen on the surface of the mats (Cornée et al., 1992). The greatest level of calcium carbonate productivity was observed in salinities between 50 and 70 g l#1 and the mineral formed was Mg calcite, with minor additions of calcite and aragonite (Ortı́ et al., 1984a). The amount of CaCO3 precipitated during evaporative concentration is negligible in comparison with the ensuing salts (gypsum, halite, and K–Mg salts). In natural environments, except for calcium carbonate precipitation that is induced by evaporation, several other nonevaporite driving mechanisms for CaCO3 deposition commonly operate within the basin. These mechanisms can be more important and can deposit a large amount of carbonate when the basin stays within the field of carbonate salinity for a long period of time while being constantly supplied with inflowing seawater (see, e.g., Decima et al., 1988). 9.17.11.2 Gypsum Crystallization Field This field ranges from the start of gypsum crystallization with the volume reduction in total water having a ratio Ver ¼ 0.2 (Logan, 1987) or 0.19 (Usiglio, 1849b), beginning at !150% and continuing up to beginning of the halite crystallization at salinity !290–320% (seawater brine density 1.20–1.26). Minor amounts of gypsum form within the lower portion of the halite field because the fields of crystallization overlap. In the lower end of the gypsum field, the first gypsum usually is a fine-grained precipitate; in more concentrated waters, it forms firm coarser-crystalline crusts commonly displaying the centimeter-to-decimeter large domal structures. When crystals show sizes larger than 2 mm, they are commonly called selenite (Warren, 1982). The interesting feature of the ‘selenite’ field is that mat-creating cyanobacterial communities do not only inhabit the sediment/water interface but actually also live within the selenite crust. They grow and remain within the photic zone, where transparent selenite crystals play a role comparable to light channels, forming endoevaporitic microbial mats (Canfield et al., 2004). The presence of a complex, living microbial community, particularly cyanobacteria, within gypsum sediments profoundly influences the geochemical microenvironment, leading, for example, to increased amounts of photosynthetically produced oxygen (up to concentration equal four times air saturation during the day), but that oxygen remains within the interstitial brine. 9.17.11.3 Halite Crystallization Field At the beginning of the halite crystallization, small amounts of gypsum still form, usually as tiny needlelike crystals, intermixed with minuscule halite cubes. Halite begins to crystallize when the standard seawater is evaporated to 0.09–0.1 of the original volume, that is, at a volume reduction ratio Ver ¼ 0.09 (Logan, 1987) or 0.095 (Usiglio, 1849b). Halite, unlike almost all other common sedimentary minerals, requires relatively low degree of supersaturation to begin precipitation (Berner, 1971). For example, in the Adriatic solar saltworks, nearly twofold supersaturation was necessary for the first gypsum to precipitate and only Geochemistry of Evaporites and Evolution of Seawater 1.3 times supersaturation for the initial halite (Herrmann et al., 1973). The halite crystallization continues up to very high salinities, passing the point where the first magnesium sulfate crystallizes (together with the halite) at salinity ! 375% and brine density 1.32. During seawater evaporation extending up to this stage, the concentrations of major ions systematically change in a predictable way – well known from geochemical studies in solar saltworks and experimental evaporation of seawater (Figure 8; e.g., Geisler-Cussey, 1997; Levy, 1977). Concentration of K+, Na+, Mg2+, Cl-, SO42- (mMol kg−1-H2O) 8000 100 Start of epsomite precipitation Ca2+ BrLi+ Start of gypsum precipitation 90 Start of halite precipitation 80 6000 Start of carnallite precipitation ClNa+ SO42Mg2+ K+ 4000 70 60 50 40 Start of kainite precipitation 2000 30 20 Concentration of Ca2+, Br-, and Li+ (mMol kg−1-H2O) 502 10 0 0 10 20 30 40 50 60 70 80 90 0 100 Degree of evaporation 1400 1350 Start of kainite precipitation Start of halite precipitation 1300 Density (g cm−3) Start of epsomite precipitation Start of gypsum precipitation 1250 1200 Start of carnallite precipitation 1150 1100 1050 1000 0 10 20 30 40 50 60 70 80 90 100 Degree of evaporation Figure 8 Major and minor ion concentrations and density rise in evaporating Caribbean seawater, after McCaffrey et al. (1987) and Warren (2006), modified. Degree of evaporation based on Mg2þ and Liþ, after McCaffrey et al. (1987). Geochemistry of Evaporites and Evolution of Seawater 9.17.12 Crystallization Sequence of K–Mg Salts The sequence of evaporite crystallization of marine K–Mg salts is known from empirical observations of the evaporating seawater brines and from laboratory and theoretical chemical studies of saline solutions. 9.17.12.1 Natural Crystallization Complete natural evaporation of seawater was rarely monitored to total dryness and/or with a necessary level of precision. The crystallization sequence strongly depends on many environmental factors, and the main ones are the temperature (both of the brine and the air) and its variations, fluctuations in humidity, rate of evaporation, rate of precipitation, depth of water, and even small deviations from standard chemical composition of the inflowing seawater brines recorded in particular regions (Garrett, 1970, 1996; Jadhav, 1985; Krauskopf, 1967; Valyashko, 1962). The natural crystallization is always polythermal. The temperature of bitterns in saltworks is most commonly between 18 and 35 " C but can reach 50 " C in the most concentrated bitterns and may drop to 5 " C, as recorded in the winter in France (Charuit and Genty, 1980; Jadhav, 1985). The natural crystallization in saltworks follows the so-called equilibrium mode crystallization (Bea et al., 2010), in which bitterns remain in permanent contact with all previously precipitated solids. During fractional crystallization (Harvie et al., 1980), mostly known from theoretical models and laboratory studies, all the precipitates are assumed to be removed from the bittern, although in fact they are also in contact with this bittern at the moment of their formation. As documented in seawater evaporation processes in coastal saltwork pans around the world, the next salt, which appears in the course of crystallization, is invariably epsomite (Cohen-Adad et al., 2002; Ortı́ et al., 1984a; Valyashko, 1962), although mirabilite may precede crystallization of that mineral, during winter or throughout evening cooling of the brine (Garrett, 1970). Interestingly, some of the detailed studies of salinas along the Mediterranean (in Spain and France) do not report mirabilite (Geisler, 1982; Ortı́ et al., 1984a,b) although the study by Charuit and Genty (1980) does. During further evaporation, water studied from evaporating ponds along the Black Sea, epsomite was transformed into sakiite (hexahydrite), which also crystallized in the primary form (Valyashko, 1962). Sakiite crystallization together with epsomite and halite was also recorded in India (Chitnis and Sanghavi, 1993). During further evaporation of the Black Sea water, the carnallite crystallization was joined to the crystallizing salts, and finally bischofite was crystallized together with these salts, up to the end of evaporation (Valyashko, 1962). During experimental evaporation of the Black Sea water at 25 " C, Valyashko (1962, p. 160) observed that sylvite started to crystallize nearly simultaneously with sakiite, and it continued the crystallization up to the start of carnallite precipitation. This result appears to support the early finding by van’t Hoff and Meyerhoffer (1899, cited by Balarew, 1993), later abandoned, that sylvite, instead of kainite, is obtained during seawater evaporation. Valyashko (1962) noted also that sylvite crystallized during cooling of the bittern (see also Garrett, 1970). Valyashko (1962) warned that sylvite may be unnoticed in 503 seawater precipitates, because of its great similarity to halite – both minerals crystallize together. The Crimean lake Saki, where the crystallization experiments were conducted, was supplied with seawater by seepage through the sandy bar and was slightly impoverished with respect to Kþ in relation to the Black Sea water (Valyashko, 1962). The brackish Black Sea water (!17% in the surface waters) is also slightly depleted in Kþ and enriched in Ca2þ in relation to the open ocean water (Carpenter, 1978). Therefore, these results are not exactly representative of the evaporation of standard oceanic water (McCaffrey et al., 1987, their Figures 3-8). The evaporating Mediterranean bitterns from saltworks of France gave the following sequence of K–Mg salts at a natural range of changing temperatures 28–35 " C: epsomite alone, then epsomite and kainite, then kainite, and, finally, kainite in association with bischofite and/or carnallite (Charuit and Genty, 1980). Krauskopf (1967) suggested that kainite (and kieserite) forms only when the rate of evaporation is sufficiently slow and the brine and precipitating salts stand together for a long time. These and associated sulfate salts (e.g., leonite) were found to crystallize from supersaturated solutions slowly and with difficulty (Bergman and Luzhnaya, 1951; Hadzeriga, 1967; Hardie, 1984, p. 207). Kainite crystallization from the Black Sea bittern was recorded only in one experiment, during a period of very slow evaporation (Il’insky 1948 in Valyashko, 1962), and it was considered as a secondary salt by Valyashko (1962). On the other hand, in the Indian saltworks, kainite joins the crystallization of epsomite, sakiite, and halite before the start of carnallite precipitation and is present volumetrically as the most significant precipitate at that interval of the crystallization path (Chitnis and Sanghavi, 1993; Garrett, 1970). In the Indian saltworks, carnallite crystallizes in the density interval 1.29–1.33 (sp. gr.) and with bischofite, being the final product of the seawater crystallization path, in the interval 1.33–1.37 (sp. gr.) (Jadhav, 1985). In the French Mediterranean saltworks, bischofite begins to crystallize from brine of the density 1.364 (Fontes and Matray, 1993). Kieserite was found crystallizing together with bischofite at the final desiccation stage of seawater in India (Chitnis and Sanghavi, 1993). Copious crystallization of bischofite was recorded as the mass of floating feather- and needlelike crystals in the density interval 1.370–1.377 (sp. gr.) at temperature 38–43.5 " C (Jadhav, 1985). Further concentration of the dense brine by solar evaporation, above the 1.377 specific gravity, was not possible due to the high viscosity of the bittern and absorption of atmospheric moisture (Jadhav, 1985). The bischofite began to dissolve when the bittern was heated over 44 " C and disappeared at 50 " C. In India, night temperature drops down to 10 " C led to precipitation of epsomite from the bitterns that at 30 " C are unsaturated with this salt, showing solubility strongly dependent on temperature (Chitnis and Sanghavi, 1993). Winter cooling of seawater bitterns is utilized for commercial production of epsomite in France. K–Mg salt crystallization is accompanied by halite, which ceases to precipitate when Na ions became exhausted from the bittern before final precipitation of bischofite (Amdouni, 2000). By heating, cooling, and freezing of the brine, the mixing of bittern from various evaporation stages, addition of bittern to formerly precipitated salts, dilution of the bittern by seawater, and other seminatural operations, a number of other salts can 504 Geochemistry of Evaporites and Evolution of Seawater precipitate in solar saltworks, including mirabilite and glauberite (Garrett, 1980, 1996; Hardie, 1985). The mixing of seawater brines leads also to precipitation of gypsum, halite (Ortı́ et al., 1984a; Raup, 1970, 1982), and sylvite and tachyhydrite (Wali, 2000). The maximum recorded density of evaporating seawater brine was 1.377 (sp. gr.) at 30 " C (Buch et al., 1993; Jadhav, 1985), and a similar artificial seawater solution – 1.339 – was created in experiments by Lychnos et al. (2010). In the highest salinity pans, the complete segregation of solids from liquid is practically impossible, and all the apparently solid phases should be considered as containing !10–20% of mother liquor (Hadzeriga, 1964). The natural sequences, particularly those from the coast of the Black Sea, are simpler than the sequences predicted by theoretical studies of mineral solubilities and numerical simulations (Braitsch, 1971; Eugster et al., 1980; Valyashko, 1962). 9.17.12.2 Theoretical Crystallization Paths The theoretical sequences of crystallization of marine K–Mg salts were established from solubility studies and determination of saturation points of these salts developed by Jacobus Henricus vant’Hoff (the first Nobel prize winner in chemistry in 1901), and his students, as well as from numerical calculations based mainly on the Pitzer ion-interaction model (Al-Droubi et al., 1980; Eugster, 1971; Harvie et al., 1980; see review by Bea et al., 2010). The sequences strongly depend not only on temperature but also on the kinetic factors of nucleation and crystallization (in particular, whether stable or unstable mineral equilibria prevail) and on whether the reactions between evolving brine and earlier formed salts actually took place. Therefore, several alternative models and paths of evaporative crystallization must be considered and at least several theoretical sequences are possible. Theoretical sequences are usually calculated for isothermal evaporation. A simulation of polythermal evaporation, for Quero Lake (Spain), was made by Sánchez-Moral et al. (1998). The theoretical stratigraphic columns showing both order and thickness of the various sequences of marine salts predicted for isothermal evaporation have been calculated and drawn by Braitsch (1971) and then by Braitsch and Kinsman (1978). Usually, the temperature sequence for 25 " C is used for comparison with the natural sequences. All the crystallization paths, including the natural sequence produced during temperature fluctuations typical of the Crimean coast of the Black Sea described earlier, can be subdivided into five steps or stages (Braitsch, 1971). In some theoretical sequences (as in the natural sequence described earlier), some steps are lacking. The complete set of steps is the following (also see Table 6): (1) Precipitation before saturation with respect of salts of the five-component system, which encloses Ca carbonate, Ca sulfate, and Na chloride stages of crystallization (2) Precipitation of Mg or Na–Mg sulfates without K salts (without sylvite and carnallite) (3) Precipitation of K–Mg salts (particularly sylvite), without carnallite (4) Precipitation of carnallite (5) Terminal precipitation with bischofite These stages are known also as the halite-, bloedite/epsomite-, kainite-, carnallite-, and bischofite-dominant stages, respectively (Eggenkamp et al., 1995). During experimental precipitation of seawater salts, the first Mg sulfates at the A–B boundary begin to crystallize when the brine is 70 times more concentrated than seawater, K-bearing salts (the B–C boundary) when it reaches 90 times the initial concentration (McCaffrey et al., 1987). Evaporation to dryness leads to the final point, “where the solution evaporates at a constant composition” (Usdowski and Dietzel, 1998, p. 70) and is simultaneously saturated with respect to all (at least two or more) dissolved solutes. This final point or state is called eutonic or drying-up (Borchert and Muir, 1964; Mullin, 2001; Sonnenfeld, 1984; Usdowski and Dietzel, 1998; Valyashko, 1962). The term eutonic coined by Kurnakov and Zhemchuzhnii in 1920 (Gamsjäger et al., 2008; Kurnakow and Žemčužny, 1924), is used not only to describe the point at saturation diagrams where the evaporation to dryness ends at the given invariant temperature and pressure (Figures 10-12) but also, less formally, to describe the final processes of crystallization or final composition of the naturally evaporating solutions. The described crystallization path concerns the present-day seawater, which evaporates in closed system without the continuous addition of fresh seawater. Such an addition, if present, could introduce calcium into the system, equally as the other ions, and this could certainly lead to precipitation of the other modified mineral suites. The predicted minerals crystallizing together with halite, and as the next solids formed after halite, would be polyhalite and glauberite (Holser, 1979a). The crystallization path of evaporated seawater, which includes the original presence of calcium, was successfully modeled by computer program written by Harvie et al. (1980). In this program (applied widely for studies of brines from halite inclusions), polyhalite is the expected mineral phase on the crystallization path that runs differently than the path predicted by solubility studies of five- Table 6 Subdivision of the crystallization sequence of the evaporating seawater into geochemical zones after various authors, increase in degree of evaporation and concentration is from 1 to 5 5 4 3 2 1 Van’ t Hoff school as summarized by Braitsch (1962, in 1971 edition) E – terminal precipitation with bischofite D – precipitation of carnallite C – precipitation of KMg-salts, without carnallite B – precipitation of NaMg or Mg sulfates without K-salts A – precipitation before saturation with respect of salts of the five-component system Zones after Valyashko (1972b) Bischofite zone Carnallite zone Sylvite zone Magnesium sulfate zone Zone of halite and zone of gypsumanhydrite Zones after Hardie (1984) MgCl2 KCl KCl MgSO4 CaSO4 Geochemistry of Evaporites and Evolution of Seawater component system of K–Mg salts described earlier (Hardie, 1984). This program is flexible and permits analysis of crystallization paths of modified marine waters of various compositions, including ancient seawaters, which showed crystallization paths different from the path of modern seawater described earlier. 9.17.13 Isotopic Effects in Evaporating Seawater Brines and Evaporite Salts 16 During evaporation, more of the light water molecules ( O as compared to the 18O) are selectively removed from the liquid, passing into vapor, leading to enrichment of the remaining liquid water in heavier isotopes (18O and 2H – deuterium: D). During evaporation of seawater to the stage of halite crystallization, parallel to this process, oxygen in sulfate ions is enriched in 18O (Pierre, 1985; Pierre et al., 1984b). Evaporation can also cause enrichment in heavier 13C (in relation to 12C) isotope (Potter et al., 2004; Stiller et al., 1985; Valero-Garcés et al., 1999). However, further evaporation of seawater brine from the start of halite precipitation leads to reverse effect, that is, depletion of heavier isotopes in the water (Valyashko et al., 1977), reflected in characteristic ‘evaporite loop’ on the dD–d18O plot (Holser, 1979b, 1992. The reverse effect is caused by rising salinity causing a progressive decrease in activity of water molecules in evaporating saline solutions and hydration of ions (Pierre, 1988, 1989). The paths of particular evaporite loops (for a given water reservoir) depend on the humidity, the isotope ratios of D and O in the air, and the cation composition of the basinal brine, which reduces the activity of H2O and its isotopic components (Sofer and Gat, 1975 ; Horita, 2005). Water molecules are used for the hydration of ions present in the brine, which introduces additional fractionation effects in the water, specific to each ionic species (Gat, 2010; Pierre, 1988). Evaporated seawater brines mixed with meteoric or other waters may have biased dD and d18O values due to mixture of H2O with different isotopic characteristic (Duane et al., 2004) and can be studied in fluid inclusions in evaporite minerals (e.g., halite; Knauth and Beeunas, 1986; Knauth and Roberts, 1991). H and O in water molecules in hydrated salts precipitated from these waters (such as gypsum) are subjected to isotopic fractionation and can be used for the interpretation of the origin of the basinal and other waters when the range of fractionation is known from experiments (Holser, 1979b; Koehler and Kyser, 1996). The same isotopic effects are recorded in evaporated nonmarine brines (e.g., Cartwright et al., 2009). Gypsum, which is most common among hydrated marine salts, is potentially an excellent marker recording the derivation of H2O in brines (Buck and Van Hoesen, 2005; Farpoor et al., 2004, 2011; Hodell et al., 2012). However, special laboratory techniques to exclude ‘nonhydration’ water (moisture and adsorbed water) from the analysis without any loss of hydration water are required to obtain the proper results (Playà et al., 2005; Rohrssen et al., 2008). Isotopic processes in brines are specific and can influence the isotopic signal leading to erroneous interpretations (Schreiber and El Tabakh, 2000). In saline waters and brines, at salinity and concentrations exceeding those of seawater, the isotopic fractionation processes are influenced by ‘salinity effects’ (Gat, 1995, 2010; Horita, 2005, 2009; Koehler and Kyser, 1996). 505 Commonly, brines are permanently stratified, which can lead to separate isotopic composition of particular brine bodies. The isotopic signals of evaporite deposits are complicated by this stratification effect and require a special approach to resolve some isotopic imbalance problems (e.g., ‘sulfur pump’ mechanism proposed by Torfstein et al., 2005). 9.17.14 Usiglio Sequence – A Summary During the evaporation of seawater, the salinity and density of water increases as well as the concentration of particular ions, leading to saturation, supersaturation, and precipitation of particular compounds. The general rule is that first compounds that precipitate are less soluble, that is, calcium carbonate (calcite and aragonite), calcium sulfate dihydrate (gypsum), and sodium chloride (halite), so the order of precipitation reflects the rise in solubility (Table 4). The composition of brine becomes simpler evolving from the initial seven-component system toward the five-component system in the field of K–Mg salt precipitation (Table 5 and Figure 14(b) and 14(c)). 9.17.15 Principles and Record of Chemical Evolution of Evaporating Seawater Evaporating seawater brine changes its composition together with the precipitation of the sequence of evaporite minerals according to laws reflected by crystallization paths on the graphic diagrams. Some of them are particularly important for the study of the evolution of evaporating brines. 9.17.15.1 Principle of the Chemical Divide for Seawater The precipitation of simple salts induced by and proceeding during evaporite concentration leads to specific changes in the concentration of ions involved in precipitation. This is the consequence of the fact that during the equilibrium precipitation, two conditions must be obeyed simultaneously (Eugster and Hardie, 1978; Eugster and Jones, 1979; Hardie and Eugster, 1970; Hardie and Lowenstein, 2003): 1. The ion activity product of the solution must remain constant at constant pressure and temperature. 2. The ions are removed from the solution in strictly appropriate molar proportions (usually in equal proportions in case of common evaporite minerals: gypsum (CaSO4*2H2O) and halite (NaCl), Ca2þ:SO4 2# ¼ 1:1, and Naþ:Cl# ¼ 1:1, respectively; see Hardie and Eugster, 1970; and Drever, 1982, for more detailed explanation of that process and Hina and Nancollas, 2000, for explanation of the role of concentrations, supersaturation, and stoichiometry in the crystal nucleation and growth, as well as for their relation to the ion activity product). Condition 1 specifies that the concentration of particular ions of the given binary evaporite salts (e.g., Ca2þ and CO3 2# in the case of calcite and Ca2þ and SO4 2# in the case of gypsum) must vary antithetically, thus the rise of concentration of one ion is accompanied with the fall of the concentration of the other (Drever, 1997, Figure 15-2; Hardie and Eugster, 1970). Condition 2 requires that the molar ratio of the 506 Geochemistry of Evaporites and Evolution of Seawater Evaporite concentration The CaCO3 divide Precipitation of CaCO3 HCO3- > Ca2+ Ca2+ > HCO3The gypsum divide Alkaline brine Na-K-Mg Cl-SO4-CO3 Precipitation of gypsum CaSO4·2H2O Ca2+ > SO42- SO42- > Ca2+ MgSO4 brine Na-K-Mg Cl-SO4 CaCl2 brine Na-K-Mg-Ca Cl (a) I. Water with SO4 > Ca Start of gypsum precipitation Mg Concentration SO4 Mg/Ca Ca Salinity II. Water with Ca > SO4 Mg Concentration Ca (b) Start of gypsum precipitation SO4 Mg/Ca Salinity Figure 9 (a) Chemical divides and chemical evolution of major types of waters during evaporative concentration of surface inflow waters (after Hardie and Lowenstein, 2003). (b) Hypothetical passage of the chloride ‘nonalkaline’ brines through the gypsum chemical divide: leading to an increase in SO4/Ca ratio, and hence the Mg/Ca ratio in case of evaporation of seawater type of brine with SO4>Ca (I), and to an increase in Ca/SO4 ratio and consequently to (II) little change in Mg/Ca ratio in case of evaporation of calcium chloride brine with Ca>SO4 (after Hardie, 1987, modified). component ions of precipitated salt must change, unless it is exactly equal to one at the beginning. Molar proportions of all the considered ions in the modern seawater are different than one and were likely different than one in the ancient seawater. Therefore, the ion with the lower concentration at the onset of evaporitic precipitation will progressively decrease in concentration, whereas the other ion showing initially higher concentration will increase in concentration, although relatively more slowly than before the onset of precipitation. The continued precipitation of the salt will lead to a drop in the concentration of the ion showing lower concentration up to the limit of the detection and practically to its selective elimination from the brine, when the precipitation of the given salt ceases. The brine irreversibly changes its composition in the strictly predicted way. Usually, from the modern seawater, at the beginning of evaporative concentration, calcite precipitation enriches the residual solution in the more abundant ionic component and depletes it in the other. Calcite precipitation induced by evaporation of seawater eliminates all HCO3# ions from the solution according to the reaction (Holland et al., 1996): Ca2þðaqÞ þ 2HCO3 #ðaqÞ ! CaCO3ðsÞ þ H2 OðlÞ þ CO2ðgÞ [6] (where aq ¼ aqueous, or soluble in water, s ¼ solid, l ¼ liquid, and g ¼ gas). The greater the initial disparity between the two ionic components, the faster is the enrichment–depletion process. In this way, “calcite precipitation acts as a branching point or chemical divide in the sense that seawaters that are initially carbonaterich will experience a further relative enrichment of carbonate and depletion of calcium and vice versa” (Eugster, 1980, p. 44). Presently seawater is calcium-rich and hence the seawater brine is relatively enriched in calcium and depleted of carbonate after calcite precipitation (Figure 9). However, in the hypothetic soda ocean water of the Archean–Proterozoic, the carbonate ion is more abundant than calcium (Kempe and Degens, 1985), and therefore, evaporation should lead to elimination of calcium within it and to evolution of such waters along the carbonate-rich path of chemical divide and precipitation of Ca-free minerals typical of the soda lakes. The next common evaporite mineral precipitating from modern seawater after calcite – gypsum – “provides a chemical divide with respect to calcium and sulfate in the same manner as calcite provided a chemical divide between calcium and carbonates” (Figure 9(b) I; Eugster, 1980, pp. 44, 46). Thus, each precipitation step acts as “a chemical divide, separating brine evaporation paths depending upon their composition prior to saturation” (Harvie et al., 1982, p. 1615). This was called ‘fractionation by mineral precipitation’ (Eugster, 1980; Eugster and Jones, 1979) or principle of chemical divide as described by Drever (1982) and is now popularized under the name of ‘chemical divide.’ The end of precipitation of a given salt, here CaCO3, is thus interpreted as the exhaustion of the calcium ion, which originally is present in the evaporating water in a relatively small amount. The evolution of the chemical composition of the waters undergoing the evaporation can be interpreted as a succession of chemical divides. Geochemistry of Evaporites and Evolution of Seawater During evaporite concentration and precipitation of successive salts, the modern seawater passes through the sequence of points – ‘chemical divides’ – in a strictly predictable way that depends on mutual molar proportions of particular ions in the seawater. Because in the seawater, as in almost all natural waters, the first mineral that precipitates is calcite, it is the cause of the first chemical divide, and then the gypsum precipitation is the cause of the next divide. Because in modern seawater Ca2þ > CO32# (in molar values), the CaCO3 ceases to precipitate when nearly the entire amount of CO32# ions in the brine is exhausted. Because Ca2þ < SO42#, the CaSO4 • 2H2O ceases to precipitate when Ca2þ ions are exhausted. Similarly, because Naþ < Cl#, NaCl would cease to precipitate when Naþ ions are exhausted (see Levy, 1977). At the turning points, the composition of brine becomes simpler because some ions involved in precipitation are eliminated – they are fixed in the sequence of evaporative precipitates. The evaporating seawater brine evolves in the following way. The seawater (Naþ > Mg2þ > Ca2þ > Kþ/Cl# > SO42# > CO32#) at the end of Ca carbonate precipitation passes into (Naþ > Mg2þ > Ca2þ > Kþ/Cl# > SO42#) within the gypsum precipitation field, and then into (Naþ > Mg2þ > Kþ/Cl# > SO42#) at the end of gypsum precipitation and within halite precipitation field. This latter halite brine contains only five major components (plus conservative Br#), in comparison with seven main components in the initial seawater, and it is this brine from which K–Mg salts are precipitated from modern seawater. Such five-component system of ions was and is the most intensively studied portion of the basic chemical system from which oceanic K–Mg salts are and were precipitated today and in the past. Seawater brines of high salinity show ratios of some ions that are different than ratios in modern seawater. During evaporitic concentration of seawater, ionic proportions remain constant only up to the start of Ca carbonate precipitation. Then, some of these proportions gradually change in the predicted way, starting with the proportions in between Ca2þ and CO32# as well as between them and the remaining ions. During evaporative crystallization of seawater salts, the concentration of some ions rises, attains a maximum, and then gradually drops in a strictly predicted way (Figure 8; Table 5). For example, in the middle of the halite precipitation field (DE !45), the concentration of Mg2þ ions (ppm) equals the concentration of Naþ ions and then becomes higher (McCaffrey et al., 1987). The evolving seawater brine thus not only has a chemical composition simpler from seawater of normal salinity but also is different from seawater, as far as the ratios of ions are concerned; however all these differences can be predicted. The evolution of seawater brine evaporating to dryness within the field of K–Mg salt precipitation is more difficult to predict, mainly because of the strong influence of temperature fluctuations on the kind of precipitated salts and other phenomena typical of high-salinity brines (Braitsch, 1971). The predicted pathways of chemical changes of evaporating brines and associated evaporitic precipitates are commonly traced on the various diagrams along lines that are called the ‘evaporation paths’ or ‘crystallization paths.’ These paths end in the final drying or eutonic point. The principle of chemical divide is successfully used to explain the chemical evolution of waters in many closed 507 basins, although, of course, it is not the only mechanism which influences the hydrochemistry of such basins (Herczeg and Lyons, 1991; Yan et al., 2002). The important limitation of the chemical divide model is the inability to add constituents to the solution (Dargam and Depetris, 1996). 9.17.15.2 Jänecke Diagrams The useful diagram showing the crystallization paths of seawater brines for the five-component system was invented by Ernst Jänecke and is now known as the Jänecke diagram or triangle. The main aim of this diagram is that, by some simplification, it enables one to trace five-component system (Mg–Na–K–Cl– SO4–H2O, stages B–E in Table 6) of evolving evaporating brine on a 2D picture (Figure 10). The preparation of the diagram requires the calculation of the relative proportions of ions in specific Jänecke units (Braitsch, 1971; Krauskopf, 1967; Zimmermann, 2001). The diagram shows relative amounts or ratios of different ions or salts, but not their concentrations in the solution. The diagram permits one to trace the relative changes of Kþ, Mg2þ, and SO42# proportions along the crystallization paths to the drying-up or eutonic point at a given constant temperature (Figure 11). The stability fields of the salt minerals along the crystallization path permit prediction of the expected sequence of crystallization at a given temperature. The diagram does not show the evolving changes in the concentration of Naþ and Cl# as well as changing water content (H2O) in the system. It is assumed that the system is always saturated with NaCl. The important feature of the diagram is that the position of modern seawater remains unchanged during halite crystallization after the precipitation of Ca carbonates and Ca sulfates, until the actual onset of precipitation of K- and MgSO4bearing salts (Horita et al., 2002). For ancient seawaters deficient in sulfates, the other Mg–Ca–2K type of diagram is used (Figure 14). The Jänecke diagram was prepared from laboratory solubility studies of salts at the constant temperature. Some authors adopted the Jänecke diagram to show the real composition of halite seawater brines undergoing natural, that is, polythermal evaporation with accompanying precipitation of K–Mg salts (Garrett, 1980). The ‘crystallization paths’ were represented by dispersed streams of points on such diagrams (Valyashko, 1962) and were used to draw some average lines separating the ‘stability’ fields of the precipitating salts (Valyashko, 1962). Such a diagram for naturally evaporating seawater, much simpler than Jänecke diagram for 25 " C, was drawn by Kurankov and supplemented by Valyashko (1962) and is known as the solar diagram (Figure 12; Holser, 1979a; Valyashko, 1972a,b). Kainite is not present on the solar diagram because this salt was not observed during evaporation of the Black Sea water. Valyashko (1962) enlarged the field of sylvite on the diagram based on the data from experimental evaporation of especially prepared solutions from which the sylvite was crystallized in the course of evaporation as well as the Black Sea water. Farther on, Valyashko (1972b) used the same solar diagram to show the predicted path of crystallization and sequences of K–Mg salts precipitated from hypothetical ancient seawater impoverished with respect to sulfate (Figure 12). Crystallization paths omitted the epsomite field 508 Geochemistry of Evaporites and Evolution of Seawater primary sylvite crystallization from brine similar to present seawater, except for the extremely low content of Mg sulfate. Modern seawater Thenardite 9.17.15.3 Increasing concentration Sylvite Glaserite Bloedite .sw K2Cl2 Na2SO4 z MgCl2 (a) SO4 K2 Thenardite Glaserite Sylvite Bloedite Picromerite sw Leonite Kainite Carnallite Epsomite Sakiite (hexahydrite) z Kieserite Mg Bischofite Crystallization path (b) sw - Modern seawater brine z - Drying-up or eutonic point Figure 10 (a) 3D phase diagram for Mg, Na, K, Cl, and SO4 at increasing concentration at temperature 25 " C. During evaporite concentration, the composition of seawater follows line marked by arrows. (b) 2D diagram known as Jänecke triangle (a projection of the 3D diagram shown in (a)). Crystallization path of seawater is marked by arrows (after Jänecke, 1929 in Dronkert, 1985, Figure 1.1; Łaszkiewicz, 1967; Krauskopf, 1967). and started with sylvite as the first K–Mg salt after halite, exactly as in ancient sequences of marine K–Mg salts, and similarly to those noted during evaporation of the Bonneville salt flat brines, Utah, USA (Hadzeriga, 1964, 1966, 1967). These can be treated as the rare present-day example of massive Spencer Triangle The principle of the chemical divide permits the prediction of the chemical evolution of natural waters undergoing evaporation depending on the initial chemical composition of the water solute (inflow waters). The ‘Spencer triangle’ (Lowenstein et al., 1989; Spencer, 2000; Spencer et al., 1990) is the best graphic technique for the prediction of the evolutionary pathways of the water solutes during evaporation (Figure 13). They are intended to show how inflow waters change, by evaporative concentration and precipitation of calcite and gypsum, into specified type of brine (Jones et al., 2009; Lowenstein and Risacher, 2009; Smoot and Lowenstein, 1991; see also Chapter 7.13). The triangle is a ternary phase diagram for the system Ca–SO4–(CO3 þ HCO3). The basic components, Ca2þ, SO42#, and (CO32# þ HCO3#), are expressed in equivalents, as units of charge concentration, and are placed at corners of the diagram (Figure 13(a) and 13(b)). The calcite (CaCO3) compositional point is placed halfway between the Ca and (CO3 þ HCO3) corners, because there are equal equivalents of Ca and (CO3 þ HCO3) in calcite. Similarly, the gypsum–anhydrite compositional point is placed halfway between the Ca and SO4 corners. Calcite is stable and can crystallize across the entire field of the triangle (similarly as Mg calcite and aragonite). Gypsum–anhydrite crystallizes along the Ca–SO4 side of the diagram. All waters may precipitate halite and other K or Mg salts at some point, particularly on later stages of evolution. There are two chemical divides on the diagram: lines from calcite point to SO4 corner and from calcite to gypsum–anhydrite point, which separate inflow waters into three types, which evolve into relative specified types of brine, depending on the initial water compositions expressed as equivalents of Ca, HCO3 plus CO3, and SO4. These are alkaline (or Na–HCO3–SO4) waters or brines, neutral (Cl–SO4) brines, and calcium chloride (CaCl) brines (Spencer, 2000). Inflow waters can be plotted on the diagram, and their chemical evolution, during evaporation and calcite and gypsum precipitation, evolves into an explicit brine type that can be specifically traced. The initial water represented by specific point on the diagram will precipitate calcite and move directly away from the calcite compositional point. Continued evaporative concentration and calcite crystallization will result in migration of the water composition to the join between (HCO3 þ CO3) and SO4 corners, or the join between Ca and SO4 corners, depending on initial water composition. Ca-poor Na–HCO3–SO4 (or Ca-poor Na–HCO3– SO4–Cl) brines form from waters with equivalents of HCO3 þ CO3 þ SO4 > Ca that evolve directly away from the CaCO3 composition during precipitation of calcite. These waters are not able to precipitate gypsum but precipitate sodium sulfate and sodium carbonate salts. Soda lakes and hypothetic soda ocean water belong to this type (Kempe and Kazmierczak, 2011). Waters in the Cl–SO4 field, such as present-day seawater, form Ca-poor, Na–Cl–SO4-rich brines following the precipitation of calcite and gypsum. Waters in the Ca–Cl field, with Ca equivalents >HCO3 þ CO3 þ SO4 (which implies that some Ca is balanced by Cl), evolve into Ca–Cl brines devoid of SO4 and HCO3 following the Geochemistry of Evaporites and Evolution of Seawater Mg Mg Mg Bischofite Bischofite Bischofite z z z Kieserite Kieserite Carnallite Carnallite Kieserite K2 Kainite Kainite Kainite Epsomite 15 °C (b) Mg Carnallite Bischofite Mg Carnallite Kieserite Sakiite EpsoSylvite mite Kieserite Kainite 0.2 Epsomite SO4 25 °C Bischofite 35 °C Mg Carnallite Bischofite Kainite Kieserite Kainite 0.8 Sakiite Leonite 0.6 Sylvite Sylvite Picromerite 0.4 Leonite sw Bloedite 0.4 sw sw 0.2 Bloedite Picromerite Bloedite 0.6 Glaserite Glaserite K2 Glaserite 0.8 Thenardite 0.8 K2 Sylvite Sylvite Sakiite SO4 Carnallite K2 Sylvite Epsomite 509 Thenardite Thenardite 0.6 Mirabilite 0.4 0.2 (a) SO4 15 °C 0.2 0.2 SO4 25 °C SO4 35 °C Crystallization path z - Drying-up or eutonic point sw - Seawater brine Figure 11 (a) Solid-solution equilibria in quinary system Na2Cl2–K2Cl2–MgCl2–Na2SO4–K2SO4–MgSO4–H2O for 15 " C, 25 " C, and 35 " C, shown on Jänecke diagrams, redrawn from Usdowski and Dietzel (1998), all stability fields saturated with NaCl, (b) enlarged Mg apices of the triangles shown in (a). precipitation of calcite and gypsum. The three types of brines can produce the three predictable distinct assemblages of evaporite minerals (Spencer, 2000; Spencer et al., 1990). 9.17.16 Evaporation of Seawater – Remarks on Theoretical Approaches The numerous modeling approaches to predict the order the evaporative crystallization of salts and the associated quantitative aspects of this process are unsatisfactory (Hardie, 1984). Usdowski and Dietzel (1998) made such a general comment to existing solubility data of marine salts; “a good many of the data are not reliable,” usually “due to the fact that equilibrium is difficult to attain experimentally and that metastable states prevail” (Usdowski and Dietzel, 1998, p. 12). Braitsch (1971) commented that the models based on solubility data are strongly dependent on assumed ideal conditions (such as constant temperature) and added that “it seems highly improbable that actual conditions existing in nature can be adequately represented by one of these models” (Braitsch, 1971, p. 84). Nevertheless, he stated “the comparison of different models with natural salt series is the only direct way of approaching the actual composition of the solutions and the conditions under which they existed” (Braitsch, 1971, p. 84). 9.17.17 Sulfate Deficiency in Ancient K–Mg Evaporites In many ancient marginal marine evaporite basins, the vertical sequence of facies follows the Usiglio sequence, that is, in ascending order, the Ca carbonate ! Ca sulfate ! Na chloride facies. However, K–Mg portions (if preserved) are always significantly different in many aspects. Some salts predicted by theory (bloedite, kieserite, and reichardtite) are rare or absent, 510 Geochemistry of Evaporites and Evolution of Seawater Figure 12 (a) Kurnakov-Valyashko solar diagram (Mg corner) and crystallization paths of seawater and seawater depleted with respect to sulfate to different degrees; the direction of change of the seawater composition is from sw to 1, 2, 3, 4, and 5. (b) corresponding columns of evaporite deposits precipitated from seawater and seawater depleted with sulfate to different degrees after Valyashko, 1972b, modified). other salts (vanthoffite, loewite, and langbeinite) are more common than expected, and some unexpected salts appear during the evaporation of seawater (polyhalite and anhydrite) (Stewart, 1963; Valyashko, 1962). The lack of the bloedite crystallization during the evaporation of seawater at 25 " C can be, at least partly, explained by the extremely difficult and slow experimental nucleation and crystallization of this salt (the appearance of first bloedite crystals in a supersaturated solution required as much as 2 months of waiting; Bergman and Luzhnaya, 1951). From the geochemical point of view, one particular difference is the most significant – the apparent lack of an ‘epsomite facies,’ that is, step B in the sequence of K–Mg salts (Table 6; Braitsch, 1971). The ancient K–Mg sequences commonly start with sylvite, instead of epsomite, and are followed by carnallite-dominated salts and show significantly smaller amounts of sulfates in B-E part of the sequence (Tables 6 and 10). What is more striking is that the deposits that contain primary halite, sylvite, and carnallite are usually entirely free from the primary Mg sulfate salts. Such deposits make up more than 60% of ancient potash deposits (Table 10; Hardie, 1990). Of the many discrepancies between real crystallization sequences and the theoretical predictions, the most remarkable feature is the deficiency of Na–Mg, Mg, and Cl- (456.6) (536.0) SO42(38.5) Ca (20.0) 2+ Mg2+ (111.1) Ca (20.0) + K (9.7) (a) Calcite su lfa Ca te su lcium s lfa tes Neutral Cl-SO4 KMg - HCO3 (2.3) Ca2+ Calcium chloride Ca-Cl Gypsum anhydrite 2+ (17.7) 511 KMg Na+ -C ac Ca hlo l c rid su ium es lfa tes Geochemistry of Evaporites and Evolution of Seawater Neutral HCO3-SO4 2- SO4 (56.2) SO42- HCO3(2.3) Sodium sulfates (b) Sodium carbonates HCO3+ CO32- Figure 13 (a) The major ions in seawater; in miliequivalents per liter (redrawn from Hite RJ (1985) The sulfate problem in marine evaporites. In: Schreiber BC and Harner HR (eds.) 6th International Symposium on Salt, Toronto, Ontario, Canada, May 24–28, 1983, vol. 1, pp. 217–230. Alexandria, VA: The Salt Institute). (b) Spencer triangle – ternary phase diagram illustrating how inflow waters evolve into brines following the principle of chemical divides. Two chemical divides (lines connecting calcite and SO42#, and calcite and gypsum–anhydrite) separate waters that will evolve (following arrows) during evaporation and precipitation of calcite and gypsum into Ca–Cl brines, Ca–SO4 brines, and Ca–HCO3–SO4 brines (redrawn from Smoot JP and Lowenstein TK (1991) Depositional environments of non-marine evaporates. In: Melvin JL (ed.) Evaporites, Petroleum and Mineral Resources. Developments in Sedimentology, vol. 50, pp. 189–347. Amsterdam: Elsevier). The divides are based on the relation of equivalents of calcium to equivalents of SO4 and HCO3 in the inflow water. The relations of these equivalents in seawater are shown on (a). K–Mg sulfates. The sulfate salts, epsomite, kieserite, and kainite, predicted to precipitate from modern seawater, and others, such as langbeinite, are rare or entirely lacking in many ancient marine evaporite deposits (Borchert, 1969). Ancient deposits show generally a much smaller than expected amount of K- and Mg-bearing sulfate minerals than should result from the evaporation of present-day seawater. This is known as the ‘sulfate deficiency.’ This deficiency has been long known and has been a central problem discussed in the geochemistry of evaporites. The ancient sequences lacking the epsomite facies are easy to obtain from the crystallization of modified seawater. This seawater is impoverished with respect to MgSO4 and is CaCl2-enriched, as was proven by both theoretical calculations and experiments (Herrmann, 1991). In particular, Valyashko (1972a,b) presented interpretation of possible paths of crystallization of K–Mg salts depending on the amount of SO4 removed from the solution (Figure 12). The explanation of sulfate deficiency in ancient evaporites splits into three main possibilities, which all can be true depending on the particular case: 1. The chemical composition of ancient ocean was different than today. 2. The deficiency is the effect of the modification of the chemical composition of marine brine in evaporite basins. 3. The deficiency is a secondary effect of the burial and diagenetic alteration of ‘normal’ evaporite sequences (Borchert, 1969; Braitsch, 1971; Dean, 1978; Harville and Fritz, 1986; Petrychenko, 1989). The localized lack of the epsomite facies can be explained by sedimentological processes acting in an open system, such as nondeposition, erosion (synsedimentary dissolution), brine reflux, changes in temperature, or brine stratification (Ayora et al., 1994). Recently, Krupp (2005) presented the concept of large-scale KCl-rich deposition of marine salts similar to the idea of mineral ‘fractionation mechanism’ (Eugster, 1980). This mechanism operates during natural recycling processes on the emerged margins of evaporite basin, where more soluble salts are carried by meteoric waters to the basin center and increase the solute load, particularly the NaCl content, in this zone (Borchert, 1969, with references). In a similar way, Krupp (2005) suggested that the selective leaching of K–Mg sulfate marine salts on the basin margin, combined with various syndiagenetic reactions and transformations during transport of the solutes down to the basin center, could lead to entirely chloride-type potash deposition there. 9.17.17.1 Sulfate Deficiency as the Secondary Feature During the last century, most investigators, with a few exceptions, believed in the near-constancy of the seawater composition at least since the Cambrian, and they mostly ignored the possibility of primary crystallization of MgSO4-poor salts directly from ancient seawater (possibility 1, mentioned earlier). The argument in favor of such a view is the fact that some K–Mg salts or brines from primary halite fluid inclusions of the same or nearly the same age from different subbasins show various degree of sulfate depletion and that the intensity of 512 Geochemistry of Evaporites and Evolution of Seawater sulfate depletion can vary strongly within the same basin (Ayora et al., 2001; Garcı́a-Veigas et al., 1995). A number of hypotheses were suggested to explain the Mg sulfate deficiency by modification of seawater brine in the evaporite environment, that is, assuming that MgSO4-poor salts represent only nonmarine-fed evaporites (Garrett, 1996; Hardie, 1984). Two of the simplest ways are considered here to produce MgSO4-poor brine from seawater. First is to remove SO42#, and the second is to add Ca2þ, which should lead to the additional removal of SO42# trapped in the gypsum (CaSO4 • 2H2O) crystallizing before K–Mg salts. The following major processes were suggested as responsible for such sulfate depletion: 1. The bacterial sulfate reduction and escape of S in the form of H2S to the atmosphere (Sonnenfeld, 1984). 2. The addition of Ca from the nonmarine sources external to the basin such as: (a) calcium bicarbonate-rich river water (Valyashko, 1972b) or (b) CaCl2-rich hydrothermal waters, particularly in rift zones (Hardie, 1990). 3. The additional flux of seawater directly to halite brine (Hite, 1985), which according to Holser (1979a), can lead to precipitation of polyhalite (2CaSO4 • MgSO4 • K2SO4 • 2H2O). 4. The addition of Ca2þ via dolomitization of the previously deposited Ca carbonate (Hite, 1985; Kendall, 1989, 2005; Levy, 1977; Schoenherr et al., 2008) (note that primary direct precipitation of dolomite would only remove Ca2þ and Mg2þ from the brine). The dolomitization of calcite or aragonite (CaCO3) produces dolomite (CaMg(CO3)2) and liberates Ca2þ ions following the ideal reaction: Mg2þ ðaqÞ þ 2CaCO3ðsÞ ! CaMgðCO3 Þ2ðsÞ þ Ca2þ ðaqÞ [7] These Ca2þ ions combine with sulfate ions to precipitate more gypsum (CaSO4 • 2H2O) or anhydrite (CaSO4) and thus lower the content of sulfate in the brine: Ca2þ ðaqÞ þ SO4 2# ðaqÞ ! CaSO4ðsÞ [8] Ca2þ ðaqÞ þ SO4 2# ðaqÞ þ 2H2 OðlÞ ! CaSO4 , 2H2 OðsÞ [9] All together, the process should lead to a decrease in both Mg2þ and SO42# and an increase in Ca2þ ions in the brine (e.g., Hardie, 1987). According to other authors, the dolomitization is also represented by the supplementary reaction (Machel, 2004, with references), which does not liberate Ca2þ ions: Mg2þ ðaqÞ þ CO3 2# ðaqÞ þ CaCO3ðsÞ ! CaMgðCO3 Þ2ðsÞ [10] and the reactions [7] and [10] can be written together as follows: Mg2þ ðaqÞ þ ð2 # xÞCO3 2# ðaqÞ þ xCaCO3ðsÞ ! CaMgðCO3 Þ2ðsÞ þ ð1 # xÞCa2þ ðaqÞ [11] According to Machel (2004), reaction [11] more realistically expresses how much Ca2þ is exported during dolomitization, which depends on particular case characterized by parameter x. 5. The polyhalitization of the previously deposited gypsum or anhydrite (Braitsch, 1971; Hardie, 1984; Harville and Fritz, 1986; Hite, 1985). The polyhalite (2CaSO4 • MgSO4 • K2SO4 • 2H2O) can form from gypsum (CaSO4 • 2H2O) according to the reaction (Hardie, 1984): 2CaSO4 , 2H2 OðsÞ þ 2Kþ ðaqÞ þ Mg2þ ðaqÞ þ 2SO4 2# ðaqÞ ! 2CaSO4 , MgSO4 , K2 SO4 , 2H2 OðsÞ þ 2H2 OðlÞ [12] or from anhydrite (CaSO4) according to the reaction (Braitsch and Kinsman, 1978): 2CaSO4ðsÞ þ 2Kþ ðaqÞ þ Mg2þ ðaqÞ þ 2SO4 2# ðaqÞ þ2H2 OðlÞ ! 