Oxygen and hydrogen isotope geochemistry of zeolites

Gewhimica ef Comsxhimica A& Vol. 54, pp. 1369-I 386

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Oxygen and hydrogen isotope geochemistry of zeolites HARALDURR. KARLSSON’**and ROBERTN. CLAYTON* ‘Department of the Geophysical Sciences, University of Chicago, Chicago, IL 60637, USA ‘Department of the Geophysical Sciences, Department of Chemistry, and the Enrico Fermi Institute, Unive~ity of Chicago, Chicago, IL 60637, USA (Received Febrzlary6, 1989; accepted in r~vi~~d~r~ February 16, 1990)

Abstract-Oxygen and hydrogen isotope ratios were obtained for natural samples of the zeolites, analcime, chabazite, clinoptilolite, laumontite, mordenite, and natrolite. Zeolites were dehydrated by heating under high vacuum to 450°C to remove all the channel water. In most cases isotopic ratios were measured both for the zeolite framework ( 6’8Of) and the channel water ( 6’80,, 6D). Most of the data were collected on analcime-49 samples from 34 world-wide locations and a variety of geological settings and ages. Dehydrated analcime has a range in @‘Of that extends from 4.3.to 26.6%, which is unusually wide for a single mineral species. There is a rough correlation between 6”Of and the inferred formation temperature such that “igneous” analcimes have low values (8.7 to 14.3%) and sedimentary analcimes have high values ( 16.6 to 24.5%~). It is, however, impossible to prove on the basis of oxygen isotope data whether an analcime is a primary igneous mineral since 6’80t values for hydrothermal analcimes (4.3 to 26.6%~) overlap with those of “igneous” analcimes. The oxygen isotope systematics of associated hydrothe~~ analcime-calcite pairs and two well-constrained samples from Surtsey Island and DSDP site 417A suggest that the analcime-water fractionation is similar to that of calcitewater. Fewer data were obtained for the other zeolites. High to intermediate 6180f values for five chabazite (20.6 to 25.8%), two soil clinoptilolites (31%), one natrolite ( 18.2%0), and one mordenite ( 17%) may indicate low-temperature origin for these samples. Laumontite (6%), however, may have formed at higher temperatures than the above zeolites. The oxygen isotope relationship between laumontite and calcite suggests that the laumontite has not preserved its original oxygen isotope ratio and exchanges more readily than calcite. Oxygen and hydrogen isotopic ratios for the channel waters indicate that the channel waters in analcime and natrolite are related to meteoric waters at the sample site. Analcime channel waters ( 6180,: -5 to -24%; 6D: -49 to -167%) fall near the meteoric water line and vary systematically with the latitude of the sample site, suggesting that the isotopic signature was inherited from the sampie locality. Natrolite channel water may also reflect that of the sample site. In contrast, isotopic com~sition of chabazite, clinoDtilolite. laumontite. and mordenite channel waters are derived from the ambient water vapor where the zktlites were last stored. INTRODUCTION

Due to experimental difficulties, little work has been done on the stable isotope geochemistry of zeolites. This study was undertaken in order to address the following questions:

ZEOLITESARE WIDESPREADminerals and occur in a large variety of rock types. More than 40 naturally occurring species have been described (e.g., BRECK, 1974; GOTTARDI and GALLI, 1985 ) . Zeolites are hydrous aluminosilicates which possess unique properties such as reversible dehydration, selective sorption, molecular sieving, ion exchange, and catalysis. While we know a great deal about their occurrence, chemistry, crystal structure, and physical properties, we know relatively little about the petrogenesis of zeolites, their role in water-rock interaction, or the structure and properties of the zeolite channel water. Stable isotope research on zeolites can yield two kinds of information. Oxygen isotope ratios of the zeolite framework may be used to constrain the conditions of zeolite formation such as temperature, fluid composition, and the nature of precursor materials. Oxygen and hydrogen isotope ratios of zeolite channel water may be used to determine the origin of the water and to study the process of dehydration.

Is it possible to obtain reproducible and accurate oxygen isotope ratios of a zeolite? Might zeolites be useful as low-temperature geothermometers? What is the origin of the zeolitic water? Is it simply the last water vapor to which the sample was exposed, or does it reflect some past atmosphe~c or fluid composition? Can oxygen isotope analysis be used to identify primary igneous analcime? PREVIOUS WORK To date, few oxygen and hydrogen isotopic measurements have been reported for natural zeolites. Oxygen isotope ratios of zeolites were first reported by SAVIN( 1967) who measured the oxygen isotope ratios of three deep-sea phillipsites but had some difficulties because of inadequate dehydration. Later SAVIN ( 1973) also reported oxygen isotope values for a deep-sea clinoptilolite. The high 6 “0 values of these deepsea zeolites indicate that they are authigenic, ruling out the possibility that they inherited oxygen from precursor materials

* Present address: PIanetary Science Branch, Code SN2, NASA Johnson Space Center, Houston. TX 77058, USA. 1369

H. R. Karlsson and R. N. Clayton

I370

such as volcanic glass or silicate minerals. without extensive isotopic exchange with the aqueous environment (SAWN, 1967: SAVIN and EPSTEIN, 1970a,b). MATTHEWS ( 1980) determined the oxygen isotope ratios of one analcime from Golden, Colorado, as a part of his study of the kinetics of the reaction: analcime + quartz = albite + H20. Matthews gave no details of the treatment of analcime prior to extraction of oxygen. KARL~~~N et al. ( 1985 ) measured fiJ80 f values for seven analcimes from various geologic environments and reported a wide range from 8.8%0 to 24.1 %o, thus demonstrating the feasibility of analcime in geothermometry. Oxygen isotopic compositions of zeolites have occasionally been reported from the Deep Sea Drilling Project (DSDP). These include two analcime d “0 values from sites 504B ( HONNOREZ et al., 1983) and 4 17A ( STAUDIGEL et al., 198 1), a single 6 IRO value for a phillipsite from site 396B ( BOHLKE et al., 1984). and hve 6’*0 values for three zeolite species (analcime, clinoptilolite, heulandite ) of pore-filling cements at site 445 (LEE, 1987). LAMBERT et al. ( 1988) reported a 6’*0 value and some Rb and Sr isotopic ratios for a vug natrolite from 42 Ma old alkali basalt in southern Texas. STALLARD and BOLES ( 1986. 1989 ) measured 6180 values for laumontite and stilbite from some veins in the marine volcanogenic sediments of the Hokonui Hills, New Zealand. LUHR and KYSER ( 1989) determined framework ( 6”Of) and the channel water ( 6180,,, 6D) values for six analcimes from various occurrences. They obtained a similar range in 6180r as KARLSSON et al. ( 1985). There are several problems with the data discussed above. First, details of the dehydration conditions and estimates of the analytical precision were given only by SAVIN ( 1967 ). LUHR and KYSER ( 1989). and STALLARD and BOLES( 1989). Second, and most importantly, none report analytical accuracy. It therefore remains to be shown whether these data yield unperturbed oxygen isotope ratios of the zeolite framework. MATERIALS

