Chlorine stable isotope fractionation in evaporites

Geochimica et Cosmochimica Acta, Vol. 59, No. 24, pp. 5169-5175, 1995 Copyright © 1995 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/95 $9.50 + .00

Pergamon

0016-7037(95)00353-3

Chlorine stable isotope fractionation in evaporites H. G. M. EGGENKAMP, I , t R. KREULEN, I and A. F. KOSTER VAN GROOS2 Department of Geochemistry, Utrecht University, P.O. Box 80.021, 3508 TA Utrecht, The Netherlands 2Department of Geological Sciences, The University of Illinois at Chicago, Chicago, Illinois 60680, USA (Received February 4, 1994; accepted in revised form September 7, 1995)

Abstract--Chlorine isotope fractionation (37C1/35C1) between NaC1, KC1, and MgC12.6H20 and their saturated solutions was determined in laboratory experiments at 22 -4- 20(2. The results are as follows: 103 In ot(NaCl--solution) = +0.26 _+ 0.07 ( h r ) , 103 In a ( K C l - - s o l u t i o n ) = - 0 . 0 9 _+ 0.09 ( l a ) , 10 3 In a(MgC12" 6HEO--solution) = - 0 . 0 6 _+ 0.10 (1cr),

where the fractionation factor a is defined as: 37C1] 35C1)precipitate Ol = -- (37Cl[35Cl)solutio n .

( 1)

These data were used to approximate the isotope fractionation factors of chloride between the saturated solution and halite, kainite, carnallite, and bischofite. From the results, the stable chlorine isotope fractionation during the formation of evaporite was calculated using a Rayleigh fractionation model. The model predicts that 6 37CI of the precipitate decreases systematically during the main phase of halite crystallization but increases again at the latest stage of evaporation. The chlorine isotope fractionation model was tested on a core from the upper Zechstein III salt formation. The salt core contains layers dominated by either halite or K-Mg salts. The K-Mg salts, which are formed during the final stages of evaporation, contain up to 75% carnallite (KMgC13-6H20) and bischofite (MgCI2.6H20). The observed chlorine isotope fractionation in the salt core is in general agreement with the Rayleigh fractionation model. During the main crystallization phase of halite, 637C1 decreases continuously, but this trend reverses during the final stages when Mg-salts begin to crystallize. It is concluded that 637C1 call be used as an indicator of evaporation cycles. In addition, it provides quantitative information on the proportion of salt that has been deposited on the input of fresh seawater and on the disturbance by postdepositional processes. 1. INTRODUCTION

no significant fractionation occurs. More recently, Kaufmann et al. (1984, 1988) measured a few halite samples. Their results showed that these halites are enriched in the heavy isotope 37C1. Vengosh et al. (1989), using negative thermal-ionization mass spectrometry to measure chlorine isotope variations in evaporites from inland ponds in China and Australia, found an extreme range of 637C1 values (32%0, with a large standard deviation ( h r ) of 0.7 to 2%0). They did not find systematic fractionation trends, although the more evaporated samples tend to deviate more, either in positive or in negative direction. Their results have not been confirmed by conventional analytical methods. Sulphates from evaporites, in contrast, show significant isotope fractionation, as they are affected by biological processes such as reduction of sulphate by bacteria (Ault and Kulp, 1959; Eriksson, 1963; Nielsen, 1966; Rees, 1970). During evaporation, some nonbiological fractionation of the sulphur isotopes also occurs, resulting in a slight change in the isotope ratio of the precipitates, when compared to the seawater from which the evaporite was formed. This effect is illustrated by the consistently lower sulphur isotope ratio's values found in K-Mg sulphate facies relative to the gypsum and anhydrite facies (Nielsen and Ricke, 1964; Holser and Kaplan, 1966). Thode and Monster ( 1965 ) determined from both experimental and calculated data that the sulphur isotope fractionation

Fractionation of chlorine isotopes in evaporites has rarely been investigated. The main reason is that earlier analytical methods lacked the analytical precision to detect the small variations in chlorine isotopic composition found in nature. That this variation is small compared to the much larger variation in sulphur isotopes, which show relatively large fractionation effects, is a consequence of the fact that chlorine does not take part in biochemical reactions. Since the early eighties, analytical methods have become sufficiently precise so that meaningful chlorine isotope measurements are now possible. In an early attempt to evaluate chlorine isotope fractionation, Hoering and Parker ( 1961 ) measured a series of evaptrite samples. However, they could not detect chlorine isotopic variations outside their limits of precision, _+0.3 to _+0.9%0 (ltr). Accompanying experiments on the isotope exchange between NaC1 crystals and a saturated solution indicated a fractionation factor a of 1.0002 _+ 0.0003 (a = ( 37C1/3sC1)precipitate/(37C1/3sCl)~olutio,). They concluded that * Present address: Postgraduate Research Institute for Sedimentology, The University of Reading, P.O. Box 227, Whiteknights, Reading RG6 6AB, United Kingdom. *Author to whom correspondence should be addressed.