2CaSO4 , MgSO4 , K2 SO4 , 2H2 OðsÞ [13] This process should remove not only sulfate but also magnesium and potassium ions from the brine. Hardie (1984) suggested that polyhalite crystallization should precede the deposition of the five-component system and successfully modeled the sulfate-deficient crystallization path under such an assumption (see Harvie et al., 1980). The presented hypotheses were questioned in the following ways: Sulfate reduction was criticized as unrealistic because the rate was too low for this process and a great volume of organic matter would be required as the energy source for the activity of sulfate-reducing bacteria (Hardie, 1985; Hite, 1985). Sulfate reduction commonly takes place in pore waters and does not affect the equilibrium concentration in the brine above the sediments. Similar to dolomitization (discussed in the succeeding text), it is a postdepositional process unable to influence the crystallization path of evaporating basinal brines (Hardie, 1985). Furthermore, Petrychenko (1989, p. 12) noted that H2S is lacking among gases present in inclusions in diagenetic halites suggesting the lack of any bacterial sulfate reduction processes in pore waters (see also Kovalevych et al., 2006a; Petrychenko et al., 2005; Siemann and Ellendorff, 2001). The influx of calcium bicarbonate-dominated river waters to evaporite basins is an expected process. In about 90% of major rivers on Earth, Ca2þ and HCO3# are the most abundant ions; in the remaining rivers, Naþ, CI#, or SO42# are dominant (Meybeck, 1976). It requires improbably large amounts of such waters to supply enough calcium, because meteoric waters are highly diluted (the contents of dissolved solids commonly range from !10 to 1000 mg l#1; Meybeck, 1976, 2003), and therefore, it seems unrealistic (Garrett, 1970; Hite, 1985). The influx of river waters would add not only calcium and carbonate but sometimes also sulfate ions. The influence of Ca from outside of the depositional basin appears to be more realistic in case of smaller basins with limited inflow of seawater or when Ca is supplied by highsalinity hydrothermal sources (Hardie, 1990). In several recent pluvial evaporite basins, supplied with meteoric waters together with the addition of CaCl2 saline waters from deep hydrothermal sources, they were shown to cause the brine modification toward the Ca–Cl field (Lowenstein and Risacher, 2009). Also, the CaCl2 brines can be carried to the surface by convective circulation promoted by thermal Geochemistry of Evaporites and Evolution of Seawater subsurface source or by topographically driven circulation (Hardie, 1990). In this case, however, the effect of the higher salinity of the source waters is apparently opposite from that of marginal marine basins (much more salts are carried in by hydrothermal sources than by meteoric waters). Polyhalitization was criticized by Hite (1985) who believed that polyhalite is found in most evaporites, with a few exceptions, in amounts too small to be responsible for effective modification of the basinal brine composition. Although Stewart (1963) considered polyhalite as the third most abundant sulfate in marine evaporites, after gypsum and anhydrite, and Garrett (1970), together with kainite, as the third most abundant salt mineral containing potash, after sylvite and carnallite, Hite (1985) believed that the polyhalite is mostly a late diagenetic product and that the host brine layer remained relatively unaffected by such polyhalitization. Dolomitization appears to be the best explanation for sulfate deficiency favored by some authors (Hite, 1985; Holland, 1978; Kendall, 2005). Hardie (1998) pointed out on several difficulties in this explanation: 1. Laboratory dolomitization of calcite or aragonite has not been achieved at normal surface temperatures (however, a heliothermal effect can facilitate the dolomitization; Aharon et al., 1977). 2. The dolomitization in marine environments also precedes with difficulty, and usually, it produces calcian dolomites mostly precipitated as cements and not being a replacement product (Hardie, 1987; Pierre et al., 1984a). 3. This modern ‘dolomitization’ has not produced CaCl2 brines (Hardie, 1987; but see Levy, 1977; Wood et al., 2002, 2005). 4. To be effective in sulfate ‘elimination,’ dolomitization should precede the gypsum precipitation (Holland et al., 1996). Hardie (1998) noted, however, that in hypersaline marine environment of the Persian Gulf sabkhas, calcian dolomite forms from hypersaline MgSO4-rich seawater brines only after Ca sulfate has been precipitated (see also Levy, 1977; Pierre et al., 1984a). In the same way, during seepage reflux of such brines, dolomite forms in the bottom sediments of the Solar Lake in Sinai (Aharon et al., 1977). Gypsum precipitation “raises the Mg/Ca ratio well above that of modern seawater, which in turn promotes the formation of a dolomite-like phase” (Hardie, 1998, p. 91). 5. The next problem with dolomitization is how the pore waters modified by dolomitization can be supplied from the subsurface up to the evaporite basin waters. Kendall (1989) suggested the reasonable model of such a process, based on topographically driven hydrological mechanism, where deep formation waters ascending into evaporite basin of the salina type (Figure 4(b)), with a deeply depressed water level, are able to dolomitize the carbonates in the subsurface, and then inflow into the basin from artesian sources. Due to the mixing of these modified waters with basinal waters or brines, gypsum can be precipitated. The lack of any volumetrically significant dolomites in many sulfate-deficient evaporite basins casts serious doubt in 513 the dolomitization hypothesis (e.g., Kovalevych and Vovnyuk, 2010). Except for the topographically driven flow, the other more realistic mechanism for supplying the mineralized water modifying the chemistry of the host water in the basin is thermal convection, driven by subsurface heat sources, commonly transporting Ca–Cl-rich hydrothermal brine up to the surface (as documented by Lowenstein and Risacher, 2009). 9.17.17.2 Sulfate Deficiency as a Record of Ancient Seawater Composition For the past few decades, a growing amount of evidence has clearly suggested that the sulfate deficiency is not merely the result of secondary changes or deposition from nonmarine brine but the primary feature inherited from the host chemistry of ancient seawaters. The justified supposition – that the chemistry of Paleozoic and Mesozoic oceans was quite different than today – has emerged in the late 1970s and 1980s from studies of carbonates. In 1975, Sandberg proved that aragonite ooids with radial structure from the Great Salt Lake are primary forms, thus also proving that the common ancient radial calcite ooids are primary as well, not secondary (from ‘recrystallization’ of aragonite) as it was thought before. Consequently, it appears that the distribution of primary calcite and aragonite ooids in geologic time is apparently regular (Figure 17(f)) and this realization was the basis for distinction of ‘aragonite’ and ‘calcite’ seas, the former precipitating aragonite, the latter calcite as the commonest mineral (Hardie, 1996; Lowenstein et al., 2003; Sandberg, 1983; Wilkinson et al., 1985). Further studies showed that the main factor that presumably promoted the basic difference in Ca carbonate mineralogy in the ancient oceans was the changing molar proportions of Mg2þ/Ca2þ ions in the evolving seawater, oscillating between ! 1 and ! 5 (Figure 17(e); Steuber and Rauch, 2005). Both experimental and observed geochemical data from various sedimentary environments strongly suggest that when the molar Mg2þ/Ca2þ ratio in seawater was high (>2) – the preferred mineralogy was aragonite (and high-Mg calcite), and when low (-2) – calcite (Hardie, 1996, 2003; Lowenstein et al., 2001; Stanley and Hardie, 1998, 1999). Some other factors, such as concentration of SO42# (Bots et al., 2011), however, could also control the aragonite–calcite mineralogy (Holland et al., 1996; Ries, 2010; Zhuravlev and Wood, 2009). A growing amount of evidence suggests that evolving seawater Mg/Ca ratio strongly influenced and also controlled the carbonate mineralogy of skeletal organisms in Phanerozoic supporting the concept of secular fluctuations of Mg/Ca ratio in seawater (Porter, 2010; Ries, 2009, 2010; Ries et al., 2008; Stanley, 2006; Stanley et al., 2002). The mineralogy of ancient marine K–Mg salts shows that the KCl-rich (sulfate-deficient) evaporites and the MgSO4-rich evaporites are also regularly distributed in time, apparently coinciding or overlapping with calcite and aragonite sea time intervals. It seems that during the aragonite seas, as today, MgSO4 salts (such as polyhalite and kieserite) were the main potash minerals, while MgSO4-poor KCl-rich potash salts were dominant in time of calcite seas. The appearance of KCl evaporites coincides with the global high-sea-level periods in the Cretaceous and some earlier parts of the Phanerozoic (Hardie, 1996). Two currently discussed and tested hypotheses explain 514 Geochemistry of Evaporites and Evolution of Seawater this coincidence, suggesting that the major driving forces of the compositional changes of seawater were 1. fluctuations in spreading rate and rate of influx/sequestration of Mg and Ca during the hydrothermal circulation in midocean ridges (Hardie, 1996; Spencer and Hardie, 1990) and 2. increased dolomitization of carbonate platforms during sea-level highstands (Holland, 2005; see, e.g., Steuber and Rauch, 2005, for further comments and information). The essential problem currently studied is how much Mg is able to ‘escape’ from ocean water due to its hydrothermal reaction with the basaltic crust, in comparison with the amount of Mg that is consumed via dolomitization of ocean sediments (Arvidson et al., 2011; Elderfield and Schultz, 1996). The other problem is geochemical evolution of sulfates in the seawater; for example, Canfield and Farquhar (2009) suggested recently that it was bioturbation, which appeared in the Phanerozoic, which caused a severalfold increase in seawater sulfate concentration, contributing to appearance of sulfate marine evaporites. The crucial element in the validation of the new emerging ideas is the reconstruction of the chemistry of ancient oceans from the available sedimentological record, and in this respect, the evaporites became a target of the very intensive studies over the past few decades. 9.17.18 Ancient Ocean Chemistry Interpreted from Evaporites Evaporites themselves supply the crucial, most significant, and direct information on the chemistry of ancient oceans (Berner, 2004; Hardie, 1984). For example, owing to the fact that isotopic fractionation of sulfur during precipitation of gypsum and anhydrite is negligible (Claypool et al., 1980; Hansen and Wallmann, 2003; Holser, 1979b; Seal et al., 2000), except in the later stages of evaporite crystallization within the halite and K–Mg sulfate fields (Strauss, 1997), the isotopic composition of sulfur in marine sulfate evaporites has been used to trace the isotopic evolution of sulfur in the Phanerozoic seawater (Kampschulte and Strauss, 2004, with references). Based on this fact, it can be assumed that marine evaporites record the sulfur isotopic composition of ancient seawater very well (Hansen and Wallmann, 2003). Before the discussion of the use of evaporites in interpreting the chemistry of ancient oceans, it is necessary to pay some attention to the ocean itself. The first assumption was that the ancient oceans (no matter what their composition) had a constant uniform composition worldwide and that they obeyed Marcet’s principle (Forschhammer, 1865), behaving much like today’s ocean, which is a consequence of its continuous mixing. The idea that ancient oceans were not fully mixed but permanently stratified however is commonly applied and accepted in modeling of the Archean, Proterozoic, and Phanerozoic oceans (e.g., Huston and Logan, 2004; Strauss, 1997). Stratification of ancient oceans was more than likely (Reddy and Evans, 2009). Geochemical modeling of the stratified oceans requires a two part, stratified ocean model with chemical (and isotopic) composition different within each part (‘a two-box model’; Holser, 1977; Holser et al., 1989, p. 31). Marginal marine evaporites would be supplied with water from the upper portion only and in such a case the restoration of the chemical composition of the ocean from the evaporites would concern the upper box only. Recently, Garcı́a-Veigas et al. (2011), investigating the geochemistry of Zechstein cyclothems, suggested that the Z2 salts were deposited from upwelling of anoxic bottom seawaters during overturn event of the stratified anoxic Panthalassa ocean (circum-Pangean ocean), and the result of their work coincides with conclusions by Luo et al. (2010). The restoration of the ancient seawater chemistry can be made indirectly from (1) the mineralogy of marine evaporites, (2) the observed vertical sequence of salt minerals in a section, (3) the geochemistry of primary evaporite minerals (trace, minor, and REE elements and isotopic composition of these minerals) and also ‘fossil’ pore fluids, and (4) more directly and precisely from the analysis of chemical composition and other geochemical features of fluid inclusions in primary salt minerals. 9.17.18.1 Implications from Evaporite Mineralogy and from Usiglio Sequence The calculation of the possible limits in the concentration of major seawater ions implied from the mineralogy of evaporites and from the preservation of the Usiglio sequence of crystallization (Ca carbonate ! Ca sulfate ! Na chloride facies) was attempted by Holland (1972, 1984) and Kovalevych (1990) based on chemical characteristics of modern seawater. The lack of sodium carbonates (trona) and bicarbonates, in all known Phanerozoic marine evaporites, implies that during the early stages of evaporation nearly all of the HCO3# has been removed by CaCO3 precipitation (Holland, 1984). In other words, the evaporite precipitation after calcite precipitation did not pass along the chemical divide containing carbonate minerals but followed the sulfate branch (Figures 10(a) and 14). This would imply that in the Phanerozoic seawater, mCaþ2 has always exceeded mHCO3#/2 (Holland, 1984). If the Ca2þ concentration ever fell below half that of HCO3#, Ca2þ would be exhausted during CaCO3 deposition. “If evaporating seawater is to precipitate first calcium carbonate and then calcium sulfate then the calcium ion concentration must exceed one half the bicarbonate ion concentration. If this were not the case, precipitation of calcium carbonate would exhaust the calcium ions in seawater leaving none to enter gypsum” (Walker, 1983, p. 520). The upper limit for Ca2þ concentration in seawater is set by the solubility of gypsum and the fact that seawater is undersaturated with respect to this mineral (Holland, 1972). According to Holland (1972), a threefold increase in the Ca2þ concentration of the modern seawater is enough to produce an ocean, which is saturated with gypsum, in which case the gypsum would be an equally common marine mineral as calcite. Fossil record indicates, however, that gypsum deposition always required some degree of seawater evaporation, so this upper limit for Ca2þ concentration was never reached in any ancient seawater comparable with the modern one. In a similar way, the concentration of Na and Cl in ancient oceans would be much higher than in the present ocean because of the high solubility of NaCl and high concentrations required for its Geochemistry of Evaporites and Evolution of Seawater fite o ch Bis ) hy xa he e( kiit Sa II te dri rite yd yh h ac T Ca III te mi so Ep te ofi ch Bis Mg 515 e llit rna Ca e init Ka I ite on Le ite ed Blo te eri e lvit ite Sy lv Sy om icr P ite ard rite se Gla en Th 2K SO 4 (a) g l-1 Ca2+ traces HCO3- traces HCO3- traces SO42- 21.0 Mg2+ 15.5 (c) traces 15.5 Mg 2+ SO42- + 40 K+ 3.3 60 HCO3- traces 60 Ca2+ traces 120 SO42- traces 120 Mg2+ 15.5 160 K 3.3 160 Na+ 104.1 200 Cl- 191.2 200 Na+ 104.1 g l-1 g l-1 0 0 (d) K+ 3.3 0 (b) 40 Na+ 104.1 40 Ca2+ 10.0 0 60 Cl- 191.2 4 HCO3- 0.15 6 Ca2+ 0.42 120 SO42- 2.77 12 Mg2+ 1.33 160 K+ 0.40 16 Na+ 11.03 200 Cl- 19.83 20 Cl- 191.2 g l-1 (e) Figure 14 (a) K–Mg–SO4 and Mg–Ca–2K type of Jänecke diagrams, and three hypothetical standard seawater brines: modern, sulfate one (I), intermediate (II), and chloride one (III); after Kovalevych (1990). (b) Ionic composition of modern seawater; (c) ionic composition of evaporating modern seawater brine at the start of the halite precipitation, representing standard sulfate brine (I); (d) ionic composition of evaporating hypothetical intermediate seawater brine at the start of the halite precipitation (II); and (e) ionic composition of evaporating hypothetical chloride seawater brine at the start of the halite precipitation (III). Redrawn from Kovalevych VM (1990) Salt Deposition and Chemical Evolution of the Ocean in Phanerozoic. 155 pp. Kyiv: Naukova dumka (in Russian), based on data from Valyashko (1962). 516 Geochemistry of Evaporites and Evolution of Seawater saturation state. The oceans would be far from saturation with NaCl even if all known halite deposits on Earth were dissolved (Holland, 1978, 1984), in which case it would result only in a doubled salinity (!70%, Knauth, 2011), whereas halite saturation of the modern evaporating seawater is at !320%. Salinity of 70% is not enough for the precipitation of halite in the ocean and such copious precipitation was not recorded. The lower limit for Ca2þ concentration in seawater can be established from the fact that the precipitation of gypsum always precedes halite in modern and ancient marine evaporites (Holland, 1972). If Ca2þ concentration in modern seawater were reduced by a factor of 30 (thirty), this water would become saturated simultaneously with gypsum and halite during evaporite concentration. If this factor were larger (>30), then halite would start to crystallize before gypsum during evaporite concentration (Holland, 1972), that is, in a way different than predicted by the Usiglio sequence. On the other hand, “the precipitation of gypsum before halite requires a minimum sulfate concentration of 2.5 mM at a present day Ca2þ concentration of 10 mM (to reach the solubility product of gypsum of 25 mM” (Holland, 1984; Reuschel et al., 2012, p. 85). The other limits can be established from the fact that evaporite gypsum deposition nearly ceases near the start of halite crystallization because the Ca2þ ions necessary for gypsum precipitation are nearly exhausted at just that time. If the Ca2þ concentration in seawater ever exceeds the sum of SO42# concentration and half of HCO3# concentration, the late salts would be enriched in Ca2þ and depleted in SO42# (Holland, 1972). Holland (1972) also claimed that the presence of primary dolomite in marine carbonates implies that the ratio of Mg2þ concentration to Ca2þ concentration has never been less than one in seawater. The methodology outlined by Holland (1972) was used by Walker (1983) and discussed by Grotzinger and Kasting (1993) to establish ranges of seawater composition in the Precambrian. Holland et al. (1986) summarized many previous estimates, including works by Eugster and Jones (1979), Eugster et al. (1980), and Harvie et al. (1980), and claimed that the concentration of nearly all of the major seawater ions could have varied by a factor of 2 to 3 (in either direction) without a modification of the Usiglio sequence of crystallization. Kovalevych (1990) summarized and supplemented these data (Table 7). Table 7 The permissible ranges of concentrations of equivalents of major ions in ocean water during Phanerozoic (in mol kg#1) (after Holland, 1974, interpreted and supplemented by Kovalevych, 1990) Ion Ancient ocean Present ocean Naþ Cl# Mg2þ Kþ Ca2þ (0.230)–0.950 (0.270)–1.100 0.01–(0.4) (0.005)–(0.02) 0.002–0.06 [0.02–0.06] 0.04–0.6 [0.02–0.056] 0.001–0.02 0.002–0.006 0.468 0.546 0.107 0.01 0.02 SO42# HCO3# 0.056 0.002 In parentheses () – very approximate data. In square brackets [ ] – according to interpretation by Kovalevych (1990). 9.17.18.2 Implications of Primary Evaporite Minerals (Excluding Implications from Fluid Inclusions) The original composition of the water, salinity and its fluctuations, can be interpreted to some extent from minor and trace element content both in ‘fossil’ brine (e.g., Vengosh et al., 2000; Boschetti et al., 2011b) and in common primary evaporite minerals (gypsum, anhydrite, and halite). In the latter, more common case, the interpretation is based on known distribution coefficients (Dean and Tung, 1974; Holser, 1979b; Kushnir, 1980, 1982a; Lu et al., 2001, 2002; Ortı́ et al., 1984a; Rosell et al., 1998), particularly for Sr, Na, K, and Mg in case of Ca sulfates (Playà et al., 2007), and Br in case of halite (Holser, 1979b; Valyashko, 1956; Zherebtsova and Volkova, 1966a,b). An attempt at restoration of the chemical composition of Messinian brine, based on trace element concentration in gypsum and anhydrite, was made by Kushnir (1982b) (see comments by Lu et al., 1997) and Mesoproterozoic seawater by Kah et al. (2001). Geochemical data, and in particular REE content and isotopic composition of minerals, can help to distinguish the marine from nonmarine evaporites and are discussed in Section 9.17.19.3 (in the succeeding text). 9.17.19 Recognition of Ancient Marine Evaporites The crucial point in the interpretation of the ancient evaporites and the ancient seawater chemistry (composition) from these evaporite deposits is to properly reveal the marine geochemical signal preserved in the marine marginal evaporite basin, which, as previously discussed, can be always modified by influx of some nonmarine waters. How to define, understand, and recognize ‘marine’ features in ancient evaporite deposits is, however, not an easy task. Braitsch and Kinsman (1978) distinguished the ‘normal marine’ and the ‘modified marine’ evaporites among the primary marine evaporite deposits, with transitional types known only from a few cases. Normal marine evaporites were defined as those “with magnesium sulfates and complex salts such as polyhalite, kainite, langbeinite, but without primary sylvite,” whereas the modified marine evaporites are “without these minerals but with primary sylvite.” As already mentioned, the ongoing studies of the saline giants and the geochemistry of ancient seawater led to the current opinion that the ‘modified marine’ evaporites were likely precipitated from ancient, unaltered seawaters that were different than today’s seawater. One of the problems faced in geochemical studies of evaporites lies in distinguishing the marine (originated from seawater) from continental evaporites derived from the waters similar to seawater in composition, which give the same structure and order of precipitated minerals as seawater. The best example is the Great Salt Lake in Utah, United States, with brine of continental derivation, very similar to seawater brine. Ancient marine and similar continental evaporites can be identical in mineralogy and facies distribution. In attempting to clarify the problem, Hardie (1984, p. 203) defined the autochthonous marine evaporite as “a sedimentary saline mineral deposit formed in situ in a marine or marginal marine depositional environment by evaporation of ocean water Geochemistry of Evaporites and Evolution of Seawater with the composition of average modern seawater, or at least with the composition similar enough to modern seawater to give the same mineral sequence on evaporation.” This definition was chosen to serve for “the purpose of revealing, quantitatively, important differences between modern and ancient ocean water chemistry” (Hardie, 1984, p. 204). Hardie (1984) further suggested the following sedimentological, mineralogical, and chemical criteria, which can help to distinguish the marine (as he defined them) and nonmarine evaporites: 1. 2. 3. 4. The nature of the associated nonevaporite facies Kinds of fossils (if present) Kinds of primary saline minerals The association and vertical succession of such minerals (sequence of crystallization) 5. Geochemical characteristics of such minerals: trace elements, isotope geochemistry, and fluid inclusions 9.17.19.1 Sedimentological Criteria Sedimentological and faunal criteria are not always univocal. Sedimentological criteria indicating the nonmarine evaporites include the geographic setting, the surrounding and intercalating of evaporites exclusively with continental facies, and the presence of nonmarine fossils. However, the reverse, that is, that evaporite deposits intercalated with marine sediments are also marine, is not necessarily valid – this is clear from the nature of every marginal evaporite basin (Hardie, 1984). In marine evaporites, during the complete desiccation and isolation from the sea, continental evaporite facies can develop in the same area immediately under/overlying marine facies (Smoot and Lowenstein, 1991). Naturally, mixed water evaporites are easy to imagine in these settings. Many sedimentary marine and continental evaporite facies are very similar and difficult to differentiate. One of the rare processes and facies not expected in continental lake sediments is tidal fluctuations and related deposits (Smoot and Lowenstein, 1991). However, the lack of evidence of tidal deposits does not exclude the marine character of the basin, which can be surprisingly similar to that in lakes (in fact, many ‘marine’ basins are just lakes supplied with seawater entering the depression through a barrier; Figure 4(b); Ba˛bel, 2004b). In the case of ancient continental evaporite basins, by definition supplied with meteoric waters, the possible influx of hydrothermal Ca–Cl waters can be detected by the following observations: (1) evidences of the volcanism and faulting contemporaneous with sedimentation in evaporite basin; (2) presence of intruded plutonic rocks below the evaporites; (3) presence of metamorphic rocks suggesting the regional hydrothermal heating (as in case of the Salton Sea, USA); (4) the hydrothermal Fe–Mn–Cu–Pb–Zn–Ba mineralization in fractures suggesting the circulation of hydrothermal fluids near the surface; (5) elevated concentrations of Ca–Cl in fluid inclusions of primary evaporite minerals; (6) the relatively high concentrations of trace elements, such as Fe, Mn, Cu, Pb, Zn, and Ba, both within the crystals and in the fluid inclusions of evaporite minerals – calcite, dolomite, gypsum/ anhydrite, and halite, sylvite, and carnallite; (7) presence of minerals typical of Ca–Cl brines – tachyhydrite (CaCl2 • MgCl2 • 12H2O) and antarcticite (CaCl2 • 6H2O), or Ca–Na 517 borate minerals (note however that tachyhydrite is expected to precipitate from Cretaceous seawater brines extremely enriched in Ca; Timofeeff et al., 2006); and (8) absence of minerals that normally crystallize from SO4-rich brines (such as seawater) but not from Ca–Cl brines (Hardie, 1990; Lowenstein and Risacher, 2009; Lowenstein et al., 1989). Based on these criteria, Hardie (1990) pointed out several ancient ‘marine’ evaporite basins as representing deposition from extremely modified seawater, unsuitable for geochemical studies of ancient ocean chemistry (Table 10). 