AND METHODS

Samples

Zeolites were chosen for the present work on the basis of I ) abundance, 2) high purity and ease of separation from host rock, and 3) structural stability during dehydration. All three requirements were met by the zeolites studied, except for laumontite and natrolite, which do not satisfy the last criterion. However, since these zeolites are both geologically extremely important, they were also included. Large hydrothermal zeolites were favored over line-grained sedimentary zeolites and the study is biased towards hydrothermal samples. Zeolites and associated phases were studied in hand specimen and under a binocular microscope. After the minerals had been tentatively identified and textural relationships among them noted, monomineralic separates were handpicked. Individual grains varied from a tenth of a millimeter to two centimeters. Visible impurities were removed by scraping or chiseling. On large crystals, surface impurities were removed by a diamond dental drill. Mineral separates were crushed lightly and the purest chips picked out for identification and further work. These were ground in an agate mortar and checked optically for impurities, such as foreign minerals or inclusions. Most samples had less than 5 ~01%impurities. A brief description of samples and their locations can be found in the Appendix. With the exceptions of SAY- 104 (analcime ) , Clino-BCk (clinoptilolite ) , and Clino-CBk2 ( clinoptilolite) , zeolite separates received no further treatment prior to isotope analysis. Samples SAY-104, Clino-BCk, and Clino-CBk2 were further purified by the use of heavy

liquids. Because the liquids contain no oxygen it is unlikely that this affected framework oxygen isotopic compositions. Sample SAY-104 was separated with CHJ, (ME1 mixture; J. F. LUHR, pers. comm., 1984) and samples Clino-BCk and Clino-CBk2 with a mixture of stetrabromoethane and bromobenzene (MING and DIXON, 1987). None of the samples was acid cleaned. since many zeolite structures either break down or lose aluminum during acid treatment ( BRECK, 1974). Minerals were identified by X-ray diffraction and electron microprobe analysis. Chemical analyses of most minerals were done with an Applied Research Laboratories ( ARL) electron microprobe using the energy dispersive system (EDS ). During a typical zeolite analysis, counts were collected for the elements Si. Al, Na, K, and Ca for 2 min using a beam current of SO nA and a 20 to 30 pm spot size. Simultaneously, loss of sodium and potassium was monitored by observing count rates on crystal spectrometers. Natural feldspars and synthetic glasses with feldspar compositions were used as standards, Oxygen and Hydrogen Isotope Analysis Oxygen was extracted from zeolites, quartz, and datolite using the BrF, method ofCLAYToN and MAYEDA( 1963). As will bediscussed later, all zeolites were already dehydrated prior to loading into the BrF, reaction vessels. Reaction vessels were outgassed at 275 to 375°C for 90 to I50 min for dehydrated zeolites and silicates, but at room temperature for hydrous analcime. All samples were reacted at 500 to 650°C for at least 12 h. According to SVEINBJ~RNS~~TTIR et al. ( 1986) analcime can react with BrFS even at room temperature. Oxygen yields are reported as Nmol O2 per mg dehydrated zeolite. Carbonates were reacted overnight with 100% phosphoric acid at 25.3”C to yield COz following the method of MCCREA( 1950). Waters (0.5 to I I mg) extracted from zeolites were either reacted with BrFS at 300°C to liberate oxygen using the method of O’NEIL and EPSTEIN( 1966) or equilibrated with small amounts of COz according to the method of KISHIMAand SAKAI (1980). With the former method, only the oxygen isotopic composition could be obtained, but with the latter. both oxygen and hydrogen isotopic compositions were obtained on the same water sample. Water was converted to hydrogen by reaction with uranium turnings at 700°C ( BIC~ELEISEN et al., 1952). The hydrogen formed during the reaction was collected by adsorption onto activated charcoal at liquid nitrogen temperature. Gas samples were analyzed isotopically using isotope ratio mass spectrometers in two different laboratories: at the University of Chicago and at the US Geological Survey in Menlo Park. The stable isotope analyses are reported in the conventional b-notation (see CRAIG, 1957). d I80 and 6D values are relative to SMOW, whereas 6 “C values are relative to PDB. The following fractionation factors (a) have been used for oxygen: 1.0407 ( ~Y(COZ-H~O))and 1.01025 (a(acid)). where a(acid) is the fractionation factor in the reaction of calcium carbonate with phosphoric acid at 25.3”C. Dehydration of Zeolites

Meaningful oxygen isotope ratios of a zeolite can be obtained only if the framework oxygens and channel water oxygens are analyzed separately ( KARLSSON, 1988). The separation must be done in such a way that it does not cause any significant oxygen isotope exchange between the framework (Or) and the channel water (0,). There are three ways in which separation might be accomplished: I ) dehydration of the zeolite; 2) preferential fluorination (see e.g., HAMZA and EPSTEIN,1980; HAIMSONand KNAUTH,1983; KARLSSON,1988); and 3) cation exchange (see KARLSSON, 1988). KARLSSON (1988) showed that dehydration is the only viable method. The dehydration must be done outside the fluorination line since the inside walls of the Ni-reaction tubes are coated with extremely hygroscopic fluoride compounds which readily adsorb water (GARLIC& 1965; SAVIN, 1967: YAPP, 1987) The channel water cannot be completely removed from zeolites by drying at room temperature. Thus zeolites cannot be dried by storage in a drybox with an atmosphere dried over PZOS prior to loading, as is commonly done with clays (SAVIN, 1967). KARLSSON( 1988) and KARLSON and CLAYTON (in prep.) demonstrated that analcime does not dehydrate significantly under high vacuum at temperatures up to 150°C. Hence it was possible to analyze

Isotope geochemistry of zeolites the total oxygen isotopecomposition of analcime (6%,) and compare it with the oxygen isotope commotion of dehyd~ted anafcime (&‘80r). They found that hy~at~ anatcime was 2 to 5.6L iower in S’*Othan analcime dehy~t~ at 450°C. Dehydration of zeolites was carried out under vacuum (N IO-’ torr) in two stages: room temperature evacuation and heating. A series of runs were conducted to determine the most suitable dehydration conditions for the zeolites analcime, chabazite, laumontite, and mordenite. Although a wealth of data exists on the dehydration of these zeolites, these experiments were deemed necessaty since most of the preexisting dehydmtions were done in air and rarely under vacuum (VAN REEUWUK,1974). Thus the effects of dehydration under high vacuum were unknown. The “best” dehydration conditions are considered to be those that allow complete dehydration at the lowest possible temperature and in a reasonable time (days or hours). Typically, a powdered sample (lo-80 mg) was evacuated at room tern~mtu~ for 1 to 5 h and then heated at 2 to 4”C/min to 45O”C, at which tem~mtu~ dehydration was complete. A detailed

description of the dehydration and channel water collection procedures is given by KARLSON ( 1988) and KARLSSONand CLAYTON (in prep.).