5169

5170

H . G . M . Eggenkamp, R. Kreulen, and A. F. Koster van Groos

b e t w e e n precipitated g y p s u m and d i s s o l v e d sulphate is 1.00165 __. 0.00012. R e c e n t l y Raab and Spiro ( 1 9 9 1 ) e x a m ined the fractionation o f sulphate during the e v a p o r a t i o n o f seawater. T h e y f o u n d that different minerals gave different fractionation factors, for g y p s u m a is 1.00165, during the halite precipitation it is -- 1, and in the precipitation o f M g - s u l phates it is 0.99844. Thus, during precipitation o f sulphate minerals the sulphur i s o t o p e c o m p o s i t i o n m a y c h a n g e with --2.5%o, w h i c h is small w h e n c o m p a r e d to the d i f f e r e n c e s f o u n d in m a r i n e evaporites ( ~ 2 0 % o , C l a y p o o l et al., 1980). Therefore, not m u c h attention has b e e n g i v e n to this type o f fractionation. T h e f o l l o w i n g study reports o n a series o f e x p e r i m e n t s on the fractionation o f chlorine isotopes b e t w e e n brines and evaporites. T h e results s h o w that the total fractionation o f chlorine isotopes in evaporites, although small, is significant. C o m p a r i s o n o f 637C1 data f r o m these e x p e r i m e n t s and a core f r o m the u p p e r Z e c h s t e i n III s h o w s that the e v a p o r a t i o n state and history o f an evaporite sequence, input o f n e w seawater, and m i x i n g with partly r e d i s s o l v e d salt, especially during the precipitation o f halite, can be evaluated. 2. ANALYTICAL METHOD 2.1. Definition of the Isotopic Standard The chlorine isotope composition is generally expressed as ~37C1, which gives the permil deviation from a chlorine isotope standard. Although no international standard is defined, Kaufmann et al. (1984) showed that the seawater isotopic composition is constant in oceanic samples. Therefore, they proposed to use seawater as the chlorine isotope standard, SMOC (Standard Mean Ocean Chloride). In our laboratory a sample from the Atlantic Ocean, taken near Madeira (800 meter depth), is used to represent SMOC. ~37C1 is defined as 6 J7ClsMoc = ( R~"mv'~ -- RsM°c ~ * l O00, • \ RsMoc /

(2)

where R~mo~is the 37C1135C1ratio of the sample and RsMoc is the 3VCl/ ~sC1 ratio of the standard SMOC. 2.2. Analytical Procedure In this study the chlorine isotope ratios are measured on chloromethane gas (Taylor and Grimsrud, 1969; Kaufmann, 1984), using a slightly modified positive ion method. This method is described in detail in Eggenkamp et ai. (1994) and summarized here. The method is slightly different from the method described by Long et al. ( 1993 ). For the measurement of salt samples solutions containing 3000 ppm chloride are prepared by dissolving the sample in distilled water. 1 mL of the solution is mixed with 4 mL of a 1 M KNO3 solution to reach a high ionic strength (Taylor and Grimsrud, 1969), and 2 mL of a citric acid-phosphate buffer (20.6 gram citric acid and 0.71 gram Na2HPO4-2H20 per liter water, after McIlvaine, 1921 ), to reach a constant low pH to avoid precipitation of Ag2S (Kaufmann, 1984). After heating the solution to ~80°C, 1 mL of a 0.2 M AgNO3 solution is added and AgCl begins to precipitate from the hot solution. Next, the suspension is filtered over a 0.7/zm glass fibre filter (Whatman ® type G F / F ) . The filter plus precipitate are dried overnight in the absence of light. Next, the filter plus precipitate are brought into a Pyrex ® reaction tube, and a capillary is drawn from the tube above the filter. Then, 100 #L CH3I is frozen onto the sample, after which the tube is evacuated and sealed at the capillary. The glass capsule is kept for two days at 80°C. At these conditions, AgCI reacts to completion forming AgI and CH3C1. Following the reaction, the capsule is placed and sealed within a flexible stainless steel tube, attached to a gas chromatograph. After the capsule is broken, the gas mixture is transferred to a cold trap cooled with liquid N2. By switching to a

