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Home All issues Volume 190 (2019) BSGF - Earth Sci. Bull., 190 (2019) 11 Full HTML
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BSGF - Earth Sci. Bull.
Volume 190, 2019
Article Number 11
Number of page(s) 9
DOI https://doi.org/10.1051/bsgf/2019010
Published online 13 September 2019

© A.-V. Bojar et al., Published by EDP Sciences 2019

Licence Creative CommonsThis is an Open Access article distributed under the terms of the Creative Commons Attribution License (http://creativecommons.org/licenses/by/4.0), which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.

1 Introduction

Evaporites consist mostly of ionic salts containing the major ions Na+, Ca2+, Mg2+, K+, Cl−, (SO4)2−, and (CO3)−2, structural water as well as other ionic constituents like Ba2+, Sr2+, Br−, Li+, I−, B3+. Approximately eighty different salt minerals have been identified for evaporite deposits, although only few of them are considered important rock formers (Holland, 1984). MgSO4 bearing marine evaporites are relatively rare in the geologic record and restricted to the late Precambrian, Pennsylvanian–Triassic, and Cenozoic (Eocene to modern) (Hardie, 1996). Generally, in evaporite deposits, minerals precipitate from sea water in reverse order of their solubilities: carbonates first, followed by sulfates, and finally chloride minerals. Polyhalite triple-cationic salt, K2Ca2Mg(SO4)4 2H2O, has a higher solubility than gypsum, polyhalite forming as near end member mineral in evaporite sequences, even after precipitation of halite.

Upper Permian potash-bearing rocks, including primary formed polyhalite and carnallite KMgCl3 6H2O, are found in the Delaware basin in the western Texas and south-eastern New Mexico (Lowenstein, 1988; Barker and Austin, 1993) as well as in NE England (Armstrong et al., 1951; Kemp et al., 2016) and the Zechstein basin of Germany (Kampschulte et al., 1998). The origin of polyhalite of Permian age from the Zdrada platform and Peribaltic Basin of Poland was investigated by Peryt et al. (1998) and Peryt et al. (2005), the authors concluding that it formed by reaction of anhydrite with marine brines. By determining the age of polyhalite by 40Ar/39Ar method, Schorn et al. (2013) and Leitner et al. (2013) linked tectonic events and mineral textures from the Northern Calcareous Alps.

Recently, formed diagenetic polyhalites replacing gypsum or anhydrite were investigated from Baja California, Mexico (Holser, 1966). Polyhalite may form as a back reaction between brine and glauberite Na2Ca(SO4)2, as demonstrated for the drill holes from the Lop Nor basin, China (Ma et al., 2010). Polyhalite occurs as end member precipitate in a recent playa lake of La Mancha, Spain (Pena et al., 1982) or in the evaporitic sequence of the Salt Lake, Tuz Gölü basin, Turkey (Camur and Mutulu, 1996). Potash salts, including polyhalite, form today also in the Qaidam basin, China (Lowenstein et al., 1989) or Sambhar Lake, Thar Desert, India (Sinha and Raymahashay, 2004).

Once formed, gypsum and polyhalite record in their texture or isotopic composition primary and/or secondary events, the last ones occurring much latter than their formation age. For example, for the Carpathian Foredeep, Hałas and Krouse (1982) investigated for Miocene gypsum the isotopic composition of water of crystallization and demonstrated that during the last glacial, isotopically depleted meteoric water was incorporated in gypsum as water of crystallization. Matsubaya and Sakai (1973) demonstrate that anhydrites and gypsum were rehydrated subsequently by various mechanisms. By using a different method, namely the 40Ar/39Ar dating, Leitner et al. (2013) conclude that the radiogenic system of Permian polyhalites from the Eastern Alps is not recording formation ages but subsequent Permian and Alpine events.

In the present study, we investigate mineral formation and subsequent events recorded in hydrated sulfates deposits by using multiple isotope systems recorded in the sulfate anionic group and water of crystallization. The investigated minerals are: gypsum and polyhalite of Upper Permian and Triassic age from the most prominent evaporitic sequences of the Alps and, in order to compare the effect of contrasting thermal history on the isotope composition of water of crystallization, selected samples from New Mexico (USA) and Kłodawa mine (Poland) and Hattberg, Hesse (Germany).