Reproducibility of Oxygen Isotope Ratios for Zeolites KARJSSON (1988) found that the oxygen isotope ratios for dehydrated analcime, chabazite, clinopti~olite, laumontite, mordenite, and natrohte can be reproduced with a precision which is similar to that of NBS-28 (9.00 It 0.20% ( 1a)). The reproducibility of the dehydration steps for analcime, chabazite, laumontite, and mordenite was studied by 1) dehydrating the same specimen for different lengths of time and at different heating rates, 2) repeated dehydration, 3) exchange with 19% water vapor at room tem~rature for two weeks, and 4) heating at 200°C in the dehydration vessel for 24 h. The effects of these treatments vary from zero to 2%. Dehydration does not at&t the framework oxygen isotope composition of analcime and mordenite markedly (cO.3%+) but has a significant effect on the 6 “Or values for chabazite and laumontite. The poor accuracy of isotopic analyses obtained for these two latter zeohtes could severely hamper their use in oxygen isotope geothermometry. For further discussion about the experimental procedure and results see KARLSSON (1988) and KARLSON and CLAYTOX(in prep.).

RESULTS AND DBCXJSSION

Oxygen and hydrogen isotopic analyses were obtained on the zeolites analcime, chabazite, clinoptilolite, laumontite, mordenite, and natrolite. Zeohte samples were classified into three different groups on the basis of their occurrence: 1) sedimentary or S-type, 2) hydrothermal or H-type, and 3) igneous or I-type. lndividual assignments were based on descriptions and localities of samples (see Appendix). The results are discussed below for individual species. Analcime

Most of the present results are for analcime (Na[AISiz06] *H20). Analcime is among the most common zeolites, occurs in a broader range of geological environment than any other zeolite, and has the widest range of inferred formation temperatures (5 to 65O’C) of any zeolite (see, e.g., GOTTARDI and GALLI, 1985 ). Since oxygen isotope fractionation depends both on temperature and on fluid composition, analcime can be expected to show a large variation in oxygen isotope ratios. Analcime has the narrowest channel diameter (2.6 A) and thus the most tightly bonded water of any zeolite. It is therefore likely that channel water in analcime is less susceptible to isotopic exchange with the surroundings

1371

than channel water of other zeolites. The analcime structure contains the smallest amount of water (8 to 9 wt’%) and the structure remains intact during dehydmtion. Hence the chances of isotopic exchange between the zeolitic framework and the channel water should be relatively small when compared with more hydrous zeolites, which undergo major structural changes during dehydration. Roughly 50 samples from 34 different locations throughout the world were analyzed isotopicaIly in order to establish the range in 6 isOr values for natural analcime. Ages of the analtimes vary from as young as -20 a for Surtsey, Iceland, to perhaps as old as 1 billion a for Keweenaw, Upper Michigan. S-type, H-type, and I-type correspond to the low-tempemture (deep-sea sediments, soiis, and alkaline lakes: -O-5O”C), inte~ediate-tem~rature (geothermal systems: +o35O”C), and high-temperature environments (magma: -600~650°C), respectively. The results for analcime framework and channel water are listed in Tables 1 and 2, respectively, and are discussed below.

Framework oxygen The range in 6 **Ovalues obtained for dehydrated analcime extends from 4.3 to 26.6%~. This is an unusually wide range in 6 “0 values for a single species. Only minerals like calcite, quartz, and feldspar-s have similar ranges. Typically, S’% values in these minerals range from 5 to 35%~(FAURE, 1986; HOEFS, 1987). Ranges of oxygen isotopic composition of S-, H-, and I-type analcime are shown in Fig. 1. It is interesting to note the distribution of 6’*Or values among these types. There appears to be a rough correlation between the 6’*Or values and the inferred formation temperatures. Thus high 6 ‘*Or values are observed for the S-type or low-temperature occurrences while low 6 ‘*Or values are observed for the ftype or high-temperature occurrences. Hydrothermai analtimes, which have intermediate temperature ranges between those of S and K-types, have 6 180 f values that span the entire range from Iow to high. It is clearly possible to distinguish most ~d~rnent~ analcimes from igneous analcimes on the basis of oxygen isotope ratios. As is discussed in more detail by KARLSSON and CLAYTON ( 1990a), oxygen isotope ratios can aid in determining the petrogenesis of igneous analcime (e.g., whether it is primary or secondary) but only to the extent of showing when it is not primary. A low &‘*Or value does not necessarily imply igneous origin since b’*Or values for I- and H-type analcimes overlap. Sedimentary (S-type). The 6 ‘*Or values for six analcimes are shown in Table I and Fig. I. These samples were all from Tertiary saline-alkaline lake deposits: one sample from the Barstow Formation, California (Pliocene); another from the Green River Formation (Laney Shate Member). Wyoming (Eocene); three from the Big Sandy Formation. Arizona (Miocene); and one from an unnamed lacustrine deposit roughly 30 km east of the Big Sandy Formation at Aquarius Cliffs, Arizona (late Cenozoic). The 6’*Or values for these analcimes range from 16.6 to 24.5%. Green River analcime has a distinctly lower isotopic composition, 16.6%, than the remaining analcimes, which range from 22.5 to 24.5f&,. Analcime from the Big Sandy and Aquarius Cliffs is remark-

H. K. Karlsson and R. N. Clayton

1371 Table 1.

Analcime framework: oxygen isotope results.