helium carrier gas line and heating the sample with hot water, the mixture is evaporated and brought onto a gas chromatographic column (75 cm filled with Poropack ® Q) to purify the gas. The purified CH3CI is stored in a glass gas vessel and measured directly on a VG SIRA 24 mass spectrometer equipped with adjustable collectors. The trap current is reduced to 100 #A in order to bring the minor beam down to values < 10 ~0 A that can be handled by the mass spectrometer while still maintaining sufficient gas pressure in the inlet system. 3. CHLORINE ISOTOPE FRACTIONATION EXPERIMENTS Three 250 mL beakers were filled with 50 mL water and 15 g NaCI, three with 50 mL water and 15 g KC1, and three with 50 mL water and 50 g MgC12" 6H20. The chlorides were dissolved in the water• All chemicals were of reagent grade. The solutions were placed in a fumehood at room temperature (22 _+ 2°C) and allowed to slowly evaporate. KCI, NaCI, and MgCI2-6H20 began to precipitate after 2 days, 3 days, and 28 days, respectively. As soon as the precipitate was observed, in general about one gram was present then, it was separated from the solution by filtering. The precipitate was rinsed with acetone to remove the remaining solution. Next, the precipitate was dissolved in distilled water. The solution of the precipitate and the coexisting solution were measured for chlorine isotopes as described above. One, two, or three chlorine isotope measurements were made of the NaC1 precipitate from each of the three beakers and of each solution (Table 1 ). The fractionation factor was determined by calculating the difference between all the measurements of the precipitates (six in total) and the solutions (also six in total). The fractionation ( 103 In a ) is the calculated average of these 36 values, the error as the ~r,~ t (n = 36 in this case). The value obtained for the NaCl precipitate-solution fractionation is +0.26 _ 0.07%o. The KCI and MgCI2-6H20 fractionation factors were determined in exactly the same way. For KC1, a total of seven measurements were made of the precipitate from three beakers and four measurements of the solution from two beakers (Table 1 ), giving 28 values and a fractionation of - 0 . 0 9 _+ 0.09%o (n = 28). For MgCI2 "6H20, six measurements were made of the precipitate from two beakers and five measurements of the solution from two beakers (Table 1 ), giving 30 values and a fractionation factor of - 0 . 0 6 _+ 0.10%¢ (n = 30). These fractionations are assumed to be equal to 103 In a, where a is the fractionation factor which is defined as OL- (

37 35 E l / C1)precipi .... (37C1/35C1 )solution "

(3)

4. 4537C1 EVOLUTION DURING PRECIPITATION OF SALTS FROM SEAWATER The sequence o f salt minerals that precipitate upon evaporation o f m o d e m seawater is as follows: g y p s u m (CaSO4' 2 H 2 0 ) , halite ( N a C 1 ) , bloedite ( N a 2 M g ( S O 4 ) 2 " 4 H 2 0 ) , e p s o m i t e ( M g S O 4 " 7 H 2 0 ) , kainite (K4Mg4C14(SO4)4" 1 1 H 2 0 ) , hexahydrite ( M g S O 4 - 6 H 2 0 ) , kieserite ( MgSO4- H 2 0 ) , carnallite ( K M g C 1 3 . 6 H 2 0 ) , and bischofite ( M g C 1 2 " 6 H 2 0 ) (Braitsch, 1962). Halite starts to precipitate after 90.9% o f the original s e a w a t e r v o l u m e has evaporated, w h e r e a s carnallite and bischofite precipitate w h e n 99.2 and 99.4%, respectively, o f the water has evaporated. M o r e relevant to the p r o c e s s o f chlorine isotope fractionation is the p r o p o r t i o n o f chloride that is precipitated f r o m the e v a p o r a t i n g seawater. Thus, 82.5% o f the original chloride c o n t e n t is precipitated as halite. Next, kainite starts to form, and, after 86.9% o f the chloride is r e m o v e d f r o m the brine, carnallite b e g i n s to precipitate. Finally, w h e n 88.9% is r e m o v e d , bischofite crystallizes (Braitsch, 1962). T h e c h a n g e in ~37C1 c a u s e d b y the precipitation o f salts from e v a p o r a t i n g s e a w a t e r is m o d e l e d here using a similar a p p r o a c h as H o l s e r and Kaplan ( 1 9 6 6 ) in their m o d e l o f sul-

Isotope fractionation of CI in evaporites

5171

Table 1: ~i37C1measurements of the precipitate and brine phase for the calculation of the fractionation factors for NaC1, KC1 and MgCI2.6H20. All values are per rail differences from the used reference gas and not against SMOC. /$37C1of the reference gas is -4.17%o relative to SMOC. The codes refer to the measurements out of a single beaker, and are built up as follows: Na, K and Mg indicate the cation of the salt, the number is the identification of the beaker, and C indicates a measurement of the precipitate and S of the remaining brine. NaC1