9.17.19.2 Mineralogical Criteria Smoot and Lowenstein (1991) pointed out that distinguishing marine from nonmarine evaporites according to mineralogical criteria is, in practice, nearly impossible. The ideal sequence of crystallization expected for seawater cannot be produced in real environments because of reasons such as (1) the changing supply of seawater and syndepositional reactions brine–minerals, (2) the change in chemistry of seawater by mixing with other waters, and (3) the ancient seawater “may not have always had the same composition” (Smoot and Lowenstein, 1991, p. 304). Mineralogical criteria commonly cannot be helpful because the most common marine evaporite minerals – gypsum and halite – are also the most common in nonmarine evaporites (Hardie, 1984; Smoot and Lowenstein, 1991). Furthermore, some minerals considered as typical of the continental brines also precipitate from seawater brine and in marine environments, for example, mirabilite or glauberite (Hardie, 1985). In particular, such minerals as glauberite (CaSO4 • Na2SO4), polyhalite (K2SO4 • MgSO4 • 2CaSO4 • 2H2O), epsomite (MgSO4 • 7H2O), bloedite (Na2SO4 • MgSO4 • 4H2O), sylvite (KCl), and tachyhydrite (CaCl2 • 2MgCl2 • 12H2O) are not diagnostic for marine/nonmarine settings (Eugster, 1980). There are, however, a few mineral assemblages that cannot crystallize from recent seawater without major modification by specific nonmarine inflow and that today occur in saline lake environments only: (1) Na carbonate minerals, such as trona, nahcolite, and shortite; (2) Na silicate minerals, such as magadiite and kenyaite; and (3) Na or Ca borate minerals (Smoot and Lowenstein, 1991). The first assemblage of minerals is however expected to form from the evaporation of the hypothetical Archean–Proterozoic soda ocean water (Kempe and Degens, 1985). Eugster (1980) listed the following carbonate minerals exclusive to continental evaporites: trona (NaH CO3 • Na2CO3 • 2H2O), gaylussite (CaCO3 • Na2CO3 • 5H2O), burkeite (Na2CO3 • 2Na2SO4), northupite (Na2CO3 • MgCO3 • NaCl), hanksite (9NaSO4 • 2Na2CO3 • KCl), and dawsonite (NaAl (OH)2CO3). 9.17.19.3 Geochemical Criteria Geochemical criteria useful for establishing the distinction between marine and nonmarine evaporites include the trace elements (e.g., Br and Rb contents in chloride minerals – halite, sylvite, and carnallite), isotope (d18O, d34S, or 87 Sr/86Sr and 34S/32S ratios in sulfate salts, 37Cl/35Cl and d11B in chlorides), and fluid inclusion studies (e.g., Boschetti et al., 2011a; Chaudhuri and Clauer, 1992; Chivas, 2007; Denison and Peryt, 2009; Denison et al., 1998; Eastoe and Peryt, 1999; 518 Geochemistry of Evaporites and Evolution of Seawater Eastoe et al., 2007; Eggenkamp et al., 1995; Flecker and Ellam, 2006; Flecker et al., 2002; Holser, 1979b, 1992; Kirkland et al., 1995, 2000; Kloppmann et al., 2001; Lu and Meyers, 2003; Lu et al., 1997; Matano et al., 2005; Palmer et al., 2004; Paris et al., 2010; Pierre, 1988; Pierre et al., 1984a; Playà et al., 2000; Raab and Spiro, 1991; Raup and Hite, 1996; Schreiber and El Tabakh, 2000; Seal et al., 2000; Taberner et al., 2000; Toulkeridis et al., 1998; Toulkeridis et al., 1998; Utrilla et al., 1992; Vengosh et al., 1992). The isotopic composition of ‘evaporative’ or associated carbonates can be helpful to some degree (Magaritz, 1987; Schreiber and El Tabakh, 2000), and 40 Ca/42Ca isotope ratios in gypsum can be used for recognition of the source of inflowing water (Nelson and McCulloch, 1989). Trace elements are useful but not unequivocal indicators of marine evaporites. It is well known that the quantitative range of the trace element Br, found both in today’s seawater brines and the marine halite, may overlap ranges present in nonmarine brines and halites (Hardie, 1984, 1985; Warren, 2006). Because of that, and in particular, the values identical or similar to recent marine halites found in ancient halites cannot be used as indicators of marine origin of such halite. The values that are much lower or higher than those present in today’s halites can only suggest but not prove nonmarine or a recycled origin of salts (developed from crystallization from nonmarine brine or from brine developed from dissolution of marine salts) (for comparison, see Schoenherr et al., 2008). Not only bromine but also the other trace elements in solid solutions within evaporite salts can be misleading in distinguishing of marine from the nonmarine evaporites (Hardie, 1984). A similar overlap in range exists, for example, for marine and nonmarine 34S/32S ratios (Smoot and Lowenstein, 1991). REE patterns of concentration in braitschite (Ca, Na2)O • REE2 • O3 • 12B2O3 • 6H2O) and gypsum were used in order to distinguish the marine from continental source of brines in evaporite deposits by Raup (1968), Toulkeridis et al. (1998) and Playà et al. (2007). High-quality trace element studies, however, require special laboratory techniques to remove contamination coming from fluid inclusions dispersed within the crystals (Lu et al., 1997; Moretto, 1988; Playà and Rosell, 2005; Raup and Hite, 1996; Schröder et al., 2003). From the listed studies, only the fluid inclusion analysis gives direct information about the brine chemistry and temperature at the time of crystal growth (Holser, 1979a; Knauth and Beeunas, 1986; Knauth and Roberts, 1991; Roedder, 1984; Siemann and Ellendorff, 2001; Smoot and Lowenstein, 1991; Timofeeff et al., 2001). The specific criteria for recognition of the marine character of brines in ancient halite fluid inclusions are described further in the text. 9.17.20 Fluid Inclusions Reveal the Composition of Ancient Brines Analyses of primary fluid inclusions supply the best, reliable, and direct information about the chemical composition of brine from which the salt crystallized (Hardie, 1984; Holser, 1979a). The advance of the modern analytical techniques has enabled considerable progress in gaining new information on the chemistry of the evaporating ancient brines, and halite appears to be the best mineral for such a study. Ancient halite is much better preserved than the other common marine evaporite mineral – gypsum, which becomes altered to anhydrite by dehydration during burial (Jowett et al., 1993). Inclusion studies, in gypsum, are rare and were made only in primary Tertiary selenites. Thus far, analyses from the gypsum of marginal marine evaporite basins showed very low values of salinity, all are actually outside of the value range for gypsum crystallizing from seawater brine, and much lower than the standard 35% seawater salinity (e.g., Attia et al., 1995; Kulchitskaya, 1982; Peryt, 2001; Petrychenko et al., 1997). In contrast to ancient crystals, fluid inclusion analyses in modern marine evaporitic gypsum yielded proper salinity values, typical of salinity range for the precipitation of that mineral from seawater (Sabouraud-Rosset, 1972). Holser (1963, 1979a) was one of the first who began the study of brines from fluid inclusions in ancient halite and showed that Mg/Cl and Br/Cl ratios of inclusion brines from the Permian halite are similar to those of modern seawater. Later, very detailed studies revealed many features of basinal halite brines such as pH, Eh, temperature, and composition of gases trapped in these brines (Kovalevych, 1990; Petrychenko, 1988). What was more significant is that they proved, without any doubt, that the chemical composition of many basinal waters of marginal marine evaporites preserved in primary halite inclusions was indeed different from recent seawater brine (at the stage of halite saturation). Further advance of modern analytical techniques has enabled collection of even more detailed and precise data from many basins, which confirmed these results. From these studies, a clear picture emerged in late 1980s that the brines in ancient marginal marine halite basins (well before the start of K–Mg salt precipitation) were more commonly Na–K–Mg–Ca–Cl type, not Na–K–Mg–Cl–SO4 type as expected for the evaporation of modern seawater to halite precipitation field (see Das et al., 1990, and summaries in Horita et al., 1996, 2002; Kovalevych et al., 1998a,b; Horita and Holland, 1998; Holland, 2003). Additional studies of halite inclusion brines have supplied more and more data confirming this picture (Khmelevska et al., 2000; Kovalevych et al., 2006a,b; Petrychenko et al., 2005, 2012; Timofeeff et al., 2006). It also appears that many halite basins lacking K–Mg salts follow the same general trends and most commonly evidence a Na–K–Mg–Ca–Cl type of brine in primary halite fluid inclusions, not Mg sulfate brine similar to modern seawater brine, as it was expected earlier (e.g., Hite, 1985). This was a new and important fact. The simplest explanation of that observation is that the composition of the ocean evolved with time, and such an interpretation was early accepted by Petrychenko (1988) and Kovalevych (1990), among others. All these studies clearly showed that the ancient Mg sulfate-poor evaporites were not the product of secondary postdepositional replacement processes, and the halite brines in many ancient evaporite basins were already Mg sulfate-deficient at the beginning of halite precipitation, which excludes the possibility of postdepositional replacements. The problem then appeared as how to illuminate the chemical composition of the ancient seawater more precisely from the studies of halite fluid inclusions. The next target and challenge in the study of fluid inclusions in ‘marine’ halite Geochemistry of Evaporites and Evolution of Seawater 519 is how to recognize the composition of the ancient seawater from the chemical composition of brine trapped in the fluid inclusions. One of the first attempts, before the use of numerical modeling described in the succeeding text, was made by Kovalevych (1990) who introduced the concept of ‘standard’ seawater and ‘standard’ seawater brines. Following the lead of some earlier investigators (Valyashko, 1962, and others), he suggested that ancient Phanerozoic seawater evolved in between three basic and ultimate types called sulfate, intermediate, and chloro-calcium waters and established the composition of major ions in these seawaters (except for the sulfate type, which contained the same array of ions as in modern seawater). He assumed, partly because of the very high residence time of some ions and also because there was no evidence concerning the scale of their variations through time, that the concentrations of major ions Cl#, Naþ, Kþ, and Mg2þ are the same in all three standard seawaters. We now know that the concentration of Mg2þ surely can vary through time because this ion is readily removed from seawater during hydrothermal circulation of seawater through mid-ocean ridges. However, at the same time, confusing the issue, the concentrations of the remaining SO42#, Ca2þ, and HCO3# varied systematically. Kovalevych then calculated the composition of the seawater brines of the three particular seawater types at the beginning of each halite precipitation stage and drew the points of the three hypothetical standard seawater brines on the Jänecke type of diagrams (Figure 14; see also Kovalevych et al., 1998b). 9.17.20.1 Criteria for Seawater Recognition in Halite Fluid Inclusions Fluid inclusions in primary halite were used for the identification of the marine evaporites based on the following test criteria (Horita et al., 2002; Timofeeff et al., 2001; Zimmermann, 2000): 1. Sampling of the earliest (first) halite in a precipitation sequence, that appears just above the last Ca sulfates (gypsum and anhydrite), well before the start of potash salt precipitation, and from the earliest salt cyclothem in a sequence (Kovalevych, 1990; Kovalevych et al., 1998b). These earliest halites are assumed to have been precipitated from the least modified seawater brine in a marine marginal evaporite basin. Such a halite also eliminates the complications of a back reaction of early formed salts with evolved brines and its influence on brine chemistry, which may bias the marine signal (Timofeeff et al., 2001). The early halite can be additionally recognized by relatively low Br content or other geochemical data (Horita et al., 1996). Sampling within the tectonically stable basins that are devoid of carbonate platforms, without an influx of clastic deposits, and is close to the seawater inlet into an evaporite basin is required (Horita et al., 2002). This criterion is crucial in a true recognition of marine signal in ancient evaporite deposits, and it was proved by Garcı́a-Veigas et al. (2011) and Cendón et al. (2008) that in the late Permian Zechstein basin and Oligocene Mulhouse basin in France, only the earliest halites bear relatively unchanged marine brines, whereas the upper parts of each section Figure 15 Primary growth zoning of the halite cube faces, created by bands of fluid inclusions, Bochnia Salt Mine, upper Zuber deposits (corridor Tesch, level V), the Badenian of the Carpathian Foredeep, Poland, length of the crystal ¼ 65 mm, photos courtesy Krzysztof Bukowski. show remarkable chemical evolution related to many processes including recycling of salts (see also Petrychenko et al., 2012, for similar interpretation). 2. Sampling the halites from deposits with features diagnostic of perennial subaqueous depositional conditions to minimize the possibilities of syndepositional recycling (Smoot and Lowenstein, 1991; Timofeeff et al., 2001). Most commonly fluid inclusions from the first developed ‘chevron’ structures, representing primary growth bands (growth zoning) of crystals, are analyzed (Figure 15). The recycling process, that is, dissolution of earlier precipitated salts, increases the proportion of the ions derived from the dissolution of these salt (e.g., Naþ and Cl# from halite, Kþ, Mg2þ, and Cl# from carnallite) in the seawater, which then will have a modified composition. Recycling is so common (e.g., Logan, 1987) that it is impossible to exclude it in any marginal marine evaporite basin. The recycling (dissolution) of halite (NaCl) and sylvite (KCl) causes the ions Na, Cl, K, and Cl to be added to basinal brines in equal molar proportions (1:1:1:1) changing the Na/Cl and K/Cl ratios in these brines. For example, brines in the Qaidam Basin has an elevated concentration of K, up to several hundred milimolal above the value predicted by the evaporation of parent waters, because the dissolution of 520 Geochemistry of Evaporites and Evolution of Seawater carnallite added K to these basinal brines (Spencer et al., 1990). High Kþ concentration in halite fluid inclusions can reflect recycling and syndepositional dissolution of potassium salts. Such inclusions should not be used for the reconstruction of ancient seawater chemistry (Lowenstein et al., 2005). “Evaporative concentration of such altered seawater will produce brines with a different chemical composition from brines evolved from pristine seawater” and halite crystallized from such significantly altered brines “cannot be used for study of ancient seawater chemistry” (Timofeeff et al., 2006, p. 1982). However, Timofeeff et al. (2006, p. 1982) noted that “the great mass of dissolved salt in large brine bodies makes modification of the major-ion chemistry by syndepositional recycling processes or nonmarine inflow waters less likely than in shallow/ephemeral systems.” Nevertheless, commonly analyzed macroscopically visible chevron structures, with primary fluid inclusion bands in halite, are created only in shallow brine (Holser, 1979a) but are absent in primary coarse crystalline and clear halite cubes growing at depth down to 250 m in the Dead Sea (Herut et al., 1998). Timofeeff et al. (2006) used the additional nonpetrographic criterium to exclude the possibility of recycling, which can be used after the fluid inclusion analyses (no. 3 in the succeeding text). 3. “The major ion chemistry of an individual fluid inclusion must lie on the evaporation curve defined by the entire suite of fluid inclusions from the same deposit. Individual fluid inclusions whose chemical compositions do not fall on the defined evaporation curve must have formed from chemically-modified parent seawaters” (Timofeeff et al., 2006, p. 1982). Such inclusions should be excluded from the calculations of seawater chemistry. The large degree of scatter in data from halite fluid inclusions suggests that they are not primary in origin (Horita et al., 2002, p. 3737). Zimmermann (2000, 2001) described the other calculations and graphic methods for distinguishing the evaporated seawater from other types of brines met in halite inclusions. 4. The comparison of major ion chemistry of halite fluid inclusions on plots (Naþ, Kþ, and SO42# against Mg2þ and Cl#) from several geographically separated evaporite basins of the same age, utilizing a significant number of analyses, will permit the tracking of the evaporation paths in separate basins. If these plots (a) fall along the same distinctive evaporation path and if (b) the paths for the various basins of the same age overlap one another, this will imply that the parent water had a uniform chemical composition and thus represents the true ancient seawater. The overlapping evaporation paths clearly indicate minimal influence of nonmarine inflow and syndepositional recycling (Lowenstein et al., 2001; Timofeeff et al., 2001). The criterion of consistency is crucial for testing the reliability of the results and it is similar to criterion introduced earlier by Nielsen (1989) for testing the sulfur isotopic age curve, based on analyses of Ca sulfates from different evaporite basins (Strauss, 1997). 5. The analysis of fluid inclusions requires the use of clearly defined criteria for primary inclusions (Vovnyuk and Kovalevych, 2007). “Primary, single-phase brine inclusions with negative crystal shapes in primary halite in which hoppers and chevrons are outlined by alternating bands of inclusion-rich and -free zones are preferable” (Horita et al., 2002, p. 3734). Additionally, the statistically significant number of samples should be analyzed to confirm the result. 6. To be sure that the sample and results really, or in the best way, reflect the ancient seawater composition, the compositional changes of the basinal brine should be ruled out by complex systematic sedimentological and geochemical studies within the whole section and basin area. This very rigorous criterion was suggested by a group of authors who proved that in many basins, including those that were sampled for ancient halite brine analysis and gave ‘positive’ results, such changes are particularly common phenomena (Ayora et al., 1994, 1995; Cendón et al., 2008; Garcı́a-Veigas et al., 1995). They pointed out that “the detection of brinerock reactions is not possible by analyzing isolated samples. Reaction detection is only possible when the brine evolution is reconstructed in detail by using systematic fluid-inclusion analyses throughout complete sequences and numerical simulation of evaporation scenarios. Moreover, this methodology distinguishes which parts of the evaporite sequences, apparently deposited in a marine setting, are in fact formed by recycling previous evaporites in an endorheic basin” (Ayora et al., 2001, p. 251). Zimmermann (2000) applied other factors that discredited the validity of some analyses and did an exemplary ‘screening’ of the various data obtained for Tertiary salt deposits. Horita et al. (2002, p. 3732) frankly commented that if one follows the criteria listed earlier, “it is difficult, if not impossible, to identify evaporite deposits that meet all these requirements.” Their screening method included comparison of the halite of the same or similar geologic age from different evaporite basins to determine whether the composition of inclusion brines bears global (i.e., seawater) or rather local or regional signal. 9.17.20.2 Reconstruction of Ancient Seawater Composition from Halite Fluid Inclusions The reconstruction of the ancient seawater compositions from the composition of brines in fluid inclusions trapped in primary evaporite mineral (halite) requires three distinct study steps (Timofeeff et al., 2001): 1. An accurate technique for the chemical analysis of fluid inclusions must be utilized. The methods used for precise halite inclusions investigation were described by Timofeeff et al. (2000) and reviewed by Vovnyuk and Kovalevych (2007) and Kovalevych and Vovnyuk (2010). The concentrations of Na and Cl in some halite fluid inclusions must be calculated (adjusted) with the help of the computer program by Harvie et al. (1984) (see Timofeeff et al., 2001). 2. The establishment, by the use of the rigorous criteria (described in the previous sections), that the fluid inclusions really contain evaporated seawater, uncontaminated by nonmarine inflow and/or syndepositional recycling. 3. The utilization of the method for the backcalculation of the chemical composition of seawater from the composition of fluid inclusion brines that have undergone evaporative concentration and have been modified by the precipitation of evaporite minerals (such as calcite and gypsum in case of ancient seawaters similar to modern seawater). Geochemistry of Evaporites and Evolution of Seawater This third point is the most difficult test to pass, because the reconstruction of original composition of seawater from halite seawater brine requires the solution of two basic difficulties: (3a) “Assumptions must be made in defining the degree of evaporation (DE)—that is, the ratio of the concentration of a conservative element in brines to that in the initial seawater” (see eqn [3]; von Borstel et al., 2000); (3b) “Uncertainties are introduced by the precipitation of mineral phases (carbonates, gypsum/anhydrite, and halite) before and during halite precipitation” (Horita et al., 2002, p. 3734). The modern analytical techniques permit detection of the concentration of Br in halite inclusion brine. This concentration was used to calculate the DE of the fluid inclusion, under reasonable assumption that Br concentration in seawater has not changed during Phanerozoic, because Br has residence time in the ocean ! 100 My (Horita et al., 2002, but see Leri et al., 2010). Furthermore, the data from fluid inclusions suggest that similarly, the concentration of K did not change significantly during Phanerozoic (Horita et al., 2002) and therefore K concentration can also be used as a measure of DE in some cases. The backcalculation is made by a simple trial-and-error fitting method with the application of the numerical computer modeling program for evaporating seawater type of brines devised by Harvie et al. (1984). The diagrams produced from geochemical data collected from halite fluid inclusions are visually compared with those calculated by computer program under some trial assumptions and the best-fit diagrams are found that are chosen to represent the original composition of ancient seawater. This methodology was successfully tested on the recent marine halite deposits (Timofeeff et al., 2001). One of the first breakthrough papers on the use of fluid inclusion studies to recognize and follow oscillations in seawater chemistry was published by Lowenstein et al. (2001). Essential in this and following interpretations is the fact that the marine halite precipitation field is broad (Figure 8). In recent evaporating seawater, halite begins precipitation from c.290–320% and continues crystallization until the start of epsomite precipitation at ! 