2 Geological setting

Sulfates occur in the halite and gypsum deposits of the Northern Calcareous Alps and in the so-called Central Alpine Mesozoic (Fig. 1). The two units of the Eastern Alps are characterised by the presence of representative and well exposed Permian-Triassic evaporitic deposits (Schorn et al., 2013; Leitner et al., 2013). Description of samples from various locations, mineralogy, texture as well as the age of the deposits are given in the supplementary material. Systematic presentation of the mineral deposits is given in Bojar et al. (2016).

Additionally, we analysed two polyhalites from New Mexico (Barker and Austin, 1993). Two others polyhalite samples are from the Zechstein facies of Europe, namely one from the Kłodawa Mine, Poland (Burliga, 2014) and the other from Hattdorf, Hesse, Germany (Kampschulte et al., 1998).

thumbnail Fig. 1

Geological overview of the Eastern Alps showing the distribution of investigated sulfates. Salt: 1-Hall in Tirol, 2-Hallein, 3-Hallstatt, 4-Altaussee, 5-Bad Ischl; Gypsum deposits, Northern Calcareous Alps: 6-Golling, 7-Wienern, 8-Lessern, 9-Tragöß, 10-Seewiesen; Gypsum deposits, Central Alpine Mesozoic: 11-Stanz, 12-Göstritz. Overprint temperature of sulfates were calculated using sulfur isotope thermometer (Bojar et al., 2016).

3 Analytical methods

The stable isotope data of sulfates were obtained on a dual inlet and triple collector mass spectrometer on SO2 prepared from the supplied BaSO4 samples by the method published by Hałas and Szaran (2001) in the case of δ34S measurements, and on CO2 prepared from these BaSO4 samples by the method of Hałas et al. (2007) in the case of δ18OSO4 measurements. NBS-127 was used as internal standard. Results are reported with respect to the international standards Canyon Diablo Troilite (V-CDT) and Standard Mean Ocean Water (V-SMOW), respectively, with standard deviation better than ±0.2‰.

Water of crystallization was extracted by heating the samples in vacuum. Firstly, the samples were treated in vacuum by heating at 50 °C for two hours in order to remove absorption water (Hałas and Krouse, 1982). Crystallization water was than collected by heating gypsum at 250 °C and polyhalite at 400 °C. The cryogenically trapped water was subsequently analysed for δ18O and δD on a Picarro L2120-i Analyzer. Laboratory standards, previously calibrated to the V-SMOW-VSLAP scale were OH-13 (δD = −2.84‰, δ18O = −0.96‰) and OH-16 (δD = −114.68‰, δ18O = −15.43‰) (Bojar et al., 2017). These standards were used to normalize the results to the V-SMOW-V-SLAP scale. The isotopic compositions of water of crystallization were synchronously measured, with a standard deviation of ±0.1‰ for δ18O and ±0.4‰ for δD.

4 Results and Interpretation

4.1 Results

For the present study, the measured δ34S and δ18O values of sulfates and δ18O and δD of water of crystallization are given in Tables 1 and 2, respectively. For the Northern Calcareous Alps, δ34S values for both gypsum and polyhalite of Late Permian age, fall in a narrow range between 10 and 13‰, except one sample with 14‰ (Tab. 1). Triassic (Carnian-Norian) sulfates from the Central Alpine Mesozoic, consistently show a more enriched isotopic composition, with δ34S values between 14 and 15‰. In contrast to sulfur, the δ18OSO4 values of gypsum and polyhalite show a broader distribution between 12 and 20‰, with one exception (H_5) showing 6.5‰ (Tab. 1).

The δ18O and δD values from gypsum crystallisation water vary from −1.8 to −11.1‰ and from −50.3 to −107‰, respectively (Tab. 2). Polyhalite δ18O and δD values vary from 13.3 to 7.4‰ and from −1.2 to −26‰, respectively.

Table 1

Isotopic composition of sulfates for which water of crystallization was measured.

Table 2

Isotopic composition of water of crystallization and brine.

4.2 Interpretation

4.2.1 Isotopic fractionation between hydrated salts and water

4.2.1.1 Fractionation between water of crystallisation in gypsum and brine

According to Gonfianti and Fontes (1963), the fractionation factor for oxygen is α = 1.0037; Fontes and Gonfiantini (1967) determined the fractionation factor for hydrogen as α = 0.98, both fractionation factors being not temperature-dependent.