KCI

precipitate

brine

precipitate

code:NalC

code:NalS

4.25 4.41

precipitate

brine

code:K4C

code:Mg7C

code:Mg7S

4.05

3.17

4.40

4.50

4.09

3.07

4.40

4.40

4.35

4.55

code:Na2C

code:Na2S

code:K5C

code:K5S

code:MgSC

code:Mg8S

4.36

4.07

3.00

3.17

4.42

4.31

4.43

4.17

2.99

3.17

4.32

4.40

3.19

4.32

4.37 code:Na3C

code:Na3S

code:K6C

code:K6S

4.14

3.11

3.25

4.08

3.20

4.32

3.23

phur isotope variations in sulphates. However, their assumption that all fractionation factors for the formed minerals are the same was changed in our application of the model to chlorine isotopes. The basic model was derived from Mclntire (1963) to calculate trace element partition coefficients in systems such as crystallizing melts. Assuming ideal Rayleigh behaviour, the 37C1/35C1 ratio of the chloride in the solution changes according to

re= [ I - m~o] exp[a -

(4)

where rc/ro is the ratio of 37Cl/35C1 in the sample and 37Cl/ 35C1 in the initial seawater, and mclmo is the weight fraction of chloride precipitated relative to the initial amount available. T h e 637C1 of the brine, from the moment the first halite is precipitated, can be calculated as ~37C1 = 1000 ~ - r_____~0,

(5)

and of the precipitate by factoring in [a - 1] : 637C1 =

MgCI2.6H20 brine

lO00(c~ - l + rc - r°)

(6)

Braitsch (1962) distinguishes five main precipitation stages from normal seawater: a halite dominant, a bloedite/epsomite dominant, a kainite dominant, a carnallite dominant, and a bischofite dominant stage. During the first two stages, halite is the only precipitating chloride mineral. In the kainite stage, both kainite and halite precipitate chloride. During this stage, approximately 23 wt% is precipitated as halite and 61 wt% as kainite. Because kainite contains 14.5 wt% C I - of which one third is allocated to K ÷ and two thirds to Mg 2÷ , at this stage 61.2 wt% of the chloride is precipitated as NaC1, 12.9 wt% as KC1, and 25.9 wt% as MgC12. During the carnallite stage,

"

halite ( 12 wt% of total) and carnallite (48 wt% of the total) precipitate chloride. In terms of the simple chlorides, 28.4 wt% is precipitated as NaC1, 23.9 wt% as KC1, and 47.7 wt% as MgC12. During the last (bischofite) stage, bischofite (99 wt%), carnallite (0.25 wt%), and halite (0.5 wt%) precipitate, or, again in terms of the simple halides, 0.9 wt% of the chloride precipitates as NaCI, 0.1 wt% as KC1, and 99.0 wt% as MgC12. From the N a / K / M g ratios of the precipitating salts a new fractionation factor was calculated on the basis of the experimentally determined fractionation factors of the simple chlorides and assuming that the fractionation factors for the components are not affected by the other precipitating components (Table 2). Combining these fractionation factors, the isotopic evolution of the evaporating brine and the precipitating salts was calculated (Fig. 1). The results show that the 637C1 of the brine would decrease from 0 to -0.45%0 during the halite stage. During the kainite and the carnaliite stage, the fractionation factor is closer to unity and the decrease in 637C1 of the brine is only 0.04%o. Because the fractionation factor in the bischofite stage is less than unity, 637C1 must increase during the last stages of the evaporation. It should be noted that the differences in the fractionation factors in these salts is not unusual. For example, the fractionation factors of sulphur isotopes in evaporate sulphates (Raab and Spiro, 1991 ) show also different fractionation factors, which may be higher or lower than unity. The isotopic compositions of precipitate deviate by 103 In a from the curve for 637C1 of the brine. Because the fractionation factor is different for each salt mineral, a discontinuity occurs in 6 37C1 of the precipitate when a different salt starts to precipitate. 637C1 of the precipitate decreases from +0.26 to -0.19%~ during the halite stage. During the kainite stage, it further decreases to about -0.34%o, and to -0.47%o during the carnallite stage. However, during the last (bischofite)

5172

H . G . M . Eggenkamp, R. Kreulen, and A. F. Koster van Groos Table 2: Fractionationfactorsfor the halite,kainite,camallite,and bischoftteprecipitationstages, based on experimentallydeterminedfractionationfactorsfor NaC1, KC1,and MgC1E.6H20(see text). Compositiont