Sample #

M3-105 Big Sandy Green River SW-3-2 USNM-95982 USNM-126762 DSDP~l?A-38-~1 DSDP~l7A-38-2#2 FMNH-M6743 ~NH-M9~5#5 FMNH-M9257 HRK-82-57 #3-l HRK-82-51#3-3 IMNH-1556 IMNH-7522 IMNH-7522 IMNH-7522 KcwlXnaw

s s ss : H H

?!J%?47 #Al UC.1547 #AZ-l UC-1547 #AZ-Z UC-1547 #AZ-S L’SNM-3536 USNM-3541 USA-4210 USNM-7~ UStiM-8363-l USNb-11884 USNM-12792 USNM-17271

USNM-83258 USNM-115625 USNM-125977 USNM-132507 USNM-134134 USNM-137136 USNM-145340 USNM-151697 USNM-155369 USNM-B17295 USNM-C3553 USNM-RI1059 USNM-RI 1062 USE-R71~ USA-R9674 &39E SAY-104

Locality

Type*

I

Barstow, California, USA Big Sandy, Arizona, USA Green River, Wyoming, USA Big Sandy, Arizona, USA Aquarius Cliffs. Arizona, USA Wtkieup, Arizona, USA South of Bermuda Rise. N. Atlantic South of Bermuda Rise, N. Atlantic Golden, Colorado, USA Brevik, Norway Edgewater. New Jersey, USA EyrarfjalI Mt.. W. Iceland Eyrarfjall Mt., W. Iceland Oddskard, E. Iceland M&horn Mt., E. Iceland Melshorn Mt., E. Iceland Melshom Mt., E. Iceland Keweenaw. Michigan, USA Surtsey Island, S. Iceland Nova Scotia, Canada Nova Scotia, Canada Nova Scotia, Canada Nova Scotia, Canada Chechy, Czechoslovakia Giant’s Causeway, Ireland Two Islands, Nova Scotia, Canada Faeme Islands Bergen Hitl, New Jersey, USA Baltimore Tunnel. Matylaod USA Ignak, Greenland Norheim, Rhineland-PfalL Germany South Table Mt., Colorado, USA Umalc, Guam, USA Sicily, Italy Mt. St. Hilaire, Quebec. Canada Kings Mt.. N. Carolina, USA Challis, Idaho, USA Cornwall, Pennsylvania USA Prospect Park, New Jersey, USA Fliiders, Victoria, Australia Antrim Ireland Great Notch, New Jersey, USA Springfield, Oregon, USA Homi ZaIezky,Czechoslovakia Arendal, Norway Old Kirkpatrick, Scotland. UK Crowsnest, Alberta Canada Crowsnest, Alberta, Canada Coliia, Mexico

Grain Size**

9 Yield

(mesh)

(&mol/mg) +10* n

100-200

100-200 ZOO-325 >325

14.6 14.4 14.5 14.4 14.5 12.7 14.7 15.5 14.8 14.7 14.2 14.8 14.7 14.5 14.7 14.6 14.7 14.4 14.x 14.8 14.6 14.6 14.0 14.4 14.5 11.7 14.0 14.0 12.6 14.9 14.4 14.7 14.9 14.9 14.5 13.8 14.4 14.8 14.3 14.4 14.8 14.6 14.4 14.3 14.5 14.4 14.1 14.4 14.3

0.06 0.12 0.47 0.30 0.15 0.29 0.81 0.32 0.08 0.10 0.10 0.10 0.10 0.07 0.19 2.43 0.32 0.30 0.29 0.07 0.08 0.16 0.19 0.12 0.08 0.10 0.21 0.10 0.08 0.04 0.08 0.15 0.12 0.18 0.08 0.30 0.12 0.09 0.10 0.14 0.09 0.51 0.11 0.09 0.11 0.17 0.09 0.28 0.12

6’80

2 3 2 2 1 1 2 1 1 1 I 1 1 1 1 3 10 2 3 1 1 2 1 1 1 1 1 1 1 1 1 1 1 1 1 2 1 1 1 1 1 2 1 1 1 1 1 3 3

(%a) *1ott 22.5 23.3 16.6 24.2 24.3 24.5 24.1 24.2 15.5 22.2 23.2 9.2 9.5 11.3 12.1 12.2 11.9 21.2 12.2 16.0 15.6 15.4 15.6 20.1

nS,

0.04 0.19 0.16 0.11

2 3 2 2 1

0.11 0.18

: 2 1 1 I

0.13 0.20 0.02 0.19 0.08 0.10

; 1 1 3 20 2 3 1 1 2 1 2

18.4 19.0 20.2

: 1

26.6 20.0 7.2 15.2 18.0 22.9 17.0 13.7 18.7 0.25 4.3 21.1 25.2 21.4 19.8 24.5 0.45

: 1 1 1 1 1 1 2 1 1

18.5 22.4 17.7 14.3 13.6 8.7

f 2 2 1 3 4

0.31 0.05 0.06 0.13

t :.

*Refers to the oczurrcnce type. I: igneous; H: hydrothermal: S: sedimentary. **For those samples that were not sieved the grain size is not given. The grain size of these samples was -200 to 325 mesh. tThe precision i. e., standard deviation, was obtained in two different ways: where more than one determination was made it was calculated according to conventional statistical formulas. otherwise it was estimated by propagation of errors through the formula used for computing the yield. t&lculated

precision, i.e., standard deviation.

§Numbcr of separate extractions.

ably constant in isotopic com~sition (23.3 to 24.5%~). The isotopic data suggest that the formation conditions for anakime in the Green River Formation differed from those in the Barstow and Big Sandy formations, as will be discussed further below. Hydrothermal (H-type). Oxygen isotope data for 40 Htype analcimes are compiled in Table I and displayed in Fig. 1. The 6 ‘*Or values span the range from 4.3 to 26.6%~ The three lowest 6 “0 r values were obtained for analcimes from the following localities: Challis, Idaho (USNM- 137 136: 4.3%0), Iganak, Greenland (USNM-12792: 7.2%), and Eyrarfjall Mt., western Iceland (HRK-82-57: 9.2 and 9.5%;0). At Chalhs and Eyrartjall, tieid evidence suggests elevated temperatures at some time during or after the formation of analtime. Ross and SHANNON ( 1925 ) described the Challis an-

alcime as flattened crystals occupying cavities between calcite lameliae in vesicles of an andesite lava flow. According to their interpretation the calcite had been dissolved away and the cavities filled with analcime. Dissolution of the cakite could have resulted from a rise in temperature. However, it is also possible that the calcite dissolved because pH or pCO:! of the fluid changed. The highest ii ‘*Or values were found in analcime from New Jersey and in DSDP core 4 17A. Four New Jersey analcimes-Edgewater, Prospect Park, Great Notch, and Bergen Hill-range from 23.2 to 26.66, while the two DSDP analcimes are 24.1 and 24.2%~ Analcimes from four different localities in Iceland have low 6 “0 r values ranging from 9.26 at Eyrarfjjl to 12.2%~ for Surtsey. The formation conditions of the Surtsey sample are the best constrained of ah the analcimes studied. This

Isotope geochemistry of zeolites Table2.