375%. At that time, in the ideal closed system, no other salts precipitate from the brine except for some gypsum at the beginning of halite crystallization field, so the majority of ions show conservative behavior within this field. When the salinity rises, the concentration of Mg and K ions in the halite brine also rises proportionally, and the molal ratios of these conservative ions are preserved, some of them reflecting or still preserving the ratios present in original ancient seawater. The Na and Cl ions will however change their proportions and concentrations because of the chemical divide principle (Cl> Na in seawater). The brine (of different salinity) trapped in fluid inclusions in halites that crystallized along the crystallization paths within the halite field will thus record these constant ratios or the gradual changes of ion concentrations. Thus, it is possible to relate these ratios and changes to changes in salinity and restore the original proportions of salts in the original ancient seawater. The start of halite crystallization on the evaporation path of recent seawater is recorded by a drop in Naþ and an accelerated increase of Mg2þ concentration at about 6000 mmol of Cl#. One of the main difficulties was however the determination of concentration of Na and Cl in small halite inclusions (investigated by ESEM X-ray 521 EDS and extraction-IC techniques), which required special adjustment with the help of computer modeling calculations (Timofeeff et al., 2001). Lowenstein et al. (2001) analyzed fluid inclusions from late Precambrian (543–544 Ma, Ara Group, Oman; Schröder et al., 2003; Schoenherr et al., 2008); Early Cambrian (520–540 Ma, Angarskaya Fm., Siberia); Silurian (418–440 Ma, Salina Group, Michigan, United States, and Carribuddy Group, Western Australia); Permian (251–258 Ma, Salado Fm., New Mexico, United States); Early Cretaceous (112–124 Ma, Muribeca Fm., Brazil, and Loeme Fm., Congo); Late Cretaceous (94–112 Ma, Maha Sarakham Fm., Laos–Thailand); and some Tertiary evaporite basins, as well as modern marine halites. All these basins were interpreted as marine in origin, based on geologic criterion. The concentrations of Mg and Na against Cl from fluid inclusions were traced on the diagrams, which revealed distinct evaporation paths of the basinal waters (Figure 16). Late Precambrian, Permian, and Tertiary paths nearly coincided, and the same feature showed the Cambrian, Silurian, and Cretaceous paths grouped together differently. These two pathway clusters reflect higher amounts of Mg in the evaporating water of the former group than in the latter. What was the most significant – the evaporation paths from geographically separate basins of the same age overlapped, which was taken as the crucial evidence for the shared marine derivation of the analyzed basinal brines. All of the ancient brines showed relatively lower concentrations of Mg than the modern seawater brine, with the late Precambrian and Permian fluid inclusions closest to modern concentrations (Lowenstein et al., 2001). Lowenstein et al. (2001) also established the Mg/Ca ratios (maximum and minimum values for each time interval) of palaeoseawaters by using both the measured concentrations of brines from the fluid inclusions (all the major ions) and the complicated but logical series of assumptions and restorations. For the Cambrian, the Silurian, and the Cretaceous, the maximum values of Mg/Ca were taken directly from the measurements of Mg and Ca concentrations in fluid inclusions. In these measurements, the concentration of Ca was interpreted as lower in relation to Mg (that is the conservative ion) due to loss of the Ca ion in the earlier precipitation of CaCO3 and CaSO4. The minimum value of Ca concentration was calculated using the appropriate backtracing procedure. The measured values of concentrations of the other major ions (Na, K, Ca, and Cl) were plotted versus the concentration of the Mg ion, which was interpreted as being the measure of the evaporite concentration until the first Mg salt precipitation. Then by using a computer program developed to monitor the changes in concentration of major ions during evaporite concentration of given water (Harvie et al., 1984) and by using the trial-anderror fitting method, the chemical composition of paleoseawater was determined (the best fit to the data). The succession of salts formed by the evaporation of modeled paleoseawater matched the observed sequence in ancient evaporites (Usiglio sequence). “All modeling was done using a present-day Cl# of 548 mmol and assuming that SO42# was 14 mmol; presentday SO42# is 28 mmol” (Lowenstein et al., 2001). The fluid inclusions in the halite of Precambrian and Permian evaporites did not contain measurable Ca. Apparently, the ancient brines were similar to modern evaporating seawater brine, the evaporation and precipitation of CaCO3 and CaSO4 522 Geochemistry of Evaporites and Evolution of Seawater 6000 Na (mmol kg-1-H2O) 5000 Primary fluid inclusions in halite Modern 4000 Late Cretaceous Early Cretaceous Permian 3000 Silurian Cambrian 2000 Precambrian Seawater evaporation 1000 0 3000 4000 5000 6000 7000 Cl 8000 9000 10 000 11 000 (mmol kg-1-H2O) 5000 Mg (mmol kg-1-H2O) 4000 Evaporating modern seawater brine up to K-Mg-salts precipitation (data after McCaffrey et al. 1987) Evaporating modern seawater brine during K-Mg-salts precipitation (data after McCaffrey et al. 1987) 3000 Primary fluid inclusions in halite 2000 Modern Late Cretaceous Early Cretaceous Permian Silurian Cambrian Precambrian Seawater evaporation 1000 0 3000 4000 5000 6000 7000 8000 9000 10 000 11 000 Cl (mmol kg-1-H2O) Figure 16 Measured concentrations of Naþ (A), and Mg2þ (B), versus Cl#, in primary fluid inclusions present in modern and ancient halite (after Lowenstein et al., 2001; Timofeeff et al., 2001; von Borstel et al., 2000) and in evaporating Caribbean seawater during fractional crystallization in saltwork pans and laboratory (based on data by McCaffrey et al., 1987), and compositional path of evaporation of modern seawater calculated by computer program written by Harvie et al. (1984). Naþ and Cl# molalities from all fluid inclusions were adjusted with the help of the same program under assumption of NaCl saturation. Diagrams reproduced from Lowenstein et al. (2001), modified and supplemented. consumed virtually all Ca2þ from the brines, leaving SO42# in excess during further evaporite concentration, that is, during halite precipitation. The Mg/Ca ratios for these waters were calculated similarly as described earlier. Ca2þ concentration equals 10 mmol (which is Ca2þ concentration in modern seawater giving the maximum Mg2þ/Ca2þ ratio on the diagram) was used for maximum value of the Mg/Ca ratio, and a Ca2þ of 15 mmol, that is, 1.5 times modern seawater – for minimum Mg2þ/Ca2þ ratio. A similar procedure of calculations was used for Tertiary seawaters based on Mg2þ concentrations determined from fluid inclusions for these waters by other authors. The results were presented on the diagram supplemented with the curve showing predicted modeled variations of Mg/Ca ratio in time previously calculated by Hardie (1996). The final point of their discussion is the restoration of the composition of ancient seawater from inclusions accepted as representing seawater brine, that is, from established coincident and overlapping evaporation path taken from fluid Geochemistry of Evaporites and Evolution of Seawater inclusions from separate basins of the same age. It is done by using the HMV computer program simulating evaporation paths of some assumed seawater. “The calculated evaporation pathways are plotted and compared with the measured inclusion brines. The parent seawater chemistry is then adjusted in an iterative process to best fit the fluid inclusion data” (Timofeeff et al., 2001, p. 2299). In the ensuing papers, increasingly detailed restorations of seawater chemistry were presented basing on the same methodology and number of the same or similar interrelated assumptions (Brennan and Lowenstein, 2002, 2004; Lowenstein et al., 2005). The calculations of the chemical composition of Permian seawater from halite fluid inclusions were made under the assumptions proposed by Lowenstein et al. (2005): 1. Salinity and concentration of Cl#, that is, chlorinity, of Permian seawater were the same as in the modern seawater. 2. HCO3# can be ignored in the modeling because in the recent seawater, the concentration of this ion (2.5 mmol kg#1 H2O) is negligible in comparison with Cl# (565 mmol kg#1 H2O) and with other major ions (Table 1). 3. The milimolal concentration of Kþ in Permian seawater was assumed to be equal 10, similarly as supposedly in the entire Phanerozoic (Horita et al., 2002). This assumption results from the fact that the recorded K/Br ratio in Phanerozoic halite fluid inclusions has been relatively constant (Horita et al., 2002) and that Br concentration in Phanerozoic seawater presumably did not change significantly, because the residence time of this element in seawater is estimated at 100 Ma. The Mg2þ concentrations were calculated from Mg2þ/Kþ ratio recorded in fluid inclusions. The ratios of Kþ/SO42# in brine inclusions were used to calculate the SO42# concentrations in Permian seawater. The concentration of Naþ was calculated from charge balance, after the concentrations of all the other ions (Cl#, SO42#, Ca2þ, Mg2þ, and Kþ) were estimated. 9.17.21 Ancient Ocean Chemistry from Halite Fluid Inclusions – Summary and Comments At present, nearly all major time intervals containing Phanerozoic saline giants are covered by the analyses and reconstructions of seawater chemistry from halite fluid inclusions (Table 8; Figures 17 and 18) and are used for testing the geochemical models of seawater evolution (e.g., Berner, 2004; Hansen and Wallmann, 2003; Holland, 2005). The main results of these restorations are the following: Permian seawater was chemically similar to modern seawater; however, it was slightly depleted in SO42# and enriched in Ca2þ in relation to present-day seawater, although the Ca2þ concentration close to modern 11 mmol kg#1 H2O cannot be excluded (Garcı́a-Veigas et al., 2011; Lowenstein et al., 2005). The Mg2þ/Ca2þ ratio was > 2 that was favorable for the precipitation of aragonite and Mg calcite as ooids and cements. Fluctuations in Mg/Ca ratio and fluctuations of Ca and Mg concentrations in time, as detected from halite fluid inclusions, are well supported by numerous geochemical studies of carbonates (Steuber and Rauch, 2005). 523 During the Early Cretaceous, the concentration of Naþ at the given concentration of Cl# in primary fluid inclusions in halite was higher than today, than in Permian, and in latest Neoproterozoic (Lowenstein et al., 2001). The Cretaceous seawater showed very high calcium concentrations, which gives the lowest Mg2þ/Ca2þ ratios (! 1) documented in Phanerozoic seawater from halite fluid inclusions (Lowenstein et al., 2003, 2005; Timofeeff et al., 2006). Such low ratio favored the precipitation of calcite from seawater rather than aragonite or high-Mg calcite. Elevated Ca2þ concentrations leading to the relation Ca2þ > SO42# at the point of gypsum saturation permitted the Cretaceous seawater to evolve into Mg–Ca–Na–K– Cl brines devoid of measurable SO42# and able to precipitate rare calcium-containing evaporite mineral tachyhydrite (CaCl2 • 2MgCl2 • 12H2O) at the end of evaporation path (Figures 8, 14, and 18; Wardlaw, 1972a). Since the Early Cretaceous, seawater evolution has been unidirectional – the concentration of Mg2þ and SO42# increased and concentration of Ca2þ decreased (Holland, 2005). Kump (2008) noted an intriguing rule that the concentration of Mg2þ and Ca2þ was inversely related when the concentration of Ca2þ was less than !20 milimol and positively related when it was above this concentration. It seems that variations in concentrations of Ca2þ and SO42# in Phanerozoic seawaters were synchronous and inversely proportional (Demicco et al., 2005; Hansen and Wallmann, 2003; Kovalevych, 1990; Kovalevych and Vovnyuk, 2010; Stanley and Hardie, 1998; Figure 17(b) and 17(c)). Potassium showed a consistent low stable concentration in all studied Phanerozoic time intervals, and its (assumed) constant level permitted the calculation for concentrations of other ions. Cl variations over time, hence also paleosalinity of seawater, were not determined from the halite fluid inclusion studies (Knauth, 2005, p. 60) and, similar to K concentration, was assumed to be constant. The interpreted concentrations (particularly Naþ) were ‘adjusted’ to paleosalinity (chlorinity or concentration of Cl) of the same level as today. The assumed constant concentration of Br in ancient seawaters served as a base of determination for the DE, as well as the base of the calculations of the concentration of other ions in ancient seawaters. The concept of a bromine constant is now doubtful in the light of investigation by Channer et al. (1997) and Gutzmer et al. (2003) and because Leri et al. (2010) showed that Br is likely not the conservative element in ancient seawater. Furthermore, some recent studies showed nonconservative behavior of bromine in both the ocean and salt pan environments (Berndt and Seyfried, 1997; Martin, 1999; Risacher et al., 2006; Wood and Sanford, 2007). The partition coefficient of Br in halite depends on the chemical composition of seawater and can be both experimentally and theoretically predicted for seawater of various compositions (Siemann and Schramm, 2000). The Br content in first marine halites from various stratigraphic intervals appears to vary exactly following the interpreted changes in Phanerozoic seawater chemistry as predicted by Hardie (1996), this being the first independent test for his model of seawater evolution (Siemann, 2003). The causes of these compositional changes and the chemical evolution of the ocean are the subject of the current ongoing scientific debate and numerical modeling. The various 524 Major-ion chemistry of ancient seawater (mmol kg#1 H2O) interpreted from chemical composition of fluid inclusions in marine halite; selected results of recent investigation Time Age (Ma) m(Cl#) m(SO42#) m(HCO3#) m(Naþ) m(Mg2þ) m(Ca2þ) m(Kþ) m(Mg2þ)/m(Ca2þ) References Modern seawater 0 565 29 2.5 485 55 11 11 5.2 Cretaceous (Albian– Cenomanian) Cretaceous (Aptian) 112.2–93.5 565 14 (8–16) 462 34 26 (20–28) 11 1.3 (1.2–1.7) Timofeeff et al. (2006); Lowenstein and Timofeeff (2008, with references) Timofeeff et al. (2006) 121.0–112.2 565 8.5 (5–12) 416 42 35.5 (32–39) 11 1.2 (1.1–1.3) Permian (Tatarian) Permian (Artinskian– Kungurian) Permian (Asselian–Sakmarian) Late Silurian 258–251 283–274 565 565 23 (18–26) 19 (13–22) Ignored Ignored 469 439 52 60 14 (9–17) 17 (11–20) 10 10 3.7 (3.1–5.8) 3.5 (3.0–5.5) 296–283 423–419 565 565 20 (15–24) 10 Ignored 461 420 52 45 15 (10–19) 33 10 11 3.5 (2.7–5.2) 1.4–2 Timofeeff et al. (2006); Lowenstein and Timofeeff (2008, with references) Lowenstein et al. (2005) Lowenstein et al. (2005); Lowenstein and Timofeeff (2008, with references) Lowenstein et al. (2005) Brennan and Lowenstein, 2002; Lowenstein and Timofeeff (2008, with references) The values were calculated under assumptions that: m(Cl#) was equal to modern seawater; m(Kþ) was equal 10 mmol kg#1 H2O for the Permian seawater, and 11 mmol kg#1 H2O for the Cretaceous and Silurian seawater; for the other assumptions, see references. Geochemistry of Evaporites and Evolution of Seawater Table 8 525 Geochemistry of Evaporites and Evolution of Seawater Mg 60 50 40 30 20 10 0 (a) 0 100 200 300 400 500 600 Ma Ca 50 40 5 4 3 2 1 0 30 (e) 0 100 100 200 300 400 500 30 m(SO42-) -1 (mmol kg H2O) 500 600 Ma Aragonite 0 m(K+) -1 (mmol kg H2O) 400 Calcite 0 (d) 300 10 (b) (c) 200 20 20 600 Ma SO4 15 20 10 10 0 Number of occurrences m(Ca2+) (mmol kg-1 H2O) Mg/Ca 6 m.(Mg2+)i/m(Ca2+)i m(Mg2+) (mmol kg-1 H2O) 70 5 0 100 200 300 400 500 20 600 Ma K (f) T Cr J T P P M D S O Cm 10 0 0 100 200 300 400 500 600 Ma Solid circle - value from analyses, open circle - assumed value Figure 17 (a–d) concentrations of major ions (Mg2þ, Ca2þ, SO42#, and Kþ) in the Phanerozoic seawater restored from halite fluid inclusions, after Horita et al., 1991 (in orange), Zimmermann, 2000 (in yellow), Brennan and Lowenstein, 2002 (in pink), Horita et al., 2002 (in blue), and (e) diagram showing oscillations of calculated Mg/Ca ratio in Phanerozoic calcite and aragonite seas, after Lowenstein et al., 2001 (in red), Horita et al., 2002 (in blue), and Timofeeff et al., 2006 (in green), with references. Circles, triangles, and thick-thin vertical bars in figures (a-e) are based on the assumption of different values for m(Ca2þ)I · m(SO42#)i, (f) occurrence of marine calcite and aragonite ooids in Phanerozoic, after Wilkinson et al., 1985, time scale modified after Ogg et al., 2008. models are tested and compared with the chemical restorations of seawater chemistry from halite fluid inclusions. However, so far, the reliable reconstructions from such inclusions are limited only to several ‘points’ on the stratigraphic scale, the large intervals in between them remain without any certain data from halite brine inclusions (Figure 17). The basic assumption of the described interpretational strategies was that seawater in the marginal marine evaporite basins reached the stage of halite crystallization strictly preserving its marine character. In modern marine evaporite environments, we can find examples that support this idea, but there also are some that do not support that view. Marine halite brine is easily modified in its composition in small peripheral evaporite pans such as karst solution basins on the coast of the Mediterranean (Nadler and Magaritz, 1980). On the other hand, seawater seeping through the barrier into one of the largest marginal marine basins – MacLeod basin (Australia) – is nearly the same in composition as the open Indian Ocean water (Logan, 1987, his Table 5). The evaporating seawater brines at various stages of concentration within this basin (up to halite saturation) are similar in composition (but not exactly the same) as the original Indian seawater brines (Figure 19; Logan, 1987). Some ions in the basinal seawater brines commonly show slight deviations from the expected values for these components, presumably due to salt recycling, particularly in case of Na and Cl (Logan, 1987, his Table 5). Similar deviations are 526 Geochemistry of Evaporites and Evolution of Seawater Figure 18 (a–d) Evolution of the composition of Phanerozoic evaporating seawater brine at the halite precipitation stage (mol%) estimated from primary fluid inclusions in marine halite shown on the Mg–2K–SO4 and Mg–Ca–2K Jänecke diagrams at 25 " C, (a) for 0–150 Ma, (b) for 150–250 Ma, (c) for 250–390/410(/530) Ma, and (d) for 390/410 (/530)–550 Ma (redrawn, with corrected Badenian age, from Horita J, Zimmermann H, and Holland HD (2002) Chemical evolution of seawater during the Phanerozoic: Implications from the record of marine evaporates. Geochimica et Cosmochimica Acta 66: 3733–3756), mSW-modern seawater brine at the halite precipitation stage. (e–f) Mg, SO4, and Ca concentrations in modern seawater of SO4rich type (e) and in ancient seawater of Ca-rich type (f), representing two extreme types of seawater in the Phanerozoic (redrawn from Kovalevych VM and Vovnyuk S (2010) Fluid inclusions in halite from marine salt deposits: Are they real micro-droplets of ancient seawater? Geological Quarterly 54: 401–410 and references cited therein). Modern seawater in (e) after Holland (1984). Ca-rich seawater in (f) calculated (based on data in Kovalevych et al., 1998a; Horita et al., 2002; Lowenstein et al., 2001, 2003) under assumption that Na and Cl contents (which made up about 90% of total ion content in modern and ancient seawater types) did not change significantly. K content was constant. recorded in lagoon-type basins – Bocana de Virrilá in Peru (Figure 20; Brantley et al., 1984) and Ojo de Liebre in Mexico (Geisler-Cussey, 1997; Pierre et al., 1984a). Similarly, in Mediterranean saltworks, Naþ, Cl#, and SO42#, unlike conservative Mg2þ, show remarkable deviations in concentrations in more saline brine, that is, they do not create perfectly coincident crystallization paths. In some modern basins, potassium shows very remarkable deviations from expected concentrations (Geisler-Cussey, 1997; Herrmann et al., 1973; Nadler and Magaritz, 1980), and is early lost during evaporation presumably through ion exchange with clay minerals (Hardie and Eugster, 1970, p. 288). However, as it was already mentioned, Timofeeff et al. (2006) strongly believed that “the great mass of dissolved salt in large brine bodies,” that is, in saline giants, “make Geochemistry of Evaporites and Evolution of Seawater 527 7000 Na+ ClMg2+ MacLeod SO422+ Ca BrClNa+ SO42Ca2+ Mg2+ K+ Br- Concentration (mMol kg-1-H2O) 6000 5000 4000 3000 Start of halite precipitation 2000 Start of gypsum precipitation 1000 0 0 5 10 15 Degree of evaporation 300 Concentration (mMol kg-1-H2O) 250 200 Start of gypsum precipitation Start of halite precipitation 150 100 50 0 0 5 10 15 Degree of evaporation Figure 19 Comparison of the crystallization paths of the Caribbean seawater based on data by McCaffrey et al., 1987 with the geochemical characteristic of the basinal brines of the MacLeod basin (after Logan, 1987). modification of the major-ion chemistry by syndepositional recycling processes or nonmarine inflow waters less likely than in shallow/ephemeral systems” (Timofeeff et al., 2006, p. 1982). These authors introduced the special criterium, described earlier (Section 9.17.20.1), for exclusion of data representing the supposed modified parent seawater from the ‘true’ seawater brine. The other problem is the use the complete set of criteria for the proper recognition seawater brine in halite inclusions. In spite of the fact that the criteria for the seawater signal in halite inclusions are numerous and rigorous in the majority of the important papers with apparent successful interpretation of the chemistry of ancient seawater (see extensive sections earlier), these criteria, particularly concerning actual sampled sections and geology of the halite basins, are not discussed. The ‘screening’ procedure for the selection of the samples was sometimes described very poorly or was omitted. Some 528 Geochemistry of Evaporites and Evolution of Seawater 7000 ClNa+ SO42Ca2+ Mg2+ K+ BrMg2+ SO42Bocana Clde + Na Virrilá Ca2+ K+ Concentration (mMol kg-1-H2O) 6000 5000 4000 3000 Start of halite precipitation 2000 Start of gypsum precipitation 1000 0 0 5 10 Degree of evaporation 15 300 Concentration (mMol kg-1-H2O) 250 200 Start of gypsum precipitation Start of halite precipitation 150 100 50 0 0 5 10 Degree of evaporation 15 Figure 20 Comparison of the crystallization paths of the Caribbean seawater (based on data by McCaffrey et al., 1987) with the geochemical characteristic of the basinal brines of the Bocana de Virrilá (after Brantley et al., 1984). restorations were made without showing overlapping crystallization paths from basins of the same age considered as crucial criterion for proper recognition of seawater derivation of brine in fluid inclusions. In fact, attempting to synthesize all the available data on halite fluid inclusions with seawater brine, Horita et al. (2002) found it very difficult (and in fact impossible) to identify the evaporite deposits that meet all the criteria for primary seawater trapped in halite inclusions, as listed in the earlier sections. Therefore, they “only consider halite from evaporite deposits whose Sr and S isotope signature indicates Geochemistry of Evaporites and Evolution of Seawater unequivocally that they are marine in origin” (Horita et al., 2002, p. 3734). Many authors described the Br contents in halite as the main argument in favor of the marine origin of salt. Such solutions were earlier criticized by several authors who have shown that isotope and trace element data are inconclusive in this respect (Hardie, 1984; Schreiber and El Tabakh, 2000; Warren, 2006). The exclusively marine derivation of many evaporites sampled for marine brine in halite inclusions still remains controversial (compare, e.g., Brennan and Lowenstein, 2004; Lowenstein et al., 2001; Schoenherr et al., 2008). The weak point of all these restorations is that they require too many uncertain assumptions concerning the ‘starting’ composition of the ancient seawater. As noted by Steuber and Rauch (2005, p. 200), “experimental data on major ion composition of palaeo-seawater are still scarce, have a coarse temporal distribution, and require assumptions on the composition of evaporating brines, resulting in some uncertainty for the reported values.” The restorations of the concentration of Mg and Ca ions in paleoseawater were based on a great number of uncertain assumptions, so Tyrrell and Zeebe (2004) called these restorations ‘best guess’ rather than proved, although they emphasized that they are well supported by many other facts from the associated nonevaporite record. The empirical data were compared with the numerical models of the evaporating seawater evolution in the ideal system. These numerical models were successfully tested on data from fluid inclusions in the modern Inagua saltworks and in a salt pan on the supratidal sabkha (Timofeeff et al., 2001). McCaffrey et al. (1987, p. 937) stated that “the sequence of mineral formation and the evaporation path of seawater defined by the Inagua brines largely corresponds to the theoretical fractional crystallization path of seawater described by Eugster et al. (1980) and Harvie et al. (1980),” which was confirmed by Timofeeff et al. (2001). The model, however, was not tested in a basin on the scale of an ancient saline giant. Saltworks cannot be exactly compared with such a basin. The important difference is that halite pans in solar saltworks are supplied by gypsum brine already stripped of calcium, that is, having composition different than seawater (McCaffrey et al., 1987; Timofeeff et al., 2001), whereas the ‘realistic’ halite basin is expected to be supplied directly by seawater (Figures 4(a) and 4(b) and 5), as in the scenario considered by Holser (1979a). Finally, there is always a danger that the ancient halite inclusions analyzed simply do not represent the evaporated seawater or basinal water (Vovnyuk and Kovalevych, 2007). Von Borstel et al. (2000) recognized that several ‘network’ fluid inclusions from modern marine halite from solar saltworks showed highly variable chemistry and higher concentration that the brine from which they crystallized. These authors explained that such inclusions evaporated after sampling (Timofeeff et al., 2001). Some ‘anomalous’ modern halite fluid inclusions from Baja California deviate from the evaporation paths predicted by computer programs, and many others show more or less broad scatter (Timofeeff et al., 2001). Ayora and other authors who are specialized in restoration of the chemical evolution of the basinal waters in marginal marine evaporite basins are warned about the uncritical acceptance of the halite fluid inclusion data, particularly from single samples, as the evidence of clear uncontaminated record of ancient ocean chemistry. The essence of their criticism is the requirement of complete and holistic sedimentological and 529 geochemical analysis of the evolution of the basinal waters to show that the signal in the sample is exclusively from uncontaminated seawater (point 6 discussed earlier; Ayora et al., 1994, 1995). Ayora et al. (2001) were able to restore the chemical evolution of brine in several Mesozoic and particularly in Tertiary evaporite basins based on the analysis of mineral associations, primary fluid inclusion analysis in halite, and numerical simulation of the model marginal marine evaporite basin (described in Section 9.17.7.2). They explained the chemical composition of brine trapped in halite exclusively by processes of alteration of the marine water of the present-day composition, that is, by sulfate depletion related to dolomitization or addition of CaCl2-rich brine to basinal waters, or other processes. They noted that sulfate depletion observed in brine from fluid inclusions varied in intensity in basins of the same age, as well as throughout the evolution of the same basin. The studied basins were relatively small in comparison with the largest saline giants. However, recently, the same features were documented in the giant sequence of Permian Zechstein cyclothems (Garcı́a-Veigas et al., 2011). Some of these variations cannot be explained by global and ‘secular’ variations in ocean chemistry, the variations are too sharp and the time spans too short to achieve global mixing of the oceans. Also young or subfossil sequences of the marginal marine basins show deviations from the sequence expected from evaporation of present-day seawater, like the potash evaporites from Dallol, Danakil Depression, whose lower part is made up of sulfate and contains kainite, whereas the upper part is primarily chloride and composed of sylvite (Hardie, 1990; Holwerda and Hutchinson, 1968). Cendόn et al. (2004) pointed out that “evaporitic successions have to be proven marine before they can be confidently used to deduce seawater palaeochemistry,” and Ayora et al. (2001, p. 251) concluded that “the solute proportion recorded in the fluid inclusions can be explained by the evaporation of present day seawater as a major recharge” and therefore the “changes in potash mineralogy and sulfate depletion in fluid inclusions are not conclusive arguments in favor of secular variations in the composition of the ocean.” Cendón et al. (2008) stated that results of their own work in Mulhouse basin proved that the chemical changes within the basinal waters were different within the individual subbasins of the similar age. These changes were also too rapid to be explained by any global secular variations of the ocean chemistry. They again stated that this “precludes the use of isolated fluid inclusions samples as a proxy of ancient ocean composition” and in particular “it precludes the use of fluid inclusions in isolated samples to reconstruct the composition of the Oligocene ocean” (Cendón et al., 2008, p. 111), because two halite samples from Mulhouse basin were used previously to back calculate the chemistry of Oligocene oceanic water from fluid inclusions by some other authors. On the other hand, in the modeling of the Mulhouse basinal water evolution during evaporation, these authors, agreeing that “identifying potential end-member water chemistry in an ancient evaporite basin is difficult,” just used the modern seawater for modeling “as there is no experimental and independent (non-evaporite based) data for Oligocene seawater composition available” (Cendón et al., 2008, p. 116). A similar integrated approach of restoration of the chemical evolution of the basinal water and seawater signal in it was recently made for the Polish Permian basin by Garcı́a-Veigas et al. (2011). 