Precipitation stage

Fractionationfactor

100% NaC1

1.00026~:0.00007 1.00013±0.00008

carnallite

61.2% NaCI, 12.9%KCI, 25.9% MgCI2 28.4%NAC1,23.9% KC1,47.7% MgCI2

bischofite

0.9% NaCI, 0.1% KCI, 99.0% MgCI2

0.99994±0.00010

halite kainite

1.00002±0.00009

~Thepercentagesgive the approximatecompositionof the chloridesin each precipitationstage after BRAITSCH(1962), as calculatedin the text. stage, 637C1 increases from -0.55%0 and may reach positive values. Several uncertainties are present in the chloride evolution as presented above. Errors in the fractionation factors are relatively large, especially for the potassium and the magnesium chlorides, where they are not significantly different from unity. Consequently, calculated variations b e c o m e larger with increasing evaporation stage. For example, the calculated 637C1 of the brine at the end of the halite stage is between -0.58%o and -0.30%0 ( a range of 0.28%,) and for the halite precipitate the error increases from 0.14 to 0.44%o, using the uncertainty in the fractionation factor ( - 0 . 2 6 + 0.07%o), see Fig. 1. Because the fractionation factor for MgC12 is close to unity, the resulting ~37C1 o f the brine and the precipitate can evolve to either positive or more negative values (Fig. 1 ). However, preferred interpretation o f the experimental data indicates that the fractionation factor for bischofite is lower than unity, and we conclude that 637C1 during the last stage is likely to increase. The results in Fig. 1 show that during the halite stage the 637C1 of both the brine and the halite precipitate change monotonically. This suggests that it is possible to use chlorine isotopes in halite to estimate the amount of evaporation that has occurred. At later stages, where K- and Mg-salts precipitate, the method will be less effective since discontinuities appear

and the low degree of fractionation appears because of the low fractionation factors. If our interpretation of the experimental data is correct, and the fractionation factor of bischofite is less than unity, than a m i n i m u m value of ~37C1 in evaporites of about -0.55%0 can be expected, and later stage bischofite rocks have higher ~37C1 values. However, if the fractionation factor in this stage is higher than unity, then it can be expected that bischofite may have lower 637C1 values than the other chlorides (Fig. 1 ). Finally it must be noted that the model assumes Rayleigh fractionation, i.e., no isotope exchange after the salt has precipitated. It is likely that in natural deposits some exchanges will take place, and that the Rayleigh fractionation model is not absolutely adhered to. 5. CASE STUDY: SALT FROM A ZECHSTEIN CORE 5.1. Sample

Location

and Geologic

Setting

The isotope effects predicted by our model were compared with a series of samples from Zechstein core TR-2 ( f r o m the V e e n d a m structure, near Groningen, Northern Netherlands, location W H C - 2 ) , made available by Billiton Refractories B.V., see Fig. 2. The salt core was drilled for magnesium exploration (e.g., Coeleweij et al., 1978; Buyze and Lorenzen, 1986). It contains carnallite and bischofite in addition to kainite and halite. The core, therefore, allows the study of the

0.4 - -

6049 ' E 0.2 "'"

0.0

~

--...

precip~atl~

I

-0.2 -0.4 -0.6

,r'

-0.8

\l

-1.0

' 0.0

I 0.2

'

I

'

0.4 Fraction chloride

I 0.6

'

precipitated

I

I

0.8

1.0

FIG. 1. Calculated 637Ci values of the precipitate and the remaining brine. The bold lines indicate the average values. The area between the short-dashed lines indicates the error for the residual brines and the area between the long-dashed lines indicates the error for the precipitates. Note the discontinuities in the curves for the precipitates.

FIG. 2. Situation of the Veendam salt structure in the Netherlands after Coeleweij et al. 1978).