An&me channelwaler: oxygenendhydrogenisotoperesults.

Sample #

Type’

Locality

Grain size** (M-h)

M3-105

Big Sandy Green River SW-3-2 USNM-95982 USNM-126762 DSDP417A 38-2#1 DSDP417A 38-2#2 FMNH-M6743 FMNH-M9255#5 FMNH-M9257 HRK-82-57 #3-l HRK-82-57 #3-3 IMNH-1556 IMNH-7522 IMNH-7522 IMNH-7522 KeweP,ll.¶V/ SurtseV UC-1547 #Al UC-1547 #A2-2 USNM-3536 USNM-3541 usNM-4210 USNM-7099 USNM-8363-1 USNM-11884 USNM-12792 USNM-17271 USNM-83258 USNM-115625 USNM-125977 USNM-132507 USNM-134134 USNM-137136 USNM-145340 USNM-151697 USNM-155369 USNM-B17295 USNMX3553

s

: : si H

USNM-R11059

USNM-RI1062 USNM-R7100 USNM-R9674 %39E SAY-104

1373

I

100-200 Barstow. California, USA Big Sandy, Arizona. USA Green River. Nevada, USA Big Sandy. Ariiona, USA Aquarius Cliffs, Arizona, USA Wiieup. Arizona. USA Southof Bermuda Rise,N. Atlantic Southof Bermuda Rise, N. Atlantic Golden, Colorado, USA Brevik. Norway Edgewater. New Jersey, USA Eyrarfjall Mt., W. Iceland Eyrarfjall Mt.. W. Iceland oddskard, E. Iceland loo-200 Melshorn Mt., E. Iceland Melshom Mt.. E. Iceland Melshorn Mt., E. Iceland Keweenaw, Michigan, USA Surtsey Island, S. Iceland Nova Smtia. Canada Nova Scotia, Canada Chechy, Czezhoslovakia Giant’s Causeway, Ireland Two Islands. Nova Scotia. Canada Faeroe Islands Bergen Hill, New Jersey, USA Baltimore Tunnel, Maryland USA Ignak, Greenland Norheim. Rhineland-Pfalz Germany South Table Mt., Colorad& USA Urn&. Guam, USA Sicily. Italy Mt. St. Hilaire, Quebec. Canada Kings Mt., N. Carolina, USA Challis, Idaho, USA Cornwall, Pennsylvania, USA Prospect Park, New Jersey, USA Flinders. Victoria. Australia Antrim, Ireland Great Notch, New Jersey, USA Springfield, Oregon. USA Homi Zalezky.Czcchoslovakia Arendal, Norway Old Kirkpatrick. Scotland. UK Crowsnest Alberta, Canada Crowsnest, Alhertq Canada Colim& Mexico

6D

6’80 BrFSt -9.0 -8.9 -12.9 -14.0

MCEW flos

MCE

-11.6

1.25

-102.2

-13.0 -11.8 -9.9 -9.5

1.16 0.73 0.85 1.39

-105.9 -97.5 -89.1 -48.8

-13.5

0.45

-119.0

-16.8

0.90

-103.6

-14.1

0.93

-100.5

-15.7

0.69

-124.7

-10.2

0.99

-78.5

-10.5

1.42

-120.7

-22.6 -12.3 -16.7 -9.2

0.73 0.61 0.79 0.80

-153.4 -94.4 -141.0 -59.5

-7.4 -12.1 -11.3 -14.3 -14.1 -15.9 -15.1 -18.2 -4.9 -17.4 -16.8 -13.9 -8.0 -11.2 -22.7

-4.8 -19.6 -9.4

0.6

-75.8

-14.2

0.96

-83.7

-10.2 -13.1 -13.9 -15.4

0.81 0.05 2.11 0.6

-66.6 -112.0 -92.9 -106.7

-14.7 -15.1

1.23 1.15

-116.8 -89.0

-23.6

0.84

-166.6 -107.6

-18.5 -11.8

1;;:; -13.1

*Refers to the occurrence type. I: igneous; H: hydrothermal; S: sedimentary. **For those samples that were not sieved the grain size is not given. The grain six of these samples was -200 10325 mesh. tOxygen extracted with BrFS method. The analytical precision (la) is fo.3%. ttMCE stands for the Micro-CQ-water

Equilibration (KISHIMAand SAKAI. 1980), which was the method used for obtaining

both oxygen and hydrogen isotope ratios from the same water sample. %Xxlated srandard deviation for the MCE method by error propagation.

analcime, which is probably the youngest naturally occurring analcime studied thus far, is important because it can yield an estimate of the analcime-water oxygen isotope fractionation factor. Surtsey, a volcanic island south of Iceland, was formed during a submarine eruption between 1963 and 1967. In the summer of 1979 a core was drilled through Surtsey in order to study the stratigraphy and hydrothermal alteration of the island (JAKOBSSON and MOORE, 1986). The core, which is 18 1 m long and extends 123 m below sea level, records the history of 12 years of alteration of a basaltic tephra. At least ten different hydrothermal minerals formed as a result of the alteration. Analcime is among the three most common ( JAKOBS~~N and MOORE, 1986). The analcime studied here was collected from a vein that occurred at 90.1 m depth, where a temperature of 145 & 5’C was measured. In order to estimate the oxygen isotope fractionation factor it is necessary to know the composition of the fluid from which the

mineral crystallized and the temperature at the time of crystallization. Since the vein occurred below sea level the fluid can be assumed to be seawater and have a &I80 value of 0.0%. This is supported by hydrogen isotope measurements of steam emanating from the surface of the island, which in two cases yielded nearly pure seawater: 6D was -4.2 and -5.2% (JAKOBSSON, 1978). Although 12 years passed between the onset of the hydrothermal activity and drilling, temperatures in the drill hole have been fairly constant. JAKOBSSON and MOORE ( 1986) estimate that a 5 to 1O“C cooling may have occurred in the 10 to 110 m depth interval. This, coupled with the observation that the analcime occurs as a pseudomorph after calcite and therefore formed at some later time during the hydrothermal activity, suggests that the measured temperature, 145 + 5 “C, is a good estimate of the analcime formation temperature. Assuming that the analcime was in isotopic equilibrium with seawater and ignoring salinity