For gypsum, a summary of available fractionation factors between oxygen in brine and oxygen in water of crystallization is given by Herwartz et al. (2017). The authors also determined a fractionation factor of 1.0034, but used for their calculations the factor of 1.0037 as determined by Gonfianti and Fontes (1963). Herwartz et al. (2017) mention that by using one of the already determined factors of 1.0034 to 1.004 the calculated oxygen composition of brines will not significantly change. Concerning hydrogen, Gazquez et al. (2017) determined a fractionation factor in good agreement to Fontes and Gonfiantini (1968). In the present case, the small temperature dependence found by Gazquez et al. (2017) does not change our interpretation. The fractionations of Gonfianti and Fontes (1963) and Fontes and Gonfiantini (1967) approximate better the present situation where we have not precise constrains for temperature.

4.2.1.2 Fractionation between water of crystallisation in polyhalite and brine

Hydrogen isotope fractionation between hydroxyl bearing minerals and water is for most of the determined minerals less than 1 (Chako et al., 2001 and reference therein). It has been shown that salt hydrates crystallization causes deuterium enrichment of brine (Barrer and Denny, 1964; Horita, 1989). For sulfate bearing minerals, both hydrogen from the hydroxyl group as well as from water of crystallization are deuterium depleted in respect to brine (Seal et al., 2000). The hydrogen isotopic fractionation of crystallization water for the complex polyhalite structure is not precisely determined. If we consider a fractionation factor less than one, for example in the range determined for alunite-jarosite for the hydroxyl group (Seal et al., 2000), than the calculated values of brines in the present study will not change in order to affect conclusions. The fractionation factor for hydrogen in gypsum (α = 0.98) is between those determined for structural water in alunite (α = 1.004) and jarosite (α = 0.050). The number of determined 18O fractionation for salt hydrates are even smaller (Horita, 1989; Seal et al., 2000) and in the order of magnitude of this determined for gypsum. According to the reasoning, we consider for polyhalite the fractionation factors as for crystallization water in gypsum and calculated the values for brines as displayed in Figures 2 and 3.

thumbnail Fig. 2

Plot of δl8O of brines versus δ18OSO4 for the investigated sulfates: blue dots represent gypsum and red polyhalite isotopic compositions. “Trend 1” indicates the thermal event affecting both anhydrite/gypsum and polyhalite δ18OSO4: “trend 2” shows that subsequently crystallization water of gypsum was replaced by meteoric water at low temperature. δ18OSO4 values are the same range for both minerals, gypsum and polyhalite, suggesting that thermal overprint shown by “trend 1” affected both to similar extent. Isotherm calculation was done according to Zeebe (2010). Blue horizontal line represents Permian oxygen isotopic composition in sulfates precipitated from seawater.
thumbnail Fig. 3

Plot of δD versus δl8O of brines (blue dots: gypsum crystallization waters; pink dots: polyhalite crystallization waters). The ellipse is displaying the field of syngenetic brines from the Kłodawa salt mine (for details see Dulinski et al., 2014). Other elements displayed are: Permian ocean water isotopic composition similar with the present one, Knauth and Beeunas (1986); the present Local Meteoric Water Line (LMWL); the mean δl8O value of the present day precipitation (red circle); the pre-Holocene meteoric water line showing a similar slope to the present day but with δl8O mean values with c. 3‰ lower; LMWL data according to the IAEA database for the Klagenfurt airport station (IAEA database, 2010), the red point from the LMWL represents the mean value.

4.2.2 Calculated isotopic composition of brines from the measured water of crystallisations in gypsum and polyhalite

Using the Gonfianti and Fontes (1963) fractionation factors for oxygen (α = 1.0037) and Fontes and Gonfiantini (1967) for hydrogen (α = 0.98), the calculated isotopic compositions of brines are shown in Table 2. The equation for the fractionation factor is given by α = (1000 + δwater crystallisation)/(1000 + δbrine). For gypsum, the δ18O and δD values of brines vary from −5.7 to −15‰ and from −30.9 to −88.8‰, respectively. For polyhalite, the δ18O and δD values of brines vary from 14.4 to 3.4‰ and 42.5 to −6.1‰, respectively (Tab. 2).

5 Discussion

As temperature of dehydration is important for the interpretation of the results, data regarding this aspect will be discussed in the following paragraph for the two investigated minerals, gypsum and polyhalite.