Isotope fractionation of CI in evaporites final stages of the evaporation process, which is especially interesting because the 637C! values in the late stage minerals may provide an indication whether the fractionation factor for the bischofite phase is either higher or lower than unity, as was discussed above. The Zechstein (Upper Permian) is the period during which the largest salt deposits in the history of the earth were deposited (Braitsch, 1962). The core is from the upper part of the Zechstein III formation (Nederlandse Aardolie Maatschappij and Rijks Geologische Dienst, 1980). It samples the subformations Zechstein III-1, -2, and -3. All three subformations are again divided in a lower part a, which contains mainly halite, and a carnallite- and bischofite-rich upper part b. Most samples come from the subformations Z-III-1 and Z11I-2. These formations were described by Coeleweij et al. (1978) using a different core. Striking differences are found between the two subformations. Subformation Z-III-2 (upper half of core) consists of only one evaporation cycle, while ZIII- 1 is much more complicated and exists of nine evaporation cycles with several subcycles. Twenty-five samples were taken from the core from 16281786 meter depth. Their stratigraphic position is shown in Fig. 3. Below 1770 m (Zechstein-III-la), the core contains massive halite with some sylvite (KCI) and langbeinite (K2Mg2 (SOn)3) layers. These minerals are not primary, thus they indicate some secondary processes. From 1770-1693 meter, a layered sequence ( Z - I l l - l b ) of carnallite, bischofite, halite, and some kieserite is found. From 1693-1654 meter, formation Z-III-2a with halite and some carnallite occurs. Between 1654 and 1640 meter, carnallite, halite, and kieserite are present. Here the core is red with some banding (Z-Ili2b). From 1640-1631 meter it contains halite with some carnaUite (Z-III-3a). Above this the core contains halite, carnallite, and kieserite (Z-III-3b, Haug, 1982). The two potassium-magnesium cycles, Z-III-lb and -2b, are very different (Coeleweij et al., 1978). Z-III-2b shows no distinct subcycles, whereas in Z-III-lb nine subcycles are present, each containing several mini-cycles. O f Z-III-2 a sample was taken every 10 meters and analysed for t~37C1. Because of the more irregular behaviour of Z-III-1 sampling was less uniform because of the irregular alternation of halite rich and halite poor layers. 5.2. Results and Discussion

Depths, salt compositions (as determined by Billiton, and indicating the average salt mineral composition per interval of the core, as shown in Fig. 3 ), and 63vC1 data, of the samples are shown in Table 3 and in Fig. 3. t~37C1data are determined on ground samples of bulk salt from the indicated depth. Samples are dissolved in distilled water and measured for 637C1 as indicated above. A distinction can be made between the halite stages (Z-III-la, Z-III-2a, Z-III-3a) and the more advanced evaporite stages (Z-III-lb, Z-llI-2b, Z-IlI-3b). Samples from the three halite stages have a simple salt chemistry, although several percent of carnallite, bischofite, and kieserite may be present, t~37C1 generally decreases within each halite stage in the upward direction. Samples from the three advanced evaporite stages have high carnalite, bischofite, or kieserite contents. Their 637C1 is generally at the lower end of

5173 Z.lll.3b ~e~

Z.lll.3a

-1640 ,.

•

-1660 m

-168o

£

halite

~1

Z.lll.2a

-1700

/

-1720

~

-1740

__

--

,,

~

•

',.;. i

Z.lll.lb

N

.

- 1760

~

--

other'salt~

I ~ •

~

minerL~=::~

-1780

0

20

40

60

80

Salt composition

',

Z.IIl.la

100

-0,50

-0.25

0.00

0.25

,5 37CI

FIG. 3. Left, stratigraphic column of drillhole TR-2 indicating the average mineral composition of the salt minerals per interval. Right, measured 637C1 values of samples taken from the indicated depths. As can be seen 637C1values in halite rich layers are generally higher than 637C1 values in carnallite and bischofite layers. Other salt minerals indicate mainly sylvite and langbeinite.

the scale. The advanced stage Z-III-lb shows a complex pattern with 637C1 variations that seem to be partly related to the salt chemistry; increases in 6~7C1 correlate with increases in halite content. The two bischofite-rich samples in Z-III-lb have relatively high 637C1 values, despite the fact that they are the most evaporated samples in the section. Because subformation Z-III-2 (upper half of the core, Coeleweij et al., 1978) represents only one single evaporation cycle, we used this section to evaluate our Rayleigh crystallization model. The subformation consists of two distinctive parts. The lower one (Z-III-2a) contains mainly halite, whereas the upper part (Z-III-2b) represents the more advanced stages of evaporation and has a complex salt chemistry. As predicted by the Rayleigh crystallization model, 637Cl decreases gradually as evaporation proceeds. Subformation Z-III-3 is represented by only two samples, one at the onset of Z-III-3a and one at the onset of Z-III-3b. Also here 637C1 in the Z-III-3b sample is lower than in the Z-III-3a sample, indicating that these samples are related by an evaporation process. However, the lower sample (1638) has a 63vC1 of -0.32%0, much lower than would be expected at an early stage in the evaporation cycle. This suggests that an influx of seawater caused substantial dissolution of salt with low 637C1

5174

H . G . M . Eggenkamp, R. Kreulen, and A. F. Koster van Groos

Table 3: Mineralogical composition (as measured by Billiton, see text) and results of 637C1 measurements of the salt samples. Error in the 637C1 value is the oo_t standard deviation, n (third column) is the number of measurements. depth

(m

below surface

637CJ( %

vs. SMOC)

n

carnallite bischofit

halit

(%)

e (%)

e (%)

kieserit e (%)

±on. t

)