H. R, Katlsson and R. N. Clayton

5

10 s’*o w,,

15

20

the same temperature (O’NEIL et al., 1969). The implications of this resuit are discussed further in a companion paper ( KARLSSON and CLAYTON, 1990b). Igmxms (I-tvpej. Three analcime samples that occur as euhedral crystals in igneous rocks were studied. Two were phenocrysts from the famous Crowsnest Formation in Alberta, Canada (PN-39E and 2a). and one was microphenotrysts from the Colima Volcanic Complex in Mexico. The 6 180r values for these analcimes ranged from 8.7%0 for Colima to 13.6 and 14.3%0for Crowsnest (Table 1 and Fig. I )_ A detailed discussion of these samples is given by KARLSSON and CLAYTON(I~~O~). UxJTgenis~~i#p~rakes ~?~[~~st~~i~~ed ~~aIci~~-~al~,~~~ pairs. Analcime is most frequently associated with calcite, quartz, and other zeolites. In order to determine whether the partitioning of oxygen isotopes between calcite and analcime would be a useful thermometer, the oxygen isotopic compositions of associated analcime-calcite pairs were measured in 16 hydrothermal samples. Both oxygen and carbon isotope data were obtained for the calcite. Associated minerals other than calcite were also analyzed. The results are displayed in Table 3 and Fig. 2. Figure 2 is a SA - bs plot, the simplest way of comparing the isotopic composition of two phases A and B. On a SA-- ijBplot ail A -- B pairs that are in isotopic equi~ib~um at the same tem~rature will lie on a single straight line with slope - 1 and intercept A, where A r;: 6,

25

Dehydrated Analcime

FIG. I. Range of 6 ‘*OI for anakime from various sources. Ranges are shown for sedimentary, hydrothermal, and igneous analcime. Each square refers to one sample. Numbers in brackets give the number of samples analyzed. Note the wide spread for hydrothermal analdme (4 to 27%0). The figure shows that igneous analcime cannot be uniquely determined on the basis of oxygen isotopes.

effects, the oxygen isotopic fra~ionation between anafcime and water is estimated to be 12.3%0at 14YC. This fractionation is afmost identical to that of calcite, which is 12.6%~at

Table 3. Stable isotope composition of hydrothermal analcime-calcite Mineral

Sample # Dehydr. Analcime 6’80

Calcite 6’80 8'3C

f%)

("/oo)VW

DSDP-417A 3%2#1

24.1

27.5

1.9

DsDP-417~

24.2

27.4

1.9

FM~-M~57

23.2

18.6

-10.5

Keweenaw

21.1

15.9

-2.3

USNM-11884

20.0

17.1

-10.9

USNM-125977

17.0

33.5

9.6

USNM-137316

4.3

0.8

-3.5

USNM-151697

25.2

19.4

-5.6

USNM-155369

21.4

26.2

-21.7

USNM-3536

15.1

20.5

-8.4

USNM-4210

19.0

19.7

-8.2

USNM-83258

18.0

18.5

-14.1

USNM-8363#1

24.5

21.9

-13.3

USNM-31729s

19.8

21.8

0.4

USNMX3553

24.5

16.1

-5.0

USNM-R11062

18.5

23.0

-17.3

USNM-R7 100

22.4

22.5

-9.3

*Datolite. **Quartz.

pairs.

3a-2#2

Other Phases

6180 PQ)

17.7*

20.2**

Isotope geochemistry of zeolites

1375

had crystallized from a fluid with a constant oxygen isotopic ~orn~~~ion, the kind of dist~bution shown in Fig, 2 could be generated, Obviously, if the fmctionation curves for analtime and calcite are similar, then analcime-calcite not useful for thermometry.

pairs are

Channel water

0

10

20

S’“0 (Q,ow

30

Calcite

FIG, 2. 6 ‘*O values of associated hydrothermal analcimecakite pairs. Diagonal lines are drawn for integral A values (%o), where A 4 6 rEO(analcime) - S ‘*O(calcite). All but two data points fall less than 5% from the A = O%O line. Also note that samples tend to be distributed regularly around tbis line depending on the crystahiration order. Thus most squares fall above the line and most circles below.

-

liB(GREGORY and CRISS, 1986). Small A values on a dA - C&plot imply either high temperatures, since A + 0 as T -, co, or similar fractionation curves for the two phases If the twa phases have identical fm~onation curves, the data will fall on a line with slope 1 and intercept 0. As can be seen from the data in Fig. 2, the difference in oxygen isotope composition between dehydrated analcime and calcite is generally small (~5%). The difference between analcime and calcite is less than 10%0in all cases except for sample USNM- 125977 from Sicily, where the difference is almost 20%~~This sample is also unusual in terms of its carbon isotopic composition. Its calcite has a positive &13Cvalue of +9.6%0, while calcites in all the other samples have negative to slightly positive 6°C values (-2 1.6 to +1.9%0; see Table 3). Ignoring sample USNM-I 25977, the data suggest that, although the difherence in oxygen isotopic composition between an~~rne and calcite is generally small (<5%0), it might still be large enough to be a useful thermometer. Miners thermometers can be used only for mineral pairs that have been in isotopic equilibrium with one another. Unfortunately, the analcime-calcite pairs studied here were rarely cocrystallized. In most specimens it was possible to deduce the order of crystallization from petrographic textures revealed under the microscope. When the sequence of c~~lli~tion is coupled with the oxygen isotopic data for these minerals, the combined results support the conclusion, drawn earlier from the Surtsey analcime, that the fractionation curve for anaicime-water is similar to that of calcite-water. Inspection of Fig. 2 shows that, when the c~stalli~tion order is taken into account, the dist~bution of data points in this plot is not random, but regular about a line with a slope = 1 and A = 0. Thus most samples in which calcite crystallized first lie above this line while most samples in which analcime crystallized first lie below it. In other words, the mineral which crystallized first tends to have a lower 6 “0 value than the mineral which crystallized second. It is possible that most mineral pairs represent a sequence crystallized from a fluid upon cooling. Then, if calcite and analcime had similar oxygen isotope fractionation curves and