5.1 Dehydration temperature of gypsum and polyhalite

Polyhalite loses crystallization water at approximately 285 °C (Fisher et al., 1996), its structure being more stable than that of gypsum, which loses crystallization water at 200 °C (Kemp, 1999). According to Wollmann et al. (2008), polyhalite dehydration starts at ∼255 °C, the reaction being completed at ∼343 °C. Polyhalite precipitates from saturated solutions at more elevated temperature than normal laboratory conditions of ∼20 °C, increase of temperature to boiling driving crystallization (D’Ans, 1908). In this context, we consider that temperature of polyhalite crystallization is an additional factor for determining the geological time-period when polyhalite mineral deposits formed. For example, large volumes of polyhalite are known for Permian. In addition to the sulfate-rich oceans present at that time, brine temperatures around 50 °C (Hałas et al., 2015) may have further enhanced the formation of the mineral.

5.2 Isotopic composition of sulfur and oxygen in the sulfate group; extent of thermal overprint

Sulfur isotope chemostratigraphy curve constructed with sulfates recovered from marine evaporites and marine carbonate deposits were for example constructed by Claypool et al. (1980), Kampschulte et al. (1998), Paytan and Gray (2012), Sial et al. (2015). The sulfur isotope data from Tab. 2 suggest that sulfates largely preserve their initial marine sulfur isotopic composition, a discussion for the Alpine realm is included for example in Longinelli and Flora (2007), Boschetti et al. (2011), Bojar et al. (2016) and references inside.

According to Longinelli (1983), the oxygen isotopic composition of marine present-day precipitated sulfates is 13‰, similar with the Permian one as determined for exemple by Boschetti et al. (2011); Bojar et al. (2016) (Fig. 2). For sulfates precipitated at ambient temperatures, the measured oxygen isotope composition is largely in disequilibrium with the oxygen isotope composition of water (Chiba and Sakai, 1985). Hałas and Pluta (2000) demonstrate that in the case of long residence time, the isotopic re-equilibration process between oxygen in gypsum sulfate group and oxygen in water molecules of brines may start at temperatures of 30–35°C, their findings being in accordance with Zeebe (2010). For isotherms calculation, we consider the oxygen fractionation factor between dissolved sulfate and water of Zeebe (2010) similar to that of Hałas and Pluta (2000), for details see also Boschetti (2013). During calculation, we also included the precipitation effect on oxygen isotope considering a fractionation factor of approximately 3% between precipitated and dissolved sulfates (Lloyd, 1968; Szaran et al., 1998). The position of isotherms will not significantly change if we use for 18O in sulfate one of the following fractionation factors: anhydrite-water (Chiba et al., 1981), alunite-jarosite-water (Stoffregen et al., 1994; Rye and Stoffregen, 1995) or barite-water (Kusakabe and Robinson, 1977). In the present case, we interpret the shift of δ18OSO4 toward higher values than these typical for marine evaporites of Permian-Triassic age resulting from partial equilibration of oxygen in sulfates with a brine, for example, this resulted from the dehydration of gypsum during the thermal overprint (Fig. 2).

5.3 Isotopic composition of water of crystallization and brine

The δ18O values of calculated brines using water of crystallization (x-axis) versus δ18OSO4 values (y-axis) are displayed in Figure 2. Gypsum and polyhalite plot in two distinct groups, with gypsum showing negative oxygen isotope values and polyhalite showing positive values of brines. Considering the measured δ18OSO4 values, the isotopic values of water of crystallisation and the isotherms presented in Figure 2, we may distinguish between several processes. Oxygen isotopic composition of gypsum shows partial equilibration of sulfate group with local brines, as their δ18OSO4 values are higher than those of the Permian marine sulfate, and this is indicated as “trend 1” (Fig. 2). Polyhalites show also a similar trend of partial equilibration between oxygen in the sulfate group with brine oxygen. A compact, fine-grained polyhalite (H_5) shows also thermal overprint, in this case at higher temperatures. The calculated value of ∼300 °C is too high and should be considered qualitatively, as indicating preservation of peak temperatures during the overprint event. A vein polyhalite (D_4) and a coarse crystalline polyhalite (D_2) show partial equilibration of sulfate oxygen isotopic composition with brines at temperatures between 100° and 200 °C (Fig. 2). In contrast, oxygen and hydrogen isotopic composition of brines plot on a mixing line suggesting a meteoric component (Fig. 3).