1628

-0.50±0.12

2

46

27

27

1638

-0.32±0.04

2

16

2

82

1648

-0.46±0.03

2

41

8

51

1658

-0.37±0.02

2

22

2

75

1668

-0.16±0.02

2

6

3

90

1678

-0.10±0.03

2

5

4

91

1688

-0.04±0.16

8

5

4

91

1692

-0.05±0.10

2

9

2

89

1710

+0.24±0.06

4

4

2

91

2

1716

-0.16±0.08

4

42

8

39

11

1718

-0.22±0.14

5

16

42

19

6

1725

-0.18±0.04

2

15

50

25

10

1736

-0.34±0.07

2

45

2

25

28

1738

-0.56±0.09

2

45

2

25

28

1740

-0.33±0.02

2

45

2

25

28

1742

-0.42±0.01

2

47

52

1

1744

-0.23±0.05

2

5

93

1

1747

-0.25±0.01

2

62

33

4

1751

-0.43±0.11

3

62

33

4

1756

-0.47±0.06

4

57

32

7

1758

-0.26±0.03

4

8

90

2

1771

-0.14±0.10

5

93

5

1775

-0.22±0.01

2

95

3

1783

-0.58±0.04

2

45

25

1786

-0.09t0.04

2

97

2

values, so that the chlorine isotope composition of the brine became relatively light. The composition of the two Z-III-3 samples, which contain substantial carnallite contents, supports this model. The lower half of the core was taken from the Zechstein m 1 subformation which is much more complicated. According to Coeleweij et al. (1978), this part of the core represents nine evaporation cycles with several subcycles. Seventeen samples were taken from this section. Obviously, the sampling cannot be completely representative of this section. Some of the samples may reflect processes not related to this evaporation. For example, consider the two lowest samples from Z-III-1 a. Sample 1789 is nearly pure halite and has a 637C1 o f - 0 . 0 9 % t ~ , which is quite normal. Sample 1783, however, has the most negative 637C1 found in this study. It contains abundant langbeinite and sylvite, which is indicative of secondary processes (Braltsch, 1962). Such processes might have changed the isotopic composition. The overlying Zechstein m - l b is an alternation of layers dominated by halite, camallite, and bischofite. In this section, 637C1 is

generally low, as expected in the advanced stages of evaporation. In this subformation, halite-rich layers tend to have a higher 637C1 than the adjacent carnallite-rich layers, which is compatible with an influx of seawater. In general, however, ¢537C1 values of the halite rich samples remain relatively low, compared to halites precipitated from unaltered seawater, suggesting that substantial amounts of the earlier precipitated salts were dissolved when the fresh seawater entered the basin. The uppermost Z-III-lb sample (1710) is almost pure halite with a very high 637C1, +0.24%0. This sample must have been crystallized from an influx of fresh seawater inflow and represents almost the first precipitate. Particularly interesting are the two bischofite-rich samples in subformation Z-Ill-lb. These samples have a relatively high t537C1 compared to other samples from similarly advanced evaporation stages. This suggests that bischofite indeed has a fractionation factor that is lower than unity as indicated by the experimental data, and that 63vC1 values of evaporites can increase during the last stages of evaporite formation. 6, CONCLUSIONS The 637C1 results on Zechstein evaporites and the 637C1 behaviour predicted by a Rayleigh fractionation model are in reasonable agreement. Crucial to our 637C1 model are the slightly different, experimentally determined fractionation factors for the different salts. These results agree with a study by Raab and Spiro ( 1991 ) who showed that sulphur isotopes in marine sulphate evaporites have various fractionation factors for the different salt minerals. Fractionation for chlorine isotopes is, however, much less that for sulphur isotopes. Striking resemblances between this study and that from Raab and Spiro ( 1991 ) are the positive 103 In a in the first evaporite stages and the negative 103 In a in the last evaporite stages. An implication of the changes in isotope fractionation factor is that highly negative 6 37C1 values cannot be generated in the final stages of evaporite formation. This finding contradicts predictions made in earlier chlorine isotope studies of evaporites. In these earlier studies, highly negative values have been looked for but were never found. The systematic changes in ~)37C1, especially during the halite stage, are a useful tool in the study of evaporites. They can be used to monitor the amount of evaporation that has taken place, detect periods of input of new seawater, and give information on mixing with partly redissoived salt.