In addition to determination of the oxygen isotopic composition of dehydmt~ analcime, the oxygen and hydrogen isotopic compositions of most channel waters were determined (see Table 2). The isotopic compositions of the channel water range from -5 to -24960 for oxygen and from -49 to - 167% for hydrogen. Al~ou~ there is a slight positive correlation between framework and channel water compositions, the correlation is poor in both cases. Thus the correlation coefficients between 6 “Or and 6 ‘“0, and 6 ‘*Of and 6D are only 0.42 and 0.54, respectively. These results indicate that the isotopic composition of the channel water and the framework are unrelated, confirming the observations of KARLSSON et al. (1985) and LUHR and KYSER (1989). In fact, the oxygen isotope ratios for the channel water are within the range that is observed for global precipitation outside polar regions, which is - 1.2 to -25k (YURTSVER and GAT, 1981). On a 6D vs. S”O plot the channel waters duster around the meteoric water line (MWL: 6D = 8&‘*0 + 10; CRAIG, 196 I), suggesting that the isotopic composition of the channel water is directly related to that of meteoric waters (Fig. 3). Most of the channel waters lie within &2%06 ‘*O of the MWLline, while meteoric waters generally scatter slightly less, +l%o 6180 from this line (SHEPPARD, 1986). It should be stressed that, although the isotopic composition of the channel waters is variable, it does not simply reflect the seasonai changes in isotopic com~sition of the atmosphe~c precipitation f and hence water vapor) in Chicago which varies considerably (see, e.g,, YIJRTSEYERand GAT, 1981). No correlations were found in the isotopic composition of the channel water either with date of last exposure to the atmosphere or with the local mean temperature that day. Repeated analysis of channel water from a single sample (IMNH-7522) showed little change in 6”0, over a period

0 -50

P

z

-100

‘i

c.e

-150

FIG. 3. Relationship between S’*Oand bD in analcime channel water. The straight line denoted MWL is the Meteoric Water Line fmm CRAIG ( 196i ). ‘Chicago” (stippledcircle) refers to water vapor composition as calculated from the weighted mean composition of meteoric waters in Chicago ( YURTSEVER and GAT, 1981).

1376

H. R. Karlsson and R. N. Clayton

-30

-25

-20

-15

-10

-5

0

ia0 (#)suow ChannelWater FIG. 4. A plot of 6 “0, in analcime vs. latitude of sample locality. 13~~0, results from MCE and BrF5 methods in Table 2. Boxes are shown for those samples where the latitude of the locality was uncertain or where there was more than one sample from the same locality. The two solid lines in the figure show the latitude effect on the mean 6 ‘*O in precipitation over the Northern hemisphere (NHL) and over the North American continent (NACL) (for data see YURTSEVER and GAT, 198I ). The dashed line represents the isotopic composition of channel waters in equilibrium with NACL-waters. A dotted curve separates samples from North America from those of Western Europe. North American samples lie below this curve and Western European above. A similar relationship also exists between dD and latitude.

of one year. Hence, the channel water in analcime has an isotopic signature which is at least older than the shelf ages of samples obtained from museums (IMNH: Icelandic Museum of Natural History; USNM: US National Museum; and FMNH: The Field Museum of Natural History); otherwise all samples obtained from the same museum would display similar isotopic compositions. Table 2 shows that this is clearly not the case. In addition, samples from the same locality, such as 2a and PN-39E from the Crowsnest formation in Alberta, have similar 6 “OEw values (see Table 2 ). Since the Crowsnest samples were obtained from two different workers, it is likely that channel water isotope composition reflects that of the sample locality. Clearly, the best way to test such a hypothesis is to directly compare the isotopic composition of the channel waters with that of the meteoric waters from the same locality. Such a comparison, however, is not possible since in the majority of cases both the sample location (i.e., latitude and altitude) and the isotopic composition of the local meteoric waters are poorly constrained. Another way to evaluate this hypothesis is to check whether or not 6”0,, and 6D correlate with latitude of the sample locality, analogous to meteoric waters. It is well established that the isotopic composition of meteoric waters becomes more negative with higher latitude. Hence a similar decrease should be observed for analcime channel waters. The data for 6’*0,, in analcime are plotted in Fig. 4. Although there is some tendency for 6 “0, to decrease with latitude, the overall trend is not as apparent as that of precipitation. The two solid lines drawn in Fig. 4 show the latitude effect on d”O of annual precipitation averaged over the Northern hemisphere and the gradient over the North American continent ( YURTSEVER and GAT, 198 1). The spread in 6 “Ocw values shown in Fig. 4 is not surprising when it is considered

that these data have not been corrected for the altitude. coastal, or amount effects. All these effects lower the isotopic composition of meteoric waters ( YURTSEVER and GAI, I98 1). This may explain why all the data points but one lie to the left of the Northern Hemisphere Line (NHL). Of these effects, the altitude effect is the best known quantitatively. It depends on the local topography and generally varies between -0.15 and -0.50%0 for 61R0 and between -5 and -4.()%0 for 6D for every 100 m of elevation ( YURTSEVER and GA-I, 198 I ). The high altitude of sample SAY-104. which was located at - 1300 m above sea level in Mexico, could explain why this sample lies far left of the NHL at 6 “Oc,+ 2 - 13. I’$&. If an altitude correction of -0.5%0/ 100 m is applied to this sample, the 6”0,, at sea level is -6.6%0, which is considerably closer to the NHL than is the uncorrected value. In addition to being affected by altitude. the isotopic composition of precipitation also changes as a function of distance from the oceans. Thus, at similar latitudes, precipitation over continents has a generally lower isotopic composition than precipitation over coastal regions or islands ( YURTSEVER and GAT, 198 1). Such a relationship is also suggested in the 6 “0, values of analcime channel water shown in Fig. 4. The continental samples (open circles) tend to have lower isotopic values than coastal and island samples (stippled circles). Another interesting feature in Fig. 4 is that channel waters of analcimes from North America and Western Europe occupy different regions in the figure. Hence all the North American samples lie below the dotted boundary marked in the figure, while all the Western European samples lie above this dividing line. Although the reasons for this division are not clear, it is interesting to note that, according to the global map of 6”O distribution in precipitation, the d”O values decrease at a much slower rate over Western Europe than over North America. In places the 6 “0 values for precipitation over Western Europe remain constant with latitude (see Fig. 2 in YURTSEVER and GAT, 198 1). The similarity between the geographic distribution of 6 “0 values in precipitation and in channel water implies that the analcime channel water records an isotopic pattern similar to that of present-day precipitation. The North American samples, excluding the one from Mexico, appear to follow a smooth trend of decreasing 6 180cw values with increasing latitude (Fig. 4). Most of these samples lie to the left of the North American Continent Line (NACL). KARLSSON et al.(1988) and KARLSSON and CLAYTON (1990b) showed that the channel water in analcime is depleted in 6 “0 by -5%0 relative to external,water nearly independent of temperature between 25 to 400°C. Hence, if the North American samples are in equilibrium with local meteoric waters then they should cluster around a line which is parallel to NACL, but shifted to lower 6 lRO by - 5%0. The equilibrium line (dotted) seems to fit most of the data. The trend improves if the sample from Greenland is included. Chabazite Five chabazite samples, four hydrothermal and one sedimentary, were analyzed for oxygen and hydrogen isotopes. Sample localities and isotopic data are summarized in Table 4.