Considering the available samples, we may distinguish for polyhalite between:

  • coarse polyhalite recording isotopic composition similar to evaporation brines evolving from marine water. The most enriched isotopic compositions are shown by samples NM1, NM2 (New Mexico) followed by K1 (Kłodawa Mine), sampled from evaporitic deposits, which were not deformed and thermal overprint in contrast to the evaporites of the Eastern Alps. In this frame, the samples from New Mexico and Kłodawa Mine may represent the isotopically enriched end-member of primary brines (Fig. 3);

  • recrystallized orange polyhalites (D_2, H_5, D_4), which indicate either isotopic equilibration between meteoric waters and polyhalite or incorporation of meteoric waters into crystal lattice of the salt; H_5 shows an unrealistic high temperature of ∼300 °C, higher than polyhalite dehydration limit;

  • fibrous polyhalite (HH1, A_6, H_1) formed in veins from salt rock. These samples record a latter imprint at estimated temperatures between 100° and 200 °C;

  • granular polyhalites grown on anhydrite or mud, which also indicate a higher temperature imprint (A_1, BI_3).

Monitoring of brines from Kłodawa Mine (Dulinski et al., 2014) has shown that enriched syngenetic water of Zechstein age is still present in closed salt pockets. The oxygen and hydrogen isotopic compositions of sygenetic brine from which Kłodawa polyhalite formed is shown in Figure 3, its composition plotting in the field of the measured brines from the salt pockets (Dulinski et al., 2014).

Permian-Triassic sulfates of Eastern Alps were thermally overprinted (Leitner et al., 2013; Bojar et al., 2016). Using illite crystallinity, vitrinite reflectance and fluid inclusions in anhydrite, Leitner et al. (2013) estimated peak overprint temperatures affecting polyhalite bearing deposits at c. 200 °C. For evaporite deposits from the Eastern Alps, Bojar et al. (2016) estimated overprint temperatures using sulfate-sulfide isotope thermometers between 190° and 260 °C (except for Tragöß gypsum deposit). As polyhalites were not thermally overprinted over 275 °C, they have preserved their water of crystallization and rather (re)crystallized at temperatures between 100 and 200 °C. On the other hand, secondary growth features were found for polyhalite of the Eastern Alps (Leitner et al., 2013; Bojar et al., 2016) showing recrystallizations younger than Permian. We may assume that during a latter process partial equilibration and /or incorporation of low amount of meteoric fluid (up to 10% for the data acquired in this study) took place at temperatures closed to ambiental ones. For example, Hałas and Krouse (1982) and Kasprzyk and Jasińska, (1998) identified various amounts of meteoric water in the crystallization water of Miocene gypsum from the Carpathian realm. Tan et al. (2014) showed that isotopic compositions of Quaternary gypsum formed in Qaidam basin in an arid environment plot on a mixing line between lake brine and recharge river water.

5.4 Relationship between the measured water of crystallization, calculated brines and local meteoric water line

Alpine overprint temperatures over 200 °C were high enough in order to drive dehydration of gypsum to anhydrite. Rehydration to gypsum is possible if fluids are available and temperature falls below the dehydration limit of the mineral. The calculated isotopic compositions of brine (using the oxygen and hydrogen isotopic compositions of gypsum crystallization water) indicate that during rehydration/recrystallization calcium sulfate (gypsum, anhydrite) incorporated meteoric water, as shown in Figure 2 by “trend 2”. The altitudes from which gypsum samples were collected for this study range between 480 and 970 m (Tab. 2). Incorporated meteoric water has δ18O and δD values from −5.7 to −15‰ and −30.9 to −89‰, respectively (Fig. 3), the isotopic values do not correlate with the altitude (Tab. 2). According to Froehlich et al. (2008), for similar altitudes, weighted annual mean δ18O values are from −10.3 to −10.6‰. These values are higher than those measured for the water of crystallization in the investigated gypsum samples. The water of crystallization for the gypsum samples is also plotting to the left of the Local Meteoric Water Line (LMWL) (Fig. 3). It has been documented that late-glacial meteoric water has lower isotopic composition than the late-Holocene one, with δ18Olate-Holocene>δ18Olate-glacial by more than 3‰ (Jasechko et al., 2015). Moreover, the Younger Dryas period represents the last major period of aquifer refill (Bath et al., 1979). In light of these findings, the lower isotopic values of crystallization water plotting left of the meteoric water line suggest that meteoric water was incorporated in gypsum in a period cooler and dryer than today, possible during the Younger Dryas, if not earlier. An alternative explanation could consider that gypsum crystallization water records the isotopic signature of meteoric water percolating downward from a much higher elevation than today into subsurface, thereby reaching the sulfate deposit. The incorporation of water characterized by lower isotopic values than present one was put in evidence also for Miocene sulfates (Hałas and Krouse, 1982). In Figure 3, gypsum samples from the halite type deposits, HT_2, H_4, A_5 (all transparent gypsum crystals, see supplementary material) show an isotopic composition of crystallization water close to the present global meteoric water line, suggesting late crystal growth.