530 Geochemistry of Evaporites and Evolution of Seawater However, the following evidences can be found in favor of secular variations of the chemical composition of seawater (Kovalevych et al., 2006a): 1. The major compositional changes of brines in fluid inclusions in ‘marine’ halites show clear stratigraphic control, irrespective of paleogeographic position of the basin (see summary by Kovalevych et al., 1998a,b). 2. The halite brines of the marine Neogene basins show such stratigraphic control and are of the SO4-rich type (the same as present-day seawater halite brine), because in Neogene time, seawater was unambiguously of the same type as today. 3. The uniform trend of changes in composition of brine was detected in marine Neogene halite inclusions showing that during last !40 My, the concentration of oceanic Mg2þ was rising (Horita et al., 2002; Zimmermann, 2000), although some data from that time interval can be questioned (Cendón et al., 2008). 4. The changes in the major element chemistry of ancient seawater coincide or overlap in time with major variations in the mineralogies of marine nonskeletal carbonates (ooids and cements) and also mineralogy of potash evaporites, changes of isotopic composition of some elements, and other geologic processes in the Phanerozoic (Kovalevych, 1990). In particular, the variations of carbonate and potash salts mineralogies apparently had the same shared causative reasons and tied to the fluctuations in Mg/ Ca ratio in seawater. 5. The apparent lack of extensive contemporaneous dolomite in many evaporite basins speaks against the importance of dolomitization for changes in basinal brine composition. 6. The variations in brine composition and especially intensity of sulfate depletion in separate basins of the similar age (as recorded by Garcı́a-Veigas et al., 1995; Ayora et al., 2001) can be explained by the influence of local factors, such as water–rock interactions and inflow of nonmarine water. 9.17.22 Salinity of Ancient Oceans Na and Cl are the most abundant elements in the seawater responsible for its salinity. NaCl is responsible for the salty taste of marine and other waters. Early geochemical calculations (Goldschmidt, 1937; Rubey, 1951, 1955) suggested that the amount of Cl in the seawater is so large that it “cannot be assumed to be derived entirely from weathered igneous rocks” but was already present in primitive atmosphere of the early Earth (Goldschmidt, 1954, p. 66). Cl is an incompatible element that presumably was outgassed as HCl together with H2O during the earliest Earth history (Holland, 1984; Knauth, 2005). Most probably, the entire inventory of Cl was present in the ionic form in the early ocean, until the time of deposition of huge evaporite formations on accreting continents in the Paleoproterozoic (Knauth, 1998). Cl# has the longest residence time among major elements present in seawater – estimated as 2.27 ( 108 year (¼227 My) by Land (1995), comparable with Br (100 My: Holland et al., 1986; Br is considered by Holland et al., 1996, as showing constant concentration in Phanerozoic seawater). Naþ, however, is lost continuously during hydrothermal circulation of seawater through the basaltic cover of the mid-ocean ridges, being sequestered in the newly formed crust mainly in the process of albitization of plagioclases (Cowen, 2000; also see Chapter 8.7). Cl# however remains relatively unchanged during this circulation – it is therefore particularly conservative element in seawater, and its concentration was probably relatively stable or changed very slowly through geologic time (cf. Holland et al., 1986). All these factors suggest that the ocean was remarkably salty, that is, contained a great deal of Cl in the Phanerozoic and perhaps since the beginning of ocean creation. The opposite position to this view is the vanished soda ocean hypothesis (Kazmierczak et al., 2004; Kempe and Degens, 1985; Kempe and Kaźmierczak, 1994, 2011) that assumes the existence of high amounts of CO2 in the oldest atmosphere and that implies the carbonic acid weathering of silicates on the early Earth, according the Urey reaction. Consequently, there was the production of huge amounts of carbonate and bicarbonate anions in seawater (HCO3# þ CO32# > Cl# þ SO42#) causing an oceanic pH as high as 11. According to this hypothesis, the main driving force for the transition of the presumed soda ocean into the present-day halite ocean during Proterozoic was the subduction of seawater (as pore water) together with oceanic crust and sediments in subduction zones (e.g., Lécuyer et al., 1998; Pope et al., 2012), and the subsequent formation of continental crust with accumulated carbonates and organic carbon (Kempe and Kaźmierczak, 1994). The concept implies a slowly but continuously growing content of Cl# beginning with the lowest values in the Hadean to today’s values and with a substantial amount of Na in the vanished early ocean (Naþ þ Kþ > Ca2þ þ Mg2þ; Kempe and Kazmierczak, 2011). Kempe and Kaźmierczak (1994) suggested that calcium concentration could rise in early ocean attaining some critical level that induced skeletogenesis in marine animals in Cambrian, in response to increased Ca2þ ‘stress.’ Morse and Mackenzie (1998) agreed with the concept of gradual calcium concentration rise in the early ocean; however, they believed that the ocean was always NaCl-dominated, as today, and that pH was lower than now due to a higher amount of CO2 in the early atmosphere. Unfortunately, a scarce Precambrian evaporite record and the lack of unquestionable marine evaporite deposits in the earliest rocks do not permit recognition of the true chemistry of the earliest ocean. Based on the other evidence and theoretical models, currently most authors believe that the chloride was dominant anion in seawater that was as salty (or saltier) as today since the Archean (Foriel et al., 2004; Hardie, 2003; Holland and Kasting, 1992; Knauth, 2011). Many authors, modeling the history of the ocean chemistry, assume that the volume of the ocean was more or less constant or that it continuously grew since the time of its creation (see Mason, 1958; Pinti, 2006; Rozanov, 2010; Schopf, 1980, and references in these publications). However, the hypotheses that the ocean volume decreased or oscillated with time are recently also accepted (Ingebritsen and Manning, 2003; Knauth, 2011; Lécuyer et al., 1998; Pope et al., 2012). The volume of ocean has oscillated slightly due to global glaciations. During Quaternary, !2% of the seawater volume was incorporated in the ice sheets (Hay et al., 2006) and could have caused an average salinity rise in the ocean from 35% up to 36–37.6% (Hay et al., 2006). According to Stevens (1977), Pleistocene salinity variations did not exceed 1.5%. The volume of the preserved marine evaporite deposits also was used for the calculation of the salinity of ancient ocean Geochemistry of Evaporites and Evolution of Seawater should lead first to lowering of SO42# concentration in the ocean, because this ion is always sequestered in the deposited gypsum before halite precipitation. Apparently, however, this sequestration does not always take place. Hansen and Wallmann (2003) suggested this as the cause of the lowered concentration of seawater sulfate and calcium !20 Ma. Wortmann and Chernyavsky (2007) also recognized the influence of such diminution, caused by the substantial Ca sulfate evaporite deposition in the Early Cretaceous (Aptian), on the global geochemical S and C cycling in that period. Deposition of the 1.125–1.6875 ( 106 km3 of salts during the Messinian salinity crisis (Ryan, 2008), ! 5% of salt content of the ocean (Ryan, 2009), could also depress the average salinity of the ocean. Holser (1984) estimated that salinity dropped rapidly four or five times between the Permian and the Cretaceous by 1–4% due to evaporite deposition (mainly NaCl). Stein et al. (2000) calculated that global acmes of evaporite deposition could disturb the isotope 87Sr/86Sr ratios in seawater by refluxing brine flowing out from the largest saline giants, such as the late Permian Zechstein basin, the Callovian Louann salts in the Gulf of Mexico, and the Messinian Mediterranean basin. under assumption that the ocean volume was constant since the beginning. If we agree with this assumption and accept that all the evaporite minerals found today in sedimentary rocks were present in dissolved form in the ocean, from its inception (with the volume comparable to today), we can calculate that the salinity of early ocean was about twice today’s salinity, that is, !70% (Knauth, 1998, 2005, 2011). Holland (1984, p. 461) roughly estimated that the average salinity of the Phanerozoic seawater was no more than 30% higher than today, that is, it was less than 45.4%. Based on the amount of evaporitic deposits in the geologic record, Hay et al. (2006) calculated that the salinity of the Phanerozoic ocean varied between 35% and 47%, and only in the Cretaceous period could it have dropped to 32–33%. They presented the model of salinity changes since Cambrian. The fluid inclusions in late Cambrian–early Ordovician carbonate cements show salinity within the range 31–47% (Johnson and Goldstein, 1993), which overlap and generally coincide with the range predicted by this model (Figure 21). Knauth (2011) suggested that the estimates by Hay et al. (2006) are probably too high and that “the idea that Paleozoic life could thrive at such high inferred salinity is likely to be resisted by marine biologists and paleontologists” (Knauth, 2011, p. 771). The extremely large volume of evaporites in Neoproterozoic and Permian possibly could temporarily lower the salinity of the ocean for several per mill (%) during those time intervals (Fischer, 1964; Holser, 1984; Vickers Rich, 2007). Stevens (1977) calculated the volume of Permian halites known at that time as nearly 1.6 ( 106 km3 and estimated that the Permian salinity dropped to 31.5%, that is, about 10%. Sulfate in Zechstein sediments is also equal to !10% of the sulfate content of the present ocean (Schaffer 1971 cited by Holland, 1972). Luo et al. (2010), based on the compilation of data by Hay et al. (2006), estimated that late Permian deposition of Ca sulfate could lower the concentration of SO42# in the ocean about 6 mM. Holser (1984) and Holland et al. (1996) suggested that the increased evaporite deposition in some periods 9.17.23 Evaporite Deposition through Time The largest saline giants are preserved in deposits formed from late Ediacaran through the Cenozoic (Table 9). According to available data, there are only four basinal areas with the volume of salts over 1 100 000 km3 recorded in that time interval, and the Gulf of Mexico basin (160 Ma) is the largest one containing 2 400 000 km3 of salt (Evans, 2006). The next in size appears to be the Messinian evaporites of the Mediterranean and Red Sea region (Rouchy and Caruso, 2006) estimated on 1 400 000 km3 in volume (mean calculated from data by Ryan, 2008). Among the pre-Ediacaran saline giants, the largest volume of recorded salts is found in Centralian Superbasin in Australia (!800– 830 Ma; the Bitter Spring Fm. and its equivalents; Lindsay, Range of salinity from various models 50 Mean salinity of ocean in ‰ 531 50 40 40 Salinity from fluid inclusions in marine calcite 30 30 Recent average salinity 34.7‰ 20 20 10 10 0 0 pCm Cm Neoproterozoic 600 O S D C P Paleozoic 500 400 T J Cr 200 N Cenozoic Mesozoic 300 P 100 0 Age (Ma) Figure 21 Reconstruction of the mean salinity of the ocean during the Phanerozoic according to Hay et al. (2006). Salinity of Cambrian–Ordovician seawater from fluid inclusions after Johnson and Goldstein (1993). 532 Geochemistry of Evaporites and Evolution of Seawater Table 9 The world’s largest evaporite basins; compilation based on various sources, repeated after Evans (2006), and supplemented after Ryan (2008) Precambrian (pre-Ediacaran; >600 Ma) Cenozoic–Mesozoic (0–250 Ma) 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 Evaporite basin Age in Ma Volume in km3 Messiniana Red Sea SW Iran S Mozambique E China Rus, Arabia N Sahara Indochina S Atlantic Hith, Arabia Central Asia Andes Gulf of Mexico Alan, Arabia Tanzania N Sahara Keuper Jilh, Arabia S China 5 10 20 20 40 50 90 100 120 150 150 160 160 180 200 200 225 230 230 2 250 000a 900 000 300 000 27 000 20 000 200 000 32 000 50 000 35 000 360 000 250 000 40 000 2 400 000 20 000 150 000 710 000 50 000 120 000 80 000 Permian–Carboniferous (250–360 Ma) 1 2 3 4 5 6 7 8 Evaporite basin Age in Ma Volume in km3 Zechstein Khuff, Arabia E European Peru–Bolivia Midcontinental USA Amazon Sverdrup Canadian Maritime 250 260 270 270 270 300 315 340 200 000 75 000 1 100 000 62 000 81 000 25 000 120 000 46 000 Evaporite basin Age in Ma Volume in km3 E European Taimyr W Canada Morsovo Michigan Canning Canadian Arctic Mackenzie Morocco–Iberia Siberia Persian Gulf Salt Range 370 370 390 400 420 440 460 500 520 520 545 !550 1 100 000 18 000 86 000 81 000 29 000 26 000 19 000 110 000 50 000 800 000 500 000 240 000 Evaporite basin Age in Ma Volume in km3 Skillogalee, S Australiab Curdimurka, S Australiac Kilian-Redstone River, Canadad ca. 770 ca. 785 ca. 770 25 000 50 000 30 000 Devonian-late Ediacaran (360–600 Ma) 1 2 3 4 5 6 7 8 9 10 11 12 Precambrian (pre-Ediacaran; >600 Ma) 1a 1b 2a (Continued) (Continued) Table 9 2b 3 4 5 6 7 8 9 10a 10b 10c 10d 10e 11 12 13 14 15 Evaporite basin Age in Ma Volume in km3 Minto Inlet, Canadae Duruchaus, Namibiaf Copperbelt, Central Africag Centralian, Australiah Borden, Canada Char/Douik, W Africa Belt basin, N America Discovery, W Australia Balbirini, N Australia Lynott, N Australia Myrtle, N Australia Mallapunyah, N Australia Corella, N Australia Stark, Canada Rocknest, Canada Juderina, W Australia Tulomozero Chocolay-Gordon Lake succession, Superior craton, Canada–USA ca. 800 ca. 800 ca. 830? 90 000 15 000 25 000 ca. 800 ca. 1200 ca. 1200? 1460 ca. 1500 1610 1635 1645 1660 1740 ca. 1870 ca. 1950 ca. 2100 ca. 2100 ca. 2250 140 000 15 000 8000 10 000 -2800 2500 3000 13 000 5000 2000 30 000 1000 1000 1000 4500 Remarks to some pre-Ediacaran basins after Evans (2006), see Table 11 for further information. Lack of or highly controversial data are marked by “?”. a After Ryan (2008). b Bedded magnesite, ca. 500 m in thickness, in rift. c Abundant pseudomorphs after anhydrite, halite, and shortite, up to 1 km in thickness over an area of ca. 50 000 km2, in the same rift. d ca. 100 m in thickness over an area of 300 000 km2. e Evaporites attain a thickness of ca. 300 m across an interpolated depositional area of 300 000 km2. f ca. 500 m thick and about 30 000 km2 evaporite succession. g Basin-scale evaporite solution megabreccias, inferred thickness of the evaporites ca. 500 m, over an area of 50 000 km2. h The most extensive pre-Ediacaran gypsum, anhydrite, and halite deposits, with a typical thickness ca. 800 m, covering an aggregate area of ca. 140 000 km2. 1987; Stewart, 1979) that contain !140 000 km3 of evaporites (Evans, 2006). The lake evaporites attain smaller but remarkable volumes, as, for example, 2.5 km thick halite deposits of the Hualapai basin, Arizona, United States, with 200 km3 volume of salt (Faulds et al., 1997), the other examples are the Dead Sea or central Andean basins. Large accumulations of K–Mg salts are rare and marine evaporites containing volumetrically important K–Mg salts occur in only about 40 Phanerozoic basins (Table 10; Goncharenko, 2006; Vysotskii et al., 1988; Warren, 2010). Precambrian evaporites are thought not to contain K–Mg salts (Muir, 1987). Glauberite occurs with halite in Sinian (late Neoproterozoic) evaporites in Sichuan, China (XiaoSong, 1987). 9.17.23.1 Late Ediacaran–Phanerozoic Marine Evaporites The marine evaporite record from late Ediacaran to Recent time shows a characteristic pattern of changes. From Cambrian to the lower half of the Permian (including Artinskian), the potash deposits were only of the chloride type, with sylvite Table 10 Major evaporite basins with potash salts deposits, their volume and chemical character, compilation of various sources, repeated after Hardie (1990), and Vysotskii et al. (1986), and supplemented (after Hardie, 1996; Harville and Fritz, 1986; Hryniv et al., 2007; Land et al., 1995; Lowenstein and Spencer, 1990; Petrychenko et al., 2005, 2012; Rahimpour-Bonab et al., 2007; Talbot et al., 2009b; Timofeeff et al., 2006; Valyashko, 1962; Warren, 2006) Location Qaidam Basin, China Danakil Depression, Ethiopia Kaidak basin, Kazakhstan Dead Sea basin Sicily, Italy Mediterranean Sea Erevan basin, Armenia Carpathian Foredeep, Ukraine–Romania Gabon Basin, West Africa Sergipe-Alagoas basin, Brazil Houston Formation with Sylvinite Member Volume of potash deposit Mineralogy (all deposits contain halite) Chemical character Holocene Pleistocene (< 1 Ma) >60 km3 >40 km3 (>30 km3 of sylvinite) ? Minor ca. 50 km3 ? ? ? ca, (sy), (mi), (Po) ka, ca, sy, ks, (Po), (rh), (bi) KCl MgSO4–KCl MgSO4-KCl KCl MgSO4 MgSO4 KCl MgSO4 ? ? ca. 22 km3 ca, mi, bi (sy), (ca) ka, ca, sy, ks, Po, Bi, lg Po, ka ca, sy ka, lg, sy, ks, Po, (ca), (gs), (bl), (lw), (mi) sy, ca, Po, (lg), (gs) (Po) sy, (ca) Vorotyshcha and Kalush formations (Ukraine) Upper Red Formation Lower Fars Formation Zone Salifere Late Pliocene? Late Pliocene to early Pleistocene Late Miocene (Messinian) Late Miocene (Messinian) Middle Miocene Early and middle? Miocene (Eggenburgian/Badenian?) Middle Miocene Early Miocene Early Oligocene Maha Sarakham Formation Late Eocene Late Eocene Late Cretaceous ? >80 km3 >1.5 ( 106 km3 ca, sy ca, sy, (Po) ca, sy, Tc, (B) KCl KCl KCl–CaCl2 * 2MgCl2 * 12H2O Early Cretaceous (Aptian) ca. 5 ( 106 km3 of K-bearing salts ? ca, sy, Tc, bi KCl-CaCl2 * 2MgCl2 * 12H2O ca, Tc, bi KCl–CaCl2 * 2MgCl2 * 12H2O 50–200 km3 of carnallite ? ca, Tc, sy KCl–CaCl2 * 2MgCl2 * 12H2O sy, (Po) KCl ? ca, sy, (rh) KCl ? ? ? ? sy ca, sy, Po, ks sy ca, sy, rh, (ks), (bi), KCl KCl–MgSO4 KCl KCl Sedom Formation Solfifera series Salt beds between Chela Series and Mavuma Beds Salt beds between Cocobeach and Madiela formations Ibura Member, Muribeca Formation Forecaucasian (Ciscaucasian) Basin Middle Asian Basin, Turkmenistan– Uzbekistan–Tajikistan Gulf of Mexico, USA Aquitanian basin, France Northern Sahara Salt Basin, Algeria Moroccan Meseta basins Age Early Cretaceous (Aptian) Early Cretaceous (Aptian) Late Jurassic (Tithonian, Kimmeridgian?) Late Jurassic Louann formation salt Trias a evaporites (salina d’Ourgla) Middle Jurassic (Callovian) Late Triassic (Keuper) Late Triassic (Carnian-Norian) Late Triassic (Carnian-Norian to perhaps early Jurassic) KCl MgSO4 KCl (Continued) Geochemistry of Evaporites and Evolution of Seawater Great Kavir Basin, Iran Iran Rhine Graben, Germany, Mulhouse basin, France Navarra Basin (Ebro Basin), Spain Catalan Basin (Ebro Basin), Spain Khorat Plateau (Khorat and Sakon Nakhon basins), Thailand Congo Basin, West Africa Formation 533 534 (Continued) Location Formation Age Volume of potash deposit Mineralogy (all deposits contain halite) Chemical character English Zechstein basin Teesside group (Z3, Leine), and Staintondale group (Z4, Aller) Zechstein (Z1, Werra; Z2, Stassfurt, Z3, Leine; Z4, Aller) Salado Formation Late Permian (early and late Tatarian) Late Permian (Kazanian to Tatarian) Late Permian (Tatarian) ? sy, (ca), (rh) KCl 2000 km3 MgSO4 Iren horizon Early Permian (Kungurian) ? Iren horizon, Berezniki Formation Early Permian (Kungurian) 119 km3 ca, sy, Po, ks, lg, ka, bl, gs, lw, (rh) sy, lg, Po, ks, ca, ka, bl, le, lw, gs ca, Po, sy, bi, ka, lg, ks, gs, bl, lw, (B) sy, ca Iren horizon Supai Formation (Upper) Early Permian (Kungurian) Early Permian (Leonardian ¼ Artinskian) Early Permian ? 0.1 km3 sy, ca sy KCl KCl ? sy, ca, ks, (bi) MgSO4–KCl Early Permian (Sakmarian) Early Carboniferous to early Permian Carboniferous, middle to late Pennsylvanian (Desmoinesian– early Virgilian) Carboniferous, middle Pennsylvanian (Desmoinesian) Carboniferous, early Mississippian (early Visean) Middle Devonian Middle Devonian (Givetian) 7.35 km3 ? ca, ks, Po, sy, bi, B, (lg) sy MgSO4–KCl KCl 0.1 km3 sy, ca KCl 450 km3 sy, ca, Po, (ks), (Rh), (B) KCl 30 km3 sy, ca, (rh), (Po), (B) KCl ? 3000 km3 sy sy, ca KCl KCl Late Devonian (Frasnian– Fammenian) Middle Devonian (Eifelian) 6000 km3 sy, ca KCl 0.02 km3 sy, (ca) KCl ? sy, (rh) KCl 1000 km3 sy, (ca), (Po), (B) KCl N.W. European Basin Delaware Basin, New Mexico and Texas, USA Pericaspian Basin, Russia, Kazakhstan Upper Kama Basin (Solikamsk, Cis-Ural trough), Russia Upper Petschora basin, Russia Supai Basin, Arizona, USA Pripyat trough, Belarus Dnipro-Donets depression, Ukraine Amazon basin, Brazil K-bearing deposits correlated with Kramators’k Formation Kramators’k Formation Nova Olinda Formation Eagle basin, Colorado, USA Eagle Valley Evaporite Paradox basin, Colorado and Utah, USA Paradox Formation (Hermosa Group) Canadian Maritimes (Moncton Basin), Canada Adavale Basin, Queensland, Australia West Canadian basin (Elk Point basin), Canada–USA Pripyat trough and Dnipro-Donets depression, Belarus–Ukraine Morsovo basin, Moscow syneclise, Russia Tuwa basin, South Siberia, Russia Cassidy Formation, Windsor Group Michigan Basin, USA–Canada Boree Salt Member, Etonvale Formation Prairie Evaporite Lower and upper salt units Morsovo Salt Member Ikhedushiihgol Formation Salina Group Middle Devonian (late Eifelian– early Givetian) Middle to late Silurian (Wenlock– Pridoli) 500 km3 KCl–MgSO4 KCl–MgSO4 KCl Geochemistry of Evaporites and Evolution of Seawater Table 10 East Siberia, Russia East Siberia, Russia Sar Pohl, Iran Angara Formation Usolye Formation, and other formations Hormoz Salt Salt Range Basin, Pakistan Salt Range Formation 25 km3 15 km3 ? ca, sy ca, sy, (rh) sy, rh KCl KCl KCl lg, ka, (sy), (Po), (ks) MgSO4 Geochemistry of Evaporites and Evolution of Seawater Lack of or highly controversial data are marked by “?”. Mineral abbreviations: B – borate minerals; bi – bischofite, MgCl2 * 6H2O; bl – bloedite, Na2SO4 * MgSO4 * 4H2O; ca – carnallite, KCl * MgCl2 * 6H2O; gs – glaserite, Na2SO4 * 3K2SO4; ka – kainite, 4KCl * 4MgSO4 * 11H2O; ks – kieserite, MgSO4 * H2O; le – leonite, K2SO4 * MgSO4 * 4H2O; lg – langbeinite, K2SO4 * 2MgSO4; lw - loeweite, 2Na2SO4 * 2MgSO4 * 5H2O; mi – mirabilite, Na2SO4 * 10H2O; Po – polyhalite, K2SO4 * MgSO4 * 2CaSO4 * 2H2O; rh – rhinneite, FeCl2 * 3KCl * NaCl; sy – sylvite, KCl; Tc – tachyhydrite, CaCl2 * 2MgCl2 * 12H2O; (sy) – within parentheses, mineral aggregates; sy – without parentheses, rock-forming mineral; sy – in bold, economically significant minerals. Chemical character of the deposits: MgSO4 – rich in magnesium sulfate; MgSO4-KCl – mixed or intermediate character, sulfates dominate; KCl–MgSO4 – mixed or intermediate character, chlorides dominate; KCl – poor in magnesium sulfate; KCl–CaCl2 * 2MgCl2 * 12H2O – poor in magnesium sulfate, and containing tachyhydrite. Early Cambrian Early Cambrian Neoproterozoic to Middle Cambrian (Cambrian?) Late Neoproterozoic to early Cambrian 535 536 Geochemistry of Evaporites and Evolution of Seawater and carnallite as the major components (Table 10; Zharkov et al., 1978). In the upper half of Permian, that is, from Kungurian to Tatarian, the chloride sedimentation was accompanied by sulfate deposition (Zharkov, 1981). The Mesozoic potash evaporites are again dominantly of the chloride type, whereas in Neogene, the K–Mg sulfate salts appeared again in the geologic record together with the chloride type (Sonnenfeld, 1984). Thus, the MgSO4-rich evaporites are confined only to the Permian, the Miocene, and the Quaternary (Hardie, 1990). The distribution of marine K–Mg salts in time appears to reflect two Phanerozoic megacycles: the Paleozoic megacycle and Mesozoic–Cainozoic megacycle that both began from long-lasting chloride type of evaporite deposition and abruptly end with deposition of chloride–sulfate evaporites in Permian and from Neogene to today, respectively (Kovalevych, 1990, with references). The Neogene K–Mg sulfate salts contain more sulfate minerals (kieserite, langbeinite, polyhalite, and kainite) than do the Permian salts (Kovalevych, 1990). In some periods of time, an enormous amount of salts was accumulated and these phases of deposition appear to have been spread across much of the planet (Table 9). The geologic record suggests at least two main intervals of increased evaporite deposition in Earth history: 180–250 Ma and 500–700 Ma (Holser, 1984; Knauth, 2005). Evans (2007) suggested two acmes of deposition for the Precambrian-to-Cambrian interval: at ! 800 Ma (in Cryogenian), which produced !350 000 km3 of evaporites, mainly Ca sulfates, and in late Ediacaran to Early Cambrian time, resulted in 1.5 million km3 of mixed Ca sulfate and Na chloride salts. In Phanerozoic, about 40% of all salts were sequestered in the Permian–Triassic interval (Knauth, 2005), and according to Trappe (2000), these evaporites together contain 35% of the world’s evaporite resources. It seems that these evaporite-rich intervals are apparently associated with the paleogeographic configurations that developed during and after the breakup of the two supercontinents: Rodinia – in late Proterozoic (Neoproterozoic) (Knauth, 2005), and Pangea – in Phanerozoic interval (Gordon, 1975). These periods were favorable for evaporite deposition because of the appearance of many enclosed, rapidly subsiding basins in equatorial and circum-equatorial settings (Knauth, 2005; Trappe, 2000). Hay (personal information in Hansen and Wallmann, 2003) calculated that the average global rate of evaporite deposition in Cretaceous and Cenozoic reached maximum in 150–140 Ma (2.315 ( 1018 kg 10#1 My) and 20–10 Ma (3.459 ( 1018 kg 10#1 My) and minimum in 70–60 Ma time interval (0.071 ( 1018 kg 10#1 My). The Messinian evaporites (!5.96–5.33 Ma) are one of the greatest evaporite events on Earth considering that their volume, at least 2.25 ( 106 km3, was deposited in a relatively short time interval – !640 ky – in much shorter time than recognized in any other saline giants (Rouchy and Caruso, 2006; Ryan, 2008). Late Neoproterozoic evaporites occupying Pangea spread from the Indian subcontinent (Rajasthan, Salt Range in Pakistan), Oman (Ara Formation), and Saudi Arabia–Iran (Hormoz, formerly Hirmuz or Hormuz, series) (Horita et al., 2002; Schoenherr et al., 2008; Talbot et al., 2009). Potash salts of both sulfate and chloride type, with polyhalite, kainite, langbeinite, and sylvite and carnallite, are known from the Hanseran Evaporite Group in Rajasthan, and Mg sulfates occur in Salt Range evaporites in Pakistan (Horita et al., 2002, with references). These are probably the oldest known volumetrically significant K–Mg salt deposits, except for the Ara Formation that contains ash beds dated radiometrically at 542.0 + 0.3 Ma and 542.6 + 0.3 Ma (Schröder et al., 2004); the age of these formations is, however, poorly constrained. From the known record of the marine and marine-related evaporites in Earth history (i.e., saline giants), it is evident that the mineralogical and chemical composition of evaporites was surprisingly stable. In particular, the K–Mg evaporites present since the Cambrian evidence both chloride chemistry (Salt Range Formation, Pakistan) and sulfate type of chemistry (Usolye Formation, Russia). That is the same chemistry that is known in majority of K–Mg deposits from the Cambrian until today (Strakhov, 1962; Vysotskii et al., 1988). Based on these observations, it was interpreted that from the end of the Precambrian time onward, both the chemical composition of the ocean and possibly the salinity of the ocean was established and was similar or remained nearly the same as today. There is no record of the irreversible evolution of evaporite composition in Phanerozoic (Strakhov, 1962); however, the record of fluctuation is recognizable both in the K–Mg facies and brine composition in halite fluid inclusions. Two periods that lasted ! 