Acknowledgments--We thank Billiton Refractories B.V., and especially ing. H. Lorenzen and ing. H. P. Rogaar for providing the samples and information about the samples. A. E. van Dijk is thanked for analytical assistance and additional measurements. Prof. A. G. Hermann and Dr. S. O. Scholten carefully read earlier versions of the manuscript and suggested many improvements. Miss D. C. McCartney is thanked for linguistic advice. Constructive comments by the reviewers, Drs. J. H o r i t a a n d J. Morris, are highly acknowledged. The mass spectrometer was partly financed by the Netherlands O r g a n i z a t i o n for the Advancement of Science ( N W O ) . This research is part of the Netherlands Foundation for Earth Science Research ( A W O N ) project Geochemistry o f c h l o r i n e i s o t o p e s ( # 7 5 1 . 3 5 5 . 0 1 4 ) with financial aid from N W O . Editorial handling: J. D . M o r r i s REFERENCES Ault W . U . and Kulp J. L. ( 1 9 5 9 ) Isotopic geochemistry of sulphur. Geochim. Cosmochim. Acta 16, 2 0 1 - 2 3 5 .

Isotope fractionation of CI in evaporites Braitsch O. ( 1962 ) Entstehung und Stoffbestand der Salzlagersttitten. Springer Verlag. Buyze D. and Lorenzen H. (1986) Solution mining of multi-component magnesium-bearing salts: a realization in the Netherlands. CIM Bull 79, 52-60. Claypool G. W., Holser W. T., Kaplan I. R., Sakai H., and Zak I. (1980) The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chem. Geol. 28, 199-260. Coeleweij P. A. J., Haug G. M. W., and van Kuijk H. (1978) Magnesium-salt exploration in the Northeastern Netherlands. Geol. Mijnb. 57, 487-502. Eggenkamp H. G. M., Kohnen M. E. L., and Kreulen R. (1994) Analytical procedures for 637C1 measurements. In Eggenkamp H. G. M. ~37C1;the geochemistry of chlorine isotopes. Ph.D. Thesis, Utrecht. Geol. Utrai. 116 13-21. Eriksson E. (1963) The yearly circulation of sulfur in nature. J. Geophys. Res. 68, 4001-4008. Haug G. (1982) Report on the results of drillhole TR-2. Internal Shell Report. Hoering T. C. and Parker P. L. ( 1961 ) The geochemistry of the stable isotopes of chlorine. Geochim. Cosmochim. Acta 23, 186-199. Holser W. T. and Kaplan I. R. (1966) Isotope geochemistry of sedimentary sulfates. Chem. Geol. 1, 93-135. Kaufmann R. S. (1984) Chlorine in groundwater: stable isotope distribution. Ph.D. thesis. Univ. Arizona. Kaufmann R. S., Long A., Bentley H., and Davis S. (1984) Natural chlorine isotope variations. Nature 309, 338-340. Kaufmann R. S., Long A., and Campbell D. J. (1988) Chlorine isotope distribution in formation waters, Texas and Louisiana. AAPG Bull. 72, 839-844. Long A., Eastoe C. J., Kaufmann R. S., Martin J. G., Wirt L., and Finley J. B. (1993) High-precision measurement of chlorine

5175

stable isotope ratios. Geochim. Cosmochim. Acta 57, 29072912. McIlvaine T. C. ( 1921 ) A buffer solution for colorimetric comparison. J. Biol. Chem. 49, 183-186. McIntire W. L. ( 1963 ) Trace element partition coefficients: A review of theory and applications to geology. Geochim. Cosmochim. Acta 27, 1209-1264. Nederlandse Aardolie Maatschappij B. V., and Rijks Geologische Dienst (1980) Stratigraphic nomenclature of the Netherlands; Verh. Kon. Ned. Geol. Mijnb. Gen. 32. Nielsen H. (1966) Schwefelisotope im marinen Kreislauf und das 634S der fritheren. Meere. Geol. Rundschau 55, 160-172. Nielsen H. and Ricke W. (1964) Schwefel-Isotopenverhaltnisse von Evaporiten aus Deutschland; Ein Beitrag zur Kenntnis von 634S im Meerwasser-Sulfat. Geochim. Cosmochim. Acta 28, 577-591. Raab M. and Spiro B. ( 1991 ) Sulphur isotope variations during seawater evaporation with fractional crystallization. Chem. Geol. ( lsot. Geosci. Sect.) 86, 323-333. Rees C. E. (1970) The sulphur isotope balance of the ocean: an improved model. Earth Planet. Sci. Lett. 7, 366-370. Taylor J. W. and Grimsrud E. P. (1969) Chlorine isotopic ratios by negative ion mass spectrometry. Anal. Chem. 41, 805810. Thode H. G. and Monster J. ( 1965 ) Sulphur-isotope geochemistry of petroleum, evaporites and ancient seas. AAPG Bull. Mere. 4, 367377. Vengosh A., Chivas A. R., and McCulloch M. T. (1989) Direct determination of boron and chlorine isotopic compositions in geological materials by negative thermal-ionization mass spectrometry. Chem. Geol. ( Isot. Geosci. Sect.) 79, 333-343.