6 Conclusions

Polyhalite and gypsum from the Eastern Alps are of marine origin, as suggested by their sulfur isotopic composition. As the δ18OSO4 values are in the same range for gypsum and polyhalite, the thermal overprint of 100 to 200 °C affected both minerals to similar extent.

The two minerals plot in two distinct groups (Figs. 2 and 3), which are not related to different precipitation stages in the marine evaporitic sequence but to different source for the incorporated crystallization water. Polyhalites from Carlsbad potash district, New Mexico and those from Kłodawa Mine (Poland) preserve an enriched syngenetic isotopic composition, characteristic for brines evolving from marine water. Polyhalites from the Eastern Alps indicate a narrow range for δD and δ18O of syngenetic brines with enriched oxygen and hydrogen isotopic composition. For a few samples mixture with meteoric water was involved, as indicated by the shift of values toward the meteoric water line.

In the case of the Eastern Alps, the crystallization water of gypsum records incorporation of meteoric water during a later event when, related cooler conditions meteoric water was depleted in heavy isotopes. A few gypsum crystals contain crystallization water whose isotopic compositions (δD and δ18O) are close to those of the present-day MWL.

Acknowledgments

The authors would like to thank to Stefan Nicolescu (Yale Peabody Museum of Natural History) for providing a polyhalite sample from New Mexico. We thank native speakers for checking the text. Christophe Lécuyer and Tiziano Boschetti are thanked for interesting comments and suggestions. An anonymous reviewer is thanked as well. An English native speaker checked the text.

We acknowledge CEEPUS mobility grant for Ana-Voica Bojar, facilitating cooperation with Mass Spectrometry Laboratory, Institute of Physics, Maria Curie-Skłodowska University, Lublin.

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Cite this article as: Bojar A-V, Hałas S, Bojar H-P, Trembaczowski A. 2019. Multiple isotope tracers from Permian-Triassic hydrated sulfates: Implications for fluid-mineral interaction, BSGF - Earth Sciences Bulletin 190: 11.

All Tables

Table 1

Isotopic composition of sulfates for which water of crystallization was measured.

In the text
Table 2

Isotopic composition of water of crystallization and brine.

In the text

All Figures

thumbnail Fig. 1

Geological overview of the Eastern Alps showing the distribution of investigated sulfates. Salt: 1-Hall in Tirol, 2-Hallein, 3-Hallstatt, 4-Altaussee, 5-Bad Ischl; Gypsum deposits, Northern Calcareous Alps: 6-Golling, 7-Wienern, 8-Lessern, 9-Tragöß, 10-Seewiesen; Gypsum deposits, Central Alpine Mesozoic: 11-Stanz, 12-Göstritz. Overprint temperature of sulfates were calculated using sulfur isotope thermometer (Bojar et al., 2016).
In the text
thumbnail Fig. 2

Plot of δl8O of brines versus δ18OSO4 for the investigated sulfates: blue dots represent gypsum and red polyhalite isotopic compositions. “Trend 1” indicates the thermal event affecting both anhydrite/gypsum and polyhalite δ18OSO4: “trend 2” shows that subsequently crystallization water of gypsum was replaced by meteoric water at low temperature. δ18OSO4 values are the same range for both minerals, gypsum and polyhalite, suggesting that thermal overprint shown by “trend 1” affected both to similar extent. Isotherm calculation was done according to Zeebe (2010). Blue horizontal line represents Permian oxygen isotopic composition in sulfates precipitated from seawater.
In the text
thumbnail Fig. 3

Plot of δD versus δl8O of brines (blue dots: gypsum crystallization waters; pink dots: polyhalite crystallization waters). The ellipse is displaying the field of syngenetic brines from the Kłodawa salt mine (for details see Dulinski et al., 2014). Other elements displayed are: Permian ocean water isotopic composition similar with the present one, Knauth and Beeunas (1986); the present Local Meteoric Water Line (LMWL); the mean δl8O value of the present day precipitation (red circle); the pre-Holocene meteoric water line showing a similar slope to the present day but with δl8O mean values with c. 3‰ lower; LMWL data according to the IAEA database for the Klagenfurt airport station (IAEA database, 2010), the red point from the LMWL represents the mean value.
In the text

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