40 My were recognized in the Phanerozoic, lacking any record of marine saline giants: in Ordovician (Zharkov, 1981) and between the Upper Coniacian and the end of Paleocene (Sonnenfeld, 2000). Ordovician is the only Phanerozoic system without recognized potash evaporites (Goncharenko, 2006). 9.17.23.2 Precambrian (Pre-Ediacaran) Marine Evaporites Precambrian evaporites (marine and nonmarine) are mostly represented by pseudomorphs after gypsum, anhydrite, and halite (Zharkov, 2005). Evans (2007) counted about 100 documented examples, including ten from the Archean (Table 11). About 20 deposits have total preserved or estimated salt volumes attaining 1000 km3, and all of them occur in the Proterozoic era (Table 9). The scarcity and lack of evaporites before 2 Ga was explained as the result of selective removal (Gordon, 1975; Hardie, 2003, among others), related at least partly to their very high solubility – they did not survive the metamorphic conditions that have affected these very old rocks. Additional reasons could be following: (1) the lack of extensive platforms at the margins of emerging continents, necessary for large-scale evaporite deposition (Strakhov, 1962), and (2) the fact that there were only a few continents in the Archean and they were relatively small and apparently did not create a supercontinent (Knauth, 2005; Walker, 1985). Note, this view is only a hypothesis (see Armstrong, 1991; Lenardie, 2006) and other authors argue for vey rapid growth of continental crust in that same span of time (Lowe and Tice, 2007). The acceptance of the slow gradual accretion of continents means that the chemistry of the early oceans (<2 Ga) was driven mainly by mantle processes with increasing influence of the input from the rivers on emerging lands only in later Precambrian and continuing up to today (Godderis and Veizer, 2000; Reddy and Evans, 2009). The other consequence is that the assumed high salinity of the Table 11 Inferred and direct evidence of earliest evaporites in Archean through early Mesoproterozoic rocks, based on the compilation by Pope and Grotzinger (2003), Bekker et al. (2006), Schröder et al. (2008), and sources given by these authors, and some additional references cited below Location Age (Ga) Units Evaporite evidence Thickness Notes References 43 Canada 0.7–1.2 Gy Pope and Grotzinger (2003) Mauritania ca. 1.1 Multiple gypsum beds up to 30 m thick ca. 50 m (?) Marine 42 Minto Inlet and Kilian formations, Shaler Group Oued Tarioufet Formation, Atar Group, Gouamir and Tenoumer formations, El Mreiti Group, Taoudeni Basin SGa (?), marine Kah et al. (2012) 41 Mauritania, Algeria ca. 1.2 Char Group, Mauritania, Douik Group, Algeria ca. 50 m Evans (2006) 40 Canada 1.2 Society Cliffs Formation, Victor Bay Group, Borden Basin SGb, marine, possibly correlate with evaporites in Atar and El Mreiti Groups (above) SGc, restricted marine 39 USA 1.15–1.3 Upper Marble, Grenville Series Metamorphosed evaporites Whelan et al. (1990) 38 USA, Canada 1.46 Waterton, Altyn, Prichard, and Wallace formations, Belt Supergroup SGd, marine, two evaporite horizons Evans (2006) 37 Australia ca. 1.5 70 m SGe, marine, several evaporite horizons Evans (2006) 36e Australia 1.61 Discovery Formation, Edmund Group, Bangemall Supergroup, Bangemall basin Balbirini Formation, McArthurMt Isa basins ? SGf, alkaline lake suggested by shortite Walker et al. (1977), Evans (2006) 36d Australia 1.635 Lynott Formation, McArthur-Mt Isa basins ca. 300 m SGf, marine sabkha Walker et al. (1977), Evans (2006) 36c Australia 1.645 ca. 200 m SGf Walker et al. (1977), Evans (2006) 36b Australia 1.66 Myrtle, Emmerugga, and other formations, McArthur-Mt Isa basins Mallapunyah, Paradise Creek, Esperanza, Staveley formations McArthur-Mt Isa basins >10 m SGf, marine sebkha Walker et al. (1977), Evans (2006) Ca-Evp, Ca-ps-Gy or Ar, ps-Ha, sc-breccias, chicken-wire texture ps-Ha Gy, ps-Ha, sc-breccias An, as lenses and beds ps-Ha, pseudomorphs after sulfates, ps-Sh, cauliflower cherts ps-Gy, ps-Ha, cauliflower cherts ps-Gy, ps-Ha ps-Gy, ps-Ha, botryoidal quartz nodules after anhydrite, massive replacement by gypsum Kah et al. (2001), Evans (2006) 537 (Continued) Geochemistry of Evaporites and Evolution of Seawater ps-Evp, ps-Gy, ps-An, ps-Ha, length-slow chalcedony, chicken-wire textures, scapolite ps-Gy, or ps-An Multiple beds a few cm’s to meter’s thick, >100 m Beds (or lenses) >40 m thick 100 m 538 (Continued) Location Age (Ga) Units Evaporite evidence Thickness Notes References 36a Australia 1.74 (1.54–1.74) Corella Formation, McArthur-Mt Isa basins ca. 500 m SGg, alkaline lake suggested by shortite Walker et al. (1977), Muir (1987), Evans (2006) 35 India >1.7 ? Canada 1.8 Marine, associated with lava flows Marine to non-marine, halite > > gypsum Pope and Grotzinger (2003) 34 Vempalle Formation, Papaghni Group Cowles Lake Formation 33 Canada 1.8 Brown Sound Formation ca. 300 m Marine to non-marine, halite > > gypsum Pope and Grotzinger (2003) 32 Russia 1.8–1.9 ? Associated with barite, sabkha Pope and Grotzinger (2003) 31 Canada 1.82–1.91 Tavani Formation, Hurwitz Group ? Coastal pans, marine-to-nonmarine Aspler and Chiarenzelli (2002) 30 Canada ca. 1.87 Stark and Hearne formations, Great Slave Lake Supergroup 200–600 m, reconstructed thickness of evaporites ¼ ca. 100 m SGh, marine to non-marine, halite > > gypsum Pope and Grotzinger (2003), Evans (2006) 29 Canada 1.8–2.0 ca. 150 m Marine Pope and Grotzinger (2003) 28 Canada ca. 1.95 Traces of evaporites dispersed within carbonates S Africa Russia 2.06 ca. 2.09 SGi (?), marine, lagoon on inner shelf, passive margin Lacustrine environment in rift SG j, passive margin, playa lake, marine sabkha, intertidal flats Evans (2006) 27 26 Kasegalik and Mc-Leary formations, Belcher Group Rocknest Formation, Coronation Supergroup, Slave craton Dewaras Group Tulomozero Formation, Upper Jatulian Group ps-Sh, ps-Gy (?), quartz-replacing anhydrite nodules ps-Ha, ps-Gy ps-Ha, ps-Gy, sc-breccias ps-Ha, ps-Gy, sc-breccias ps-Gy, ps-An Q-ps-Gy, Dol-ps-Gy, halite moulds ps-Ha, silicified hopper casts, and pagoda halite, ps-Gy, sc-megabreccia ps-Gy, ps-Ha Dol-ps-Gy, ps-An, ps-Ha Gy (?) Ca-ps-Gy, Dol-ps-Gy, Si-ps-Gy, An (relics), pseudomorphs after anhydrite and gypsum crystals and nodules, ps-Ha, sc-breccias, enterolithic and chicken wire structures >200 m Multiple units >20 m, within ca. 500 m of total thickness Pope and Grotzinger (2003) Pope and Grotzinger (2003) Melezhik et al. (2005), Brasier et al. (2011), Reuschel et al. (2012) Geochemistry of Evaporites and Evolution of Seawater Table 11 Russia ca. 2.1 Fedorovka (Fedorov) Formation (Aldan Shield) Norah Formation, Deweras Group Francevillian C Formation, Francevillian Group Lower part of the Nash Fork Formation, Snowy Pass Supergroup Laparre Formation, Peribonca Group, Otish Supergroup 24 Zimbabwe ca. 2.15 23 Gabon ca. 2.0–2.2 22 USA ca. 2.15 21 Canada ca. 2.15 20 S Africa ca. 2.15 (2.10–2.20) Lucknow Formation, Olifantshoek Group and Transvaal Supergroup 19 Australia ca. 2.15 (2.2)? Bubble Well Member, Juderina Formation, Yerrida Group 18 S Africa 2.2 Pretoria Group 17 Australia ca. 2.2 Bartle Member, Killara Formation, Yerrida Group 16 USA ca. 2.22–2.3 Kona Dolomite, Chocolay Group 15 Canada ca. 2.22–2.3 Gordon Lake Formation, Huronian Supergroup 14 S Africa 2.52–2.56 Campbellrand-Malmani carbonate platform, Transvaal Supergroup An, as layers and veins ? Passive margin An, as layers ? Intracratonic rift basin Ca-ps-An, Ca-ps-Gy Molds after anhydrite nodules and gypsum crystals ? Préat et al. (2011) ? Marine, supratidal-sabkha environment Passive margin Dol-ps-Gy, Dol-ps-An (after crystals and nodules) Q-ps-Gy, Q-ps-An, molds after gypsum and anhydrite Si-Evp, Q-ps-Gy, Q-ps-An ps-Mir ? Passive margin Bekker et al. (2006) ? Marine, passive margin Bekker et al. (2006), Schröder et al. (2008) ca. 100 m SGk, Marine or marginal marine, associated with volcanics Sodic lake deposits in a playa setting Playa lake (alkaline?) El-Tabakh et al. (1999a) Si-ps-An, Si-ps-Gy, Kao-ps-Gy or An, An (relics), ps-Sh?, ps-Tro? Si-ps-Gy, Si-ps-An, ps-Ha (moulds), sc-breccias Ba (as beds), silicified and pristine anhydrite and gypsum nodules and layers, Si-tr-An; beds of anhydrite nodules Si-ps-Ha, Ca-ps-Gy (?), sc-breccias <2 m ? 30–1000 m Multiple horizons in >300 m >500 m Zharkov (2005), Bekker et al. (2006) Bekker et al. (2006) Bekker et al. (2006) Pope and Grotzinger (2003) Pirajno and Gray (2002) SGl, marine, associated with volcanics, intracratonic basin, open to passive margin, correlated with Gordon Lake Formation SGl, marine? passive margin, supratidal and sabkha zone, correlated with Chocolay Group Bekker et al. (2006) Marine Sumner and Grotzinger (2004), Gandin et al. (2005) Cameron (1983), Bekker et al. (2006) (Continued) Geochemistry of Evaporites and Evolution of Seawater 25 539 540 (Continued) Location Age (Ga) Units Evaporite evidence Thickness Notes References 13 S Africa ca. 2.58 ps-Ha (moulds), Ca-ps-Ar ? Supratidal flat or sabkha Eriksson et al. (2005) 12 Australia 2.4–2.8 (2.6) Black Reef and Oaktree formations, Transvaal Supergroup Carawine Formation (Carawine Dolomite), Hamersley Group <20 m Marine, considered as the earliest undoubtful selenite deposits Simonson et al. (1993), Sumner and Grotzinger (2000) 11 Australia 2.6–2.7 Black Flag Beds 40 m Canada 2.7 Steeprock Group 9 Zimbabwe 2.7 Cheshire Formation, Belingwe Greenstone Belt 8 Australia 2.7 ps-Ha ca. 320 m 7 6 S Africa Australia 2.8 2.97–3.19 Ga Lacustrine Continental margin Australia 3.3–3.5 4 S Africa 3.4 Witkop Formation, Nondweni Greenstone Belt ps-Nat ps-Evp, Si-ps-Nah ps-Ha, ps-Gy Ba-ps-Gy ? Several horizons 5 Tumbiana Formation, Fortescue Group Ventersdorp Supergroup Farrel Quartzite, George Creek Group Rocklea Dome Tidal flats (?) associated with volcanics Marine, carbonate platform with stromatolites, supposedly non-evaporite carbonate (Ar) deposits Marine, carbonate platform with stromatolites, supposedly non-evaporite carbonate (Ar) deposits Lacustrine or marine Pope and Grotzinger (2003) 10 Dol-ps-Gy, Si-ps-Gy, or Dol-ps-Ar (?), ps-Ha Ank-ps-Gy, Ank-ps-An (?) Ca-ps-Ar, or Ca-ps-Gy (?), Si-tr-Gy, Si-tr-An (?) Ca-ps-Ar, or Ca-ps-Gy (?) 3 S Africa 3.4 Buck Reef Chert, Kromberg Formation, Onverwacht Group, Barberton Greenstone Belt Ba, Ba-ps-Gy (?), ps-Nah, Si-Evp, molds, silicified sc-breccias Pseudomorphs dispersed within carbonates 10 m <20 m <1.5 m 5–40 m Grotzinger (1989), Sumner and Grotzinger (2000), Hardie (2003) Grotzinger (1989), Hardie (2003) Buick (1992), Awramik and Buchheim (2009) Pope and Grotzinger (2003) Sugitani et al. (2003, 2007) Boulter and Glover (1986) Associated with volcanics Wilson and Versfeld (1994), Hofmann and Wilson (2007) Byerly and Palmer (1991), Lowe and Worrell (1999), Lowe and Byerly (2007) Geochemistry of Evaporites and Evolution of Seawater Table 11 2 Australia ca. 3.4 (3.35– 3.43), or (3.346–3.459) Strelley Pool Chert (Strelley Pool Formation), Kelly Group 1 Australia ca. 3.49 (3.447– 3.496) Dresser Formation (North Pole Chert), Warrawoona Group Ba, Ba-ps-Gy, Si-ps-Nah (?), Si-ps-Ba (?), ps-Ha, Si-ps-Ar (?) Ba (as beds), Ba-ps-Gy, Si-ps-Gy (crystal rosettes), ps-Ha Multiple beds, <10 to 25 m Associated with volcanics, interpreted as marine Lindsay et al. (2005), Warren (2006), Allwood et al. (2007, 2009), Van Kranendonk (2007) Multiple beds, <10 to 25 m Associated with volcanics, considered as non-marine Lambert et al. (1978), Lowe (1983), Buick and Dunlop (1990), Shen and Buick (2004), Runnegar et al. (2001), Allwood et al. (2007), Lowe (1983), Grotzinger (1989), Warren (2006), Van Kranendonk (2007) Geochemistry of Evaporites and Evolution of Seawater Saline giants (SG), with volume ) 1000 km3, are distinguished and shortly described (in footnotes) after Evans (2006), except of the Mesoproterozoic Taudeni Basin (see Table 9). An, anhydrite; Ank-ps-Gy, ankerite pseudomorphs after gypsum; Ba, barite; Ba-ps-Gy, barite pseudomorphs after gypsum; Ca-Evp, calcitized evaporites; Ca-ps-Ar, carbonate pseudomorphs after aragonite; Ca-ps-An, carbonate pseudomorphs after anhydrite; Ca-ps-Gy, carbonate pseudomorphs after gypsum; Ca-ps-Gy or -Ar, carbonate pseudomorphs after gypsum or aragonite; Dol-ps-An, dolomite pseudomorphs after anhydrite; Dol-ps-Gy, dolomite pseudomorphs after gypsum; Gy, gypsum; Kao-ps-Gy or An, kaolinite pseudomorphs after gypsum or anhydrite; ps-An, pseudomorphs after anhydrite; ps-Evp, pseudomorphs after evaporites; ps-Gy, pseudomorphs after gypsum; ps-Mir, pseudomorphs after mirabilite; ps-Nah, pseudomorphs after nahcolite; ps-Nat, pseudomorphs after natron; ps-Ha, pseudomorphs after halite; ps-Sh, pseudomorphs after shortite; ps-Tro, pseudomorphs after trona; Q-ps-An, quartz pseudomorphs after anhydrite; Q-ps-Gy, quartz pseudomorphs after gypsum; sc-breccias, solution collapse breccias; Si–Evp, silicified evaporites; Si-ps-An, silicified pseudomorphs after anhydrite; Si-ps-Ar, silicified pseudomorphs after aragonite; Si-ps-Gy, silicified pseudomorphs after gypsum; Si-ps-Nah, silicified pseudomorphs after nahcolite; Si-tr-An, quartz filled traces after anhydrite; Si-tr-Gy, quartz filled traces after gypsum. Lack of or highly controversial data are marked by “?”. Notes to saline giants (SG): a Traces of calcitized evaporites within a few tens of m thick stratigraphic interval traced at the distance >1500 km. b The basin area 800 ( 200 km. c The basin area ca. 140 000 km2. d Two intervals of vanished or metamorphosed evaporites on an area of 300 ( 200 km; restoration of tectonic shortening suggests a basin ca. 100 000 km2 in size. e The basin area ca. 40 000 km2. f The basin area at least 5000 km2. g Scapolite-albite-tourmaline association, total ca. 500 m thickness across an area of 200 ( 20 km. h The basin area ca. 300 000 km2. i Traces of evaporites within ca half of the thickness of a carbonate facies, which spans an area of about 250 ( 50 km, and in a lagoon zone about 200 km wide. j Relict evaporitic textures within ca. 500 m of the dolomitic section. k The basin area 100 ( 100 km. l Pseudomorphs in ca. 100 m of the section over a ca. 400 ( 100 km in Chocolay Group (USA); a 40 m thick basal part of the Gordon Lake Formation (Canada) contains anhydrite nodules and breccias, interpreted as a sabkha environment. 541 542 Geochemistry of Evaporites and Evolution of Seawater initial ocean began to decrease due to the accumulation of evaporites on continental shelves not earlier than 2.5 Ga (Knauth, 2005). Probably, the oldest known record of marine evaporites is represented by silicified pseudomorphs after beds of some unknown bottom-grown evaporite crystals occurring in 3.43 billion-year-old Strelley Pool Chert, in Pilbara Craton, Australia (Table 11 and Figure 22). These beds show traces of synsedimentary dissolution, and of syntaxial growth over dissolution surfaces, and are associated with stromatolitic structures and solution–collapse breccias. Lowe (1983) suggested gypsum or aragonite, and Lowe and Tice (2004) – nahcolite, as the original mineralogy. Lindsay et al. (2005) recognized postevaporite ‘chicken-wire’ textures formed by quartz aggregates and described 30-cm-long quartz pseudomorphs after supposed barite and also aragonite crystals. However, Allwood et al. (2007), who recently restudied these outcrops, did not specify what evaporite mineral crystallized in the early Archean environment, although in later work, they agreed that it was “probably originally aragonite” (Allwood et al., 2009, p. 9548), as it was also suggested by van Kranendonk (2006). Allwood et al. interpreted the environment as “an isolated, partially restricted, peritidal marine carbonate platform, or reef, where there is virtually no trace of hydrothermal or terrigenous clastic input” (Allwood et al., 2007, p. 198). Marine environment of these deposits was earlier proved by geochemical studies (van Kranendonk et al., 2003). The traces of evaporites (pseudomorphs) from marine deposits are known from several formations dated c.2.8– 2.4 Ga (Table 11; Pope and Grotzinger, 2003). However, the oldest (! 2250 Ma) saline giant, recognized by the presence of copious pseudomorphs after gypsum and anhydrite within 100 m thick interval of the section, occurs in Chocolay– Gordon Lake succession, on Superior craton, Canada and United States, and is estimated as 4500 km3 of vanished evaporite salts (Table 9; Evans, 2006). The Gordon Lake Formation in Ontario presumably contains the earliest preserved Ca sulfate deposits as thin beds of anhydrite nodules within laminated mudstone (Cameron, 1983; Huston and Logan, 2004). Most of the pre-Ediacaran saline giants are known only from the accumulation of pseudomorphs within thick stratigraphic intervals and associated collapse breccias. The earliest wellpreserved sequence of extensive bedded evaporites, including gypsum, occurs in the latest Paleoproterozoic to early Mesoproterozoic (! 1.6 Ga) rocks of the McArthur Basin of Australia (Tables 9 and 11; Walker et al., 1977). The younger mentioned Centralian Superbasin in Australia (800–830 Ma), about 140 000 km2 in size, contains the most extensive and relatively well-preserved gypsum, anhydrite, and halite beds with typical thickness 800 m (Table 9; Evans, 2006, with references). The first recorded bedded Ca sulfate deposits occur in Mesoproterozoic, proving that sulfate concentration in water has been high enough for more abundant gypsum precipitation at least since that time (Kah et al., 2004). The mentioned first recorded bedded Ca sulfate deposits in the Mesoproterozoic McArthur Basin (1.6 Ga) also marks the limit of the existence of the hypothetic early soda ocean according to the other concept of ocean chemistry evolution (Kempe and Kaźmierczak, 1994). The assumed slow rise in sulfate concentration in the Archean and Mesoproterozoic is consistent with C isotope record from that time and with a presumed increase in oxygenation recorded in the biosphere (Kah et al., 2004). The predicted concentration of sulfate in Mesoproterozoic was presumably as low as 2.7–4.5 mM (Kah et al., 2004). At that time, the availability of Ca was apparently limited by excess precipitation of calcium carbonate resulting from elevated carbonate saturation, and therefore, a great DE would be required to attain gypsum saturation (Kah et al., 2004, with references). In the presence of high amounts of Cl and Na in the Mesoproterozoic seawater, halite would precipitate before gypsum during evaporation, which is consistent with the scarce geologic record. It seems likely that very low concentration of sulfate ions before Mesoproterozoic (Reuschel et al., 2012), even though the concentration of calcium ions could be high (Rouchon et al., 2009), resulted in lack or very sparse deposition of Ca sulfates during the early time of Earth history (Eriksson et al., 2005; Foriel et al., 2004; Kah et al., 2004; Zentmyer et al., 2011). It is estimated that the concentration of sulfate ions was less than 200 mM in the Archean, and it rose to over 1 mM in the early Paleoproterozoic and to more than 2.5 mM in the mid-Paleoproterozoic (Reuschel et al., 2012, with references), Grotzinger (1989), Grotzinger and Kasting (1993), Grotzinger and Knoll (1995), Sumner and Grotzinger (2000), and Pope and Grotzinger (2003) questioned the occurrence of marine bottom-grown gypsum crystals and massive evaporite deposition in the Archean and Paleoproterozoic and challenged the earlier interpretation that the chemistry of the ocean was likely the same as today since the (late) Archean (Hardie, 2003; Walker, 1983). They assumed that bicarbonate ion concentration exceeded twice that of calcium in Precambrian seawater (Grotzinger, 1989, see Rouchon et al., 2009, for more up-todate information on contents of calcium and carbonate– bicarbonate ions in the Archean seawater). This would cause the Ca ion to become exhausted during evaporite concentration by calcite/aragonite precipitation well before the stage of gypsum precipitation was achieved. Therefore, the precipitation of gypsum was bypassed during early salinity rise, and halite precipitated directly after Ca carbonates in all the early Precambrian marine evaporite successions (Eriksson et al., 2005; Pope and Grotzinger, 2003). According to Grotzinger (1990), the rare occurrences of pseudomorphs after gypsum in earliest Precambrian were restricted to deltaic settings, where locally higher concentration of calcium could appear. Grotzinger and coauthors assumed that NaCl concentration was high, in agreement with the presence of halite pseudomorphs. The high chloride concentration, together with much other evidence, suggests that the seawater had relatively low pH at that time (e.g., Foriel et al., 2004; Holland and Kasting, 1992; Pinti, 2006; Rouchon et al., 2009; Sugisaki et al., 1995). This would explain why seawater was unable to precipitate Na carbonates (Grotzinger and Kasting, 1993), otherwise expected to form in the hypothetic soda ocean (Kempe and Degens, 1985). We would like to point out that the use of mineralogy as the only evidence or the lack of evidence of the Usiglio sequence might lead to significant misinterpretation in the case of limited data. Modern halite deposits can start an evaporite sequence without or only with minor crystallization of earlier gypsum, as it is proved by both modeling of the marginal marine evaporite basins (Sanford and Wood, 1991) and a similar record from some modern environments – for example, halite pans on supratidal flats or the Taxada halite from the MacLeod basin (Logan, 1987). Indeed, evaporite sequences can be modified by a replacement processes, which include 543 Era Period Quaternary Neogene Phanerozoic Paleozoic Mesoz. C. Neoproterozoic Ter. Eon Geon Geochemistry of Evaporites and Evolution of Seawater Paleogene 60 ± 0.5 Ma Cretaceous Jurassic Triassic 251 ± 0.1 Ma Permian Carboniferous Devonian Silurian Ordovician Cambrian 542 ± 1.0 Ma Ediacaran Mesoproterozoic Proterozoic 850 Tonian 1200 43 42 39 40 41 Ectasian 1400 Calymmian 1600 Statherian Paleoproterozoic Centralian superbasin, Australia, the largest preserved pre-Ediacaran saline giant, ~140 000 km3 800–830 Ma Stenian Common, well preserved evaporites ~635 Ma Cryogenian 1000 1800 Orosirian 2050 Rhyacian 2300 38 37 36d 36e 36b 36c 35 36a Saline giant with 33 34 abundant traces 32 30 31 of Ca-sulfates, 29 2.1 Ga Tulomozero Formation, 28 27 Upper Jatulian Group, Russia 2500 Neoarchean 26 20 21 22 23 24 17 18 19 15 16 First undisputable vanished saline giant with Ca-sulfates and halite, Gordon Lake Formation, Huronian Supergroup, Canada and Kona Dolomite, Chockolay Group, USA 2.3 Ga 14 11 12 13 8 2800 First common Ca-sulfate evaporites 25 Siderian 9 10 7 ca 3.06 Ga Archean Gulf of Mexico evaporites, the largest saline giant on Earth, ~2 400 000 km3 Mesoarchean 2.7 Ga 2.06 Ga Great Oxidation Event 2.35 Ga Pseudomorphs after halite, Tumbiana Formation (lacustrine or marine), Fortescue Group, Australia Silicified pseudomorphs after nahcolite? Farrel Quartzite, George Creek Group, Australia 6 3200 Paleoarchean 3600 Eoarchean ~4000 Hadean Barite pseudomorphs after gypsum? Halite molds 3.4 Ga Witkop Formation, 3.4 Ga in chert, Nondweni Rocklea Dome, Greenstone Australia Belt, 2 3 4 5 S. Africa Silicified 1 pseudomorphs 3.4 Ga after nahcolite? ca 3.49 Ga Buck Reef Chert, Kromberg Formation, Barite and/or Barberton Greenstone barite pseudomorphs Belt, S. Africa after gypsum? Dresser Formation, ca 3.4 Ga Warrawoona Group, Australia Silicified pseudomorphs after evaporite mineral 4.03 Ga The oldest rocks; (aragonite? gypsum? Acasta gneisses, nahcolite?) Canada Strelley Pool Chert, Kelly Group, Australia 4.4 Ga The oldest minerals; detrital zircons from Jack Hills metaconglomerate, Australia (early hydrosphere) ~4600 1 15 Precambrian evaporites listed and numbered in Table 11 Saline giants with volume of evaporites ³1000 km3 Figure 22 Precambrian (pre-Neoproterozoic) record of evaporite deposits, after data by Bekker et al., 2006; Evans, 2006; Bekker and Holland, 2012, and other references in Table 11. Age of the Great Oxidation Event after Bekker and Holland, 2012; time scale after Ogg et al., 2008. 544 Geochemistry of Evaporites and Evolution of Seawater the early replacement of less soluble gypsum by more soluble halite taking place in halite-producing environments having supersaturated brine warmed to 35–50 " C, which is typical of the heliothermal effect (Hovorka, 1992; Schreiber and Walker, 1992; Schröder et al., 2003). 2. 3. 4. 5. Prevalence of symmetrical (wave) ripples Lack of unidirectional current structures Paucity of dolomite An evaporite succession that goes from carbonate directly to halite (no sulfates were found) His interpretations of this evidence are the following: 9.17.23.3 Nonmarine Evaporites in Precambrian According to Frimmel and Jiang (2001), most of the known but scarce records of metamorphosed Proterozoic evaporites represent nonmarine playa lake environments in rift grabens. They are recognized mainly because of mineralogical and geochemical data such as their low 11B/10B ‘nonmarine’ ratios (Byerly and Palmer, 1991). These values are recorded in a number of Proterozoic borate deposits, including the (2.1 Ga) of the Liaohe Group in Liaoning, China (Jiang et al., 1997; Peng and Palmer, 1995); the )1.7 Ga Thackaringa Group in New South Wales, Australia (Slack et al., 1989); and the Neoproterozoic Duruchaus Formation in the Damara Belt in Namibia (Porada and Behr, 1988). Additionally, there are data from some other occurrences (Frimmel and Jiang, 2001; Grew et al., 2011). On the other hand, high 11B/10B ratios in tourmalines associated with vanished silicified Archean (! 3.5 Ga) evaporite deposits in Barberton greenstone belt (Table 11) suggest derivation of the boron from marine evaporites (Byerly and Palmer, 1991). Evaporites cannot help directly in recognition of the salinity and chemistry of the Precambrian oceans because it is difficult to recognize truly marine evaporites in Precambrian, except for the saline giants described earlier. However, even the scale of a large basin, there is not universal criterion because the brine in particular subbasins can evolve in its own pathway, as proved by Garcı́a-Veigas et al. (1995) and Ayora et al. (2001). Only carefully integrated, multimethodological geochemical studies can lead to the recognition of the marine signal in halite fluid inclusions (Garcı́a-Veigas et al., 2009, 2011). The separation of marine from lacustrine deposits in Precambrian settings, even by using the criteria listed by Kelts (1988) and Eriksson et al. (2004), is more difficult than in the Phanerozoic and requires complex analysis (Awramik and Buchheim, 2009; Brasier, 2011). Southgate et al. (1989) suggested the following criteria helpful in the recognition of marine and lacustrine Precambrian evaporites: 1. Arrangement of facies 2. The assemblage of evaporite minerals or their pseudomorphs 3. Any fossils that may have been present 4. Geochemical evidence Buick (1992) argued that the Late Archean evaporite sequences of the Tumbiana Formation in Western Australia (Table 11), which pass from Ca carbonates to halite, without intercalated or present traces of gypsum, are different than the Usiglio sequence and that this suggests a lacustrine environment. Buick (1992) further supported lacustrine environments for these deposits based on several lines of evidence: 1. Interfingering relationships found between the nondetrital sediments and the terrestrial basalts and between the alluvial fan and fluvial sediments 1. The interfingering relationships are unlikely to occur in marine transgressions. 2. Unidirectional sedimentary structures such as asymmetrical ripples and herringbone cross-bedding would suggest tidal activity, but they are absent. 3. Dolomite is very common in ancient marginal marine carbonates and is usually indicative of saline waters with high Mg/Ca ratios. The rarity of dolomite in the Tumbiana Formation suggests nonmarine conditions consistent with a lacustrine depositional environment. 4. Buick (1992) did not find gypsum, its pseudomorphs, or other sulfate evaporite evidence, but sulfates are known from the Archean marine deposits elsewhere (Zharkov, 2005). Its apparent absence from the Tumbiana Formation, !2.7 Ga, was conspicuous for Buick (1992) and was used as strong evidence for the nonmarine origin of this deposit. By contrast, this and similar formations were treated by Reddy and Evans (2009) as marine in origin. The halite pseudomorphs present within calcilutites, without any traces after dissolved gypsum in the underlying sequences – as expected in the Usiglio sequence, were used as evidence of the anomalous chemistry of the early ocean, following earlier interpretation by Grotzinger (1989), Pope and Grotzinger (2003). The origin of this well-exposed formation – marine or nonmarine, tideless sea, or lake – is the subject of continuing debate (Awramik and Buchheim, 2009; Bolhar and van Kranendonk, 2007; Sakurai et al., 2005). This is a good example of the weakness of any interpretations of the Precambrian seawater chemistry based on a nonfossiliferous sedimentary record, lacking of clear diagnostic features of marine environment. 9.17.23.4 Pseudomorphs after Evaporite Minerals in Precambrian Pseudomorphs after various salt minerals are common in the Precambrian (Table 11). The most common appear to be of gypsum (in Neoarchean; Gandin et al., 2005) and halite (Pope and Grotzinger, 2003). Microcline pseudomorphs (up to 10 cm long) after shortite and gaylussite were recognized in the Callanna Group of the Willouran Ranges, Australia, late Proterozoic 1.4–0.8 Gy. These are interpreted as evaporite playa lake deposits (Rowlands et al. 1980 in Muir, 1987). Silicified pseudomorphs (up to 40 cm long) after bottom-grown evaporite crystals (nahcolite) occur in >! 2.97 Ga Farrel Quartzite, in George Creek Group, Pilbara Craton, Australia (Sugitani et al., 2003, 2007). Possible pseudomorphs after shortite occur in Corella Formation (1.74–1.54 Gy) in Australia (Muir, 1987). The Precambrian pseudomorphs interpreted as postevaporite (after gypsum, nahcolite, etc.) in origin have been commonly treated as evidence of the chemistry of the ancient oceans. Such opinions seem to be invalid until two facts are proved: first that these are true pseudomorphs after the given evaporite mineral, Geochemistry of Evaporites and Evolution of Seawater and second, that the evaporite crystal is marine in origin. In the case of Precambrian deposits, it is not easy, if ever possible. Probably, the oldest well-recognized record of vanished evaporites on Earth is silicified pseudomorphs of the bottom-grown crystals, up to 40 cm long, interpreted as supposed nahcolite (NaHCO3) from the Kromberg Formation (3.416–3.334 Ga), Barberton Greenstone Belt, South Africa (Table 11; Lowe and Worrell, 1999). The authors were not sure about that interpretation based on the measurements of interfacial angles with “an error of from +2" , where the faces were well preserved, to as much as +10" between poorly preserved faces” (Lowe and Worrell, 1999, p. 180). They concluded that the silicified pseudomorphs “most closely resemble nahcolite” (Lowe and Worrell, 1999, p. 181). Nahcolite is a typical evaporite mineral of the soda lakes (Batalin et al., 1973; Warren, 2010). The other comparable example of equally old (! 3.4 Ga) vanished evaporites are barite and quartz pseudomorphs after gypsum, as well as negative crystals of that mineral, forming stellate aggregates and single euhedral forms in the Witkop Formation, the Nondweni Group, Kaapvaal Craton, South Africa, well documented by measurements of interfacial angles of the crystals (Wilson and Versfeld, 1994). The primary mineralogy of many pseudomorphs is commonly identified by a superficial comparison of crystal habit, which can lead to serious mistakes. Many lens-shaped pseudomorphs interpreted as postgypsum forms could mimic many other salt minerals: ikaite, gaylussite (Warren, 2006), or glauberite. In particular, glauberite is remarkably similar in morphology to gypsum (Salvany et al., 2007). Trona and gypsum are both monoclinic and form radial sprays (Smoot and Lowenstein, 1991). The following groups of minerals can form similar pseudomorphs: ikaite–gypsum–gaylussite-glauberite, barite– siderite–gypsum, aragonite–gypsum, pyrite–halite-sylvite, and anhydrite–gypsum (Warren, 2006, supplemented). Hydrohalite (NaCl • 2H2O) crystallizing in temperatures below 0 " C shows hexagonal shape that can be attributed to many minerals including gypsum (Roberts et al., 1997). The Precambrian and Permian marine deposits are famous for their seafloor crystal crusts showing grasslike structures (radial bundles) and interpreted as calcite pseudomorphs after primary bottom-grown aragonite (e.g., Sumner, 2002). Many of these crusts were formerly interpreted as calcitized grasslike gypsum deposits, based in part on the large crystal sizes (Cassedanne, 1984; references in Sumner and Grotzinger, 2000; Riding, 2008). Some occurrences were interpreted as pseudomorphs after trona (Jackson et al. 1987 cited by Winefield, 2000). Recently, the majority of such crusts have been interpreted and/or reinterpreted as aragonite, based, among other reasons, on the high amount of strontium present (up to 4169 ppm; Grotzinger, 1989; Peryt et al., 1990; Sumner and Grotzinger, 2000). Aragonite commonly contains a great deal of strontium, whereas gypsum and calcite typically incorporate less strontium because the strontium partition coefficient for aragonite is much larger (1.13) than strontium partition coefficients for calcite and gypsum (<0.2). These crystal pseudomorphs reach centimeter to decimeter in length (some upright crystals attain 1.60 m in length, Sumner and Grotzinger, 2000), and show elongated rodlike habit (‘rays in 2D view’) with hexagonal cross sections. The aragonite mineralogy also is suggested by squared-off growth-off zones and 545 squared terminal crystal apices (Peryt et al., 1990). Another revisitation of these crystal pseudomorphs should be made utilizing the results given by Riccioni et al. (1996) in which aragonite crystals were examined closely and the hexagonal cross sections (in that deposit) were found to be the result of penetration twinning (aragonite is orthogonal) and commonly show diagnostic indentations on some faces. One of the seafloor crystal crusts, from the 2.6 Ga Carawine Dolomite in Australia (Table 11; Simonson et al., 1993), represents carbonate pseudomorphs after crystals, considered as the earliest unquestionable gypsum precipitates (Eriksson et al., 2005). Hardie (2003) and Gandin et al. (2005) questioned at least one of the ‘aragonite’ interpretations (from Neoarchaean Kogelbeen and Gamohaan formations of the Campbellrand Subgroup, South Africa, !2.5 Ga), suggesting that the pseudomorphs can represent selenite crystals creating large domal structures similar to those known from the Messinian of the Mediterranean. Hardie further noted that careful measurements of interfacial angles are needed to support these interpretations. Indeed, at least three minerals with elongated habit can show very similar hexagonal cross sections: aragonite, gypsum, and nahcolite. The distinction requires very careful measurements of interfacial angles, which however not always bring the conclusive results. Sumner (2004) stated that “measurements of interfacial angles are consistent with either an aragonite or gypsum precursor due to the sensitivity of results to small errors in cross section orientation.” “Errors were estimated to be 5–10" , which are too high to aid in primary mineral identification” (Sumner, 2004). The additional misinterpretation can result from vicinal character of natural crystal faces or unnoticed compactional or tectonic deformation, particularly acting during replacement process (Hovorka, 1992). The proper identification of primary minerals forming pseudomorphs requires the statistical measurement of interfacial angles of the pseudomorphs and their comparison with the ideal crystal form (Smoot and Lowenstein, 1991). This was rarely made in the case of Precambrian evaporite pseudomorphs. The rare exceptions concern pseudomorphs after gypsum (Walker et al., 1977; Dunlop, cited in Lambert et al., 1978; Wilson and Versfeld, 1994), after the aragonite (Winefield, 2000), and after nahcolite (Lowe and Worrell, 1999; Sugitani et al., 2003). Nevertheless, even such measurements in case of barite crystals being supposed pseudomorphs after gypsum gave conflicting results (Buick and Dunlop, 1990; Lambert et al., 1978; Runnegar, 2001; Runnegar et al., 2001). These include the Archean (3.47 Ma) North Pole Chert, Warrawoona Group, Australia (Figure 23; Buick and Dunlop, 1990; Lambert et al., 1978; Runnegar, 2001; Runnegar et al., 2001; Shen and Buick, 2004; Shen et al., 2001, 2006; see comments by Buick, 2008; Warren, 1997, 2006, pp. 106, 559). Barite is, however, apparently also primary in these Archean deposits (Nijman et al., 1998). Runnegar (2001) and Runnegar et al. (2001) proved it by measurements of the interfacial angles in some crystals. These authors used X-ray computer tomography (CT) and questioned the occurrence of primary gypsum in these beds (the documentation of these studies was not published, however). Shen et al. (2009) commented, “X-ray CT only images density contrasts in the sample, so this technique cannot reveal original crystal morphology in partially silificified sediments where the 546 Geochemistry of Evaporites and Evolution of Seawater postrift passive margins, and (4) continental collision zones and foreland basins. Hudec and Jackson (2007, their Figures 1–4) showed the distribution of these evaporite basins on Earth. The evaporites are important in Earth history and some of their selected significant features are as follows. 9.17.24.1 Figure 23 Laminated chert draping the bottom-grown barite crystals interpreted as pseudomorphs after gypsum by Buick and Dunlop (1990) and as primary barite by Runnegar et al. (2001); North Pole Chert (3.47 Ga), Warrawoona Group, Australia, scale in centimeters. Photo courtesy Roger Buick. microquartz–barite boundary is within the original crystal boundary along barite cleavage planes. By contrast, the universal-stage petrographic technique used in previous studies (Buick and Dunlop, 1990; Lambert et al., 1978) differentiates between the inclusion-rich microquartz of the surrounding silicified sediment and the inclusion-poor microquartz of the silicified sulfate crystals, thus yielding accurate interfacial angle measurements” (Shen et al., 2009, p. 384). Buick (2008), similar to earlier authors (Lambert, 1978; Shen et al., 2006), believed that the necessary sulfate ions were produced from H2S by anaerobic photosynthesizers, being a part of ancient sulfuretum ecosystem, and using this gas as their electron donor. This view, however, is highly controversial – inorganic origin of sulfates (by photochemical and other reactions in atmosphere and surface of the hydrosphere from volcanic emanations of hydrogen sulfide and sulfur dioxide, see, e.g., Grotzinger and Kasting, 1993; Holland, 2002; Huston and Logan, 2004; Lambert, 1978; Walker, 1983) is equally or more probable (according to Johnston, 2011, with references). These probably are the oldest recorded bottomgrown sulfate crystals and remain one of the most important and intensively studied objects in pursuit of an understanding the Archean sulfur cycle and chemistry of the Archean atmosphere and ocean, despite the fact that it is unclear if they are lacustrine or ‘modified’ marine deposits (Grotzinger, 1989). Irrespective of the primary mineralogy of these precipitates (barite or gypsum), they indicate that the Paleoarchean seas or lakes “were at least locally sulfate bearing” (Golding et al., 2010, p. 42). 9.17.24 Significance of Evaporites in the Earth History There are many evaporite deposits present among the sedimentary rocks on Earth. Kozary et al. (1968) estimated that approximately 25% of continental areas are underlain by ancient evaporite deposits and more large salt deposits have been discovered since that time, for example, the Messinian saline giant at the bottom of the Mediterranean. Evaporites are encountered in (1) cratonic basins, (2) synrift basins, (3) Paleogeographic Indicators Evaporite formation requires a relatively restricted climate, supporting a negative water balance. Ancient saline giants also needed specific paleogeography related to tectonic pulses of Earth and accommodation space for deposition. The evaporite deposits form and are preserved mainly in areas of limited rainfall and elevated evaporation. Today, they are largely concentrated within two ‘dry’ subtropical high-pressure belts between 15" and 35" latitude from the equator (Borchert and Muir, 1964; Lotze, 1938). Compilation by Ziegler et al. (2003), Zharkov (1998, 2001), and Chumakov and Zharkov (2003) proved that from the Permian until now, the evaporite deposition prevailed within the two belts 10–40" N and 10–40" S, apparently related to hot and dry climate conditions created by the descending branches of the Hadley cells. This appears to indicate that global atmospheric circulation was approximately the same as today at least since Permian. In particular, Ziegler et al. (2003) noted that the distribution of evaporites in time was related to availability of epeiric and shelf sea basins within lower latitudes rather than to global events of dry climate. Some ancient evaporite basins, like the giant Permian basins of north hemisphere, extended up to ! 50" latitude (Trappe, 2000; Ziegler et al., 2003). The evaporites are good paleogeographic indicators suggesting the position of the Phanerozoic basins did not extend beyond 35–50" latitude (Briden and Irving, 1964; Parrish et al., 1982; Rees et al., 2004; Zharkov, 1998). They can be used for the correction of paleogeographic reconstructions based on paleomagnetic data (e.g., Drewry et al., 1974; Evans, 2006). 9.17.24.2 Seals for Hydrocarbons and More (Evaporites and Hydrocarbons) Unlike the other sedimentary rocks, evaporite deposits, particularly halite, lose their porosity very early and rapidly, mostly due to early cementation (Garrett, 1970). Casas and Lowenstein (1989) have shown that saline pan halite is nearly entirely cemented by the burial depth of 10 m, where porosity is less than 10%. Halite cementation is promoted by evaporative concentration of groundwater brines and/or cooling of sinking surface NaCl-saturated brines. This very important feature enables the evaporites to create the seal necessary for hydrocarbon accumulations (Warren, 2006). This also enables the study of halite in underground mines in galleries, most of each is dry or nearly devoid of water, which is an unusual feature in mines excavated in sedimentary rocks. Large exploration chambers appear to be devoid of any influence of underground waters and therefore considered as perfect sites for the storage of radioactive waste (Roedder, 1984). However, the other specific feature – the ability of salt to flow due to its ‘halokinetic’ properties – decides that such potential sites may not be safe. This flow property can be rapid, in human terms, requiring only months or years, and can compromise any valid seal safety. Geochemistry of Evaporites and Evolution of Seawater 9.17.24.3 Halotectonics When during burial, halite deposits reach the depth where the density of overlying sediments surpasses the density of the halite, the halite may deform plastically and move slowly as a fluid, unless that overlying sediment is already cemented and rigid, acting as a cap to salt movement. Particularly, significant effects of salt deformation may occur in areas where the overlying load is unevenly distributed or the slope of the deposits gradually changes during subsidence. The moving salt is a powerful tectonic force able to completely modify the structure of overlying sedimentary cover. The salt becomes mobilized, and deforms or creeps downslope, out from beneath the overlying load, and may destroy or create entire sedimentary basins. In areas of rising salt diapirs, the salt is able to move to the Earth’s surface (diapiric structures), and in the zones of dry climate, the salt may flow out and build salt mountains. At the surface, in such zones, the rising salt (forced out by diapirism) even may begin flowing down the mountain slopes as namakiers or ‘salt glaciers’ (Talbot et al., 2009; Talbot and Pohjola, 2009). In a similar way, the huge masses of salt together, with their overlying sedimentary cover, are able to flow down the slope of the basins and inclined continental shelves. Flowing salt, squeezed out from below and between such flows, may form salt walls and canopies, as in the Gulf of Mexico, where such processes are responsible for creating numerous traps for hydrocarbons (see the review in Warren, 2006). Similar huge-scale structures occur in many areas of Earth and possibly on Mars (Hudec and Jackson, 2007; Montgomery et al., 2009). Halite also has another feature in that it has a very high heat capacity. Salt diapirs that rise toward the surface from great depths are, perforce, much warmer than nearer surface sediments (Mello et al., 1995) and conduct heat upward from depth. This process has two disparate results. First, the subsalt sediments are kept cooler than a regional geothermal gradient would suggest (slowing maturation of organic matter), and second, the rising salt brings warmth and migrating fluids through the sediment through which it passes, raising rates of diagenesis. Regionally, these two aspects of the effect of mobilized halite are rarely considered, but certainly, they may prove very important in many areas. 9.17.24.4 Diagenesis and Metamorphism of Evaporites In surface and subsurface, evaporites are prone to dissolution and are typical rocks responsible for development of karst. Subsurface dissolution of evaporites, particularly chloride salts, is considered as the main sources of saline fluids in many buried sedimentary basins (Carpenter, 1978; Hanor, 1994), and the composition of these fluids depends on the mineralogy of vanished buried and/or metamorphosed evaporite deposits (Lowenstein et al., 2003). Such fluids are particularly active in the areas of salt diapirs (McManus and Hanor, 1993). During burial, hydrated salts and particularly the most common gypsum undergo dehydration and gypsum passes into anhydrite releasing large volumes of water (Jowett et al., 1993), which is then an active agent in diagenetic transformations of surrounding (overlying) evaporites, carbonates, and other rocks (Borchert, 1969; 547 Borchert and Muir, 1964). Sometimes, incongruent dissolution of hydrated minerals, such as carnallite, releases water (Harville and Fritz, 1986). Due to the high solubility and reactivity, the evaporites themselves do not survive long in the zone of metamorphism. In some cases, however, both anhydrite and halite (that are part of well known sedimentary sequences) are known to remain preserved in thick beds up to temperatures well over 300 " C (Lugli, 1996a,b), preserving primary isotopic signals (Boschetti et al., 2011a), and when these data are pressurecorrected, the actual temperatures were closer to 380–400 " C (Lugli et al., 2002). However, hydrothermal salts, formed from hot circulating waters, may precipitate to infill voids but do not appear as bedded sequences along with layers of dolomites and mudstones. In the metamorphic realm, the evaporites usually are completely remobilized and create high-salinity subsurface fluids or brines. These and related fluids, on the other hand, are able to carry and precipitate sedimentary metallic ore deposits as for example in the Mississippi Valley-type Zn–Pb deposits. The deep subsurface basinal fluids are nearly entirely represented by high-salinity Ca–Na–Cl brines (e.g., Carpenter, 1978; Lodemann et al., 1997). Although the origin of these brines remains controversial, one of the accepted explanations is the infiltration of NaCl-rich evaporite brines and their interactions with Ca-rich rocks (e.g., Derome et al., 2007; Möller et al., 2005, with references). The quantitative Ca–Na relations in these brines may be explained by dolomitization or albitization of plagioclases (Boschetti, 2011; Carpenter, 1978; Davisson and Criss, 1996). Recently, Lowenstein et al. (2003) suggested that these brines represent a true relic of evaporated ancient seawater of the Ca chloride type, but the concept remains highly controversial as pointed out by Kharaka and Hanor (2003) and Hanor and McIntosh (2006). That view is a logical implication of the halite inclusion studies, which strongly suggested that there were long periods in the Phanerozoic history of the ocean when its halite brines were nearly devoid of sulfates and that oceans producing sulfate brines, such as today, were possibly rare events in the Earth history. In recent studies, Lowenstein and Timofeeff (2008) showed that this concept of relic Ca chloride brines is only partly true. 9.17.25 Summary Marine evaporites are the chemical deposits, which represent the direct record of the chemistry of ancient oceans and the soluble ionic load present in their waters. Crucial to our understanding of the origin of these evaporites is the order of salts precipitated from evaporated modern seawater: the Ca carbonates ! Ca sulfates ! Na chlorides ! Mg–K sulfates and chlorides. This mineralogical order is usually preserved in the Phanerozoic evaporites, although Mg–K salts are rarely present. This sequential relationship is the cornerstone of the widespread idea that the chemistry of the ocean was stable and remained constant throughout Phanerozoic time and presumably was similar to today’s ocean even earlier, in the Proterozoic and possibly the Archean. The advance of studies over the past few decades led to the emergence of the new and evolving picture of the ocean. What was previously expressed by a few forgotten investigators appeared to be the true and we are now quite certain that 548 Geochemistry of Evaporites and Evolution of Seawater the chemistry of the ancient oceans changed with time, and the changes were profound and relatively rapid. The pair of ions Mg2þ and Ca2þ apparently was crucial in the evolution of the sedimentary record. In the Phanerozoic, the molar ratio of these ions Mg2þ/Ca2þ changed with time, varying from a minimum estimated value of about one in Cretaceous (Aptian) to today’s maximum known value !5. The change in the ratio of these ions, as well as the other remaining ions (particularly SO42#), had an oscillating character and is unmistakably reflected in the clear changes in the depositional record of potassium and magnesium salt deposits. Such KCl salts, dominant in much of the Phanerozoic, alternated with two (or three) phases of magnesium sulfate salt deposition: first (controversial) in late Neoproterozoic, then in the Permian and Cainozoic. The last period appears to be the most peculiar considering the well-documented extremely high rate of these changes, the fact that Mg2þ/Ca2þ ratio has approached the very high value (5.2 – the highest recorded in Phanerozoic), and that the driving forces of these changes remain controversial. Evaporites cannot help much in our understanding of the chemistry of early Precambrian oceans; they were removed from the fossil record and are mostly known from poorly preserved pseudomorphs, commonly of a controversial derivation. True, well-preserved marine evaporites are unrecognized from that early, very long time interval. Many facts suggest that the Precambrian ocean was different than today. It is however highly possible that it behaved like that of the Phanerozoic and the chemical changes in evaporite composition were even more rapid and drastic, and possibly also oscillating in nature. Acknowledgments The authors thank Krzysztof Nejbert for helping in the preparation of the diagrams, David A.D. Evans for supplying the information about saline giants, and Fred T. Mackenzie for editorial improvement of the text of this paper. References Aharon P, Kolodny Y, and Sass E (1977) The hot brine dolomitization in the ‘Solar Lake,’ Gulf of Elat, isotopic, chemical, and mineralogical study. Journal of Geology 85: 27–48. Al-Droubi A, Fritz B, Gac J-Y, and Tardy Y (1980) Generalized residual alkalinity concept; application to prediction of the chemical evolution of natural waters by evaporation. 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