Isotope geochemistry of zeolites

Table4. Oxygenand hydrogenisotopiccompesitionof chabazite,~1~~~. Sample#

Type*

l-J=@

1377

laumontite,mordenite,and natrolite.

GrainSize**

Framework 02 yield

(mesh)

i$allite

FMNH-M6743 FMNH-M9250 IMNH-2456 IMNH-7926

S H

Durkee, Oregon, USA

150-200

@mol/mg) f lu

(%) It la?

w)

(5%

-78.7

14.8 13.8 13.3 13.6 13.0

0.2 0.4 0.2 0.5 1.0

20.6 21.6 25.8 23.9 24.7

0.0(2) 1.2(2) 0.2(4) 0.2(4) 1.2(4)

-9.6 -_;K;

0.1

31.5 30.5

O.lf2) 0.0(2)

-8.7 -9.0

H H H

Golden,Colorado,USA?? Bay of Fundy,Nova Scotia,Canada &hlid, Vestfkrlir,N.W. Iceland Mj&dalm,Vestfirdir,N.W. Iceland Houla, Texas, USA Houla, Texas, USA

14.1

2

Laumontite IMNH-6162 H

&tafjklur, E.Iceland

15.1 0.0

Mordenite IN-7556

H

Sttivarfj6rdur,E. Iceland

Natrolite m-M3340

H

Cape Split, Nova Scotia, Canada

<325

ChannelWater 8180, 6D

PO,

-11.2 -15.0

~~oo~~~lite CliIlo:CBk2

100-200

6.1 0.3(5)

-16.4 ttt

14.0

0.0

17.8

0.1(3)

-9.1

13.9

0.3

18.0

0.3(6)

-7.8t

-81.5

-19.5

*Refersto the occurrencetype. H: hydrothermal;S: sedimentary. **For those samples that were not sieved the grabt size is not given. The grain size of these samples was -200-325 mesh. %%andatddeviation calculated on the basis of the number of anaiyses given in parentheses. ~tAnaicime was also analyzed in thii sample.It had&t*O,= 15.5%and &tat&.= -13.5%

tt@ne d for the MCE method cakxdated by error propagation is 0.7%. For the remaining samples the lo is iO.3% *Average of two measurements with a standard deviation of iO.2%

Framework oxygen

Qifloptilolite

The S ‘*Of values of chabazite range from 20.6 to 25.8% (Table 4). The lowest of these was obtained for S-type chabazite from Durkee, Oregon (SD-4-2 ) . These high S ‘*Of values are compatible with low formation temperatures. The small variation observed in 6’*Of values for chabtite may suggest that these samples formed undepsimilar conditions.

Two S-type clinoptilolite samples were analyzed isotopitally. These samples were cleaned mineral separates obtained from a soil horizon in south Texas (MING and DIXON, 1987). Oxygen isotope data for the clinoptilolite are given in Table 4. Both samples have high 6 ‘*Or values of comparable magnitude, -3 I k. The high 6 IsOr value is consistent with Iowtemperature origin. In addition, the similarity of the 6’*Of values for the two samples suggests similar formation conditions.

Channel water Oxygen isotopes were measured in channel waters from all five samples. The S “0, range from -9.6 to - 1S.OL with an average of - 11.7%0.Both elements were measured on a single split water sample (SD-4-2), and yielded -9.6 for oxygen and -78.7% for hydrogen. This composition lies close to the MWL,, suggesting meteoric origin (Fig. 5). The narrow spread in 6 ‘*O, values and the mean composition are both consistent with a local source of the channel watei, i.e., ambient water vapor picked up by the samples in the Chicago laboratory. KARL~SON f 1988 ) and KARLSSON et al. ( f 989) have shown that the channel water in chabazite can be readily exchanged with ambient water vapor.

-50 Laumontite *

i $y

-100

-

CG

Relationship with associated minerals The chabazite from Golden, Colorado (FMNN-M6743 ), was of particular interest since it is associated with analcime (see Tables 1 and 2). The chabazite, which has a 6’*Or = 2 1.6% and a 6 ‘*Ocw= -11.9%, coats the analcime crystal, which has a 6”Or = 15.5%~and a 6’*0, = -13.5%~

FIG. 5. Relationship between 6“0 and 6D in chabazite, laumontite, and natrolite. MWL is the Meteoric Water Line (CRAIG, 196 I), and “Chicago” is atmospheric water vapor composition as calculated from the weighted mean isotopic composition of meteoric waters in Chicago (YURTSEVERand CAT, 1981). The close proximity of the channel waters in these zeolites to MWL suggests that they are related to meteoric waters.

H. R. Karlsson and R. N. Clayton

1378

If the clinoptilolite had formed in the soil, then it is possible to estimate the oxygen isotope fractionation between clinoptilolite and water at 10 to 35°C (D. W. MING, pets. comm., 1989). The 6180 value for the fluid in the soil is probably close to that of local precipitation. Assuming that the 6 180 value of the local precipitation is the same as the weighted mean hi80 value at Waco, Texas (YURTSEVER and CAT, 1981),wecanassigna6’80 - -4%’ to the soil fluid. Hence we get the following estimate for the clinoptilolite-water fractionation: A’8(clinoptilolite-water)

= 3 I - (-4)

= 35%0.

For comparison another estimate of A (clinoptilolite-water) can be made from 6 180r values of 27%~ and 30%0 given by SAVIN ( 1973) and LEE ( 1987), respectively, for a deep-sea clinoptilolite. Assuming seawater = 00/w,we obtain A18(clinoptilolite-water)

= 27-30%0

at “sedimentary temperatures” (O’C?) from their data. The discrepancy between the two A values is not too surprising considering the uncertainties in sample purity, fluid composition, and temperature. The agreement would be reasonable if the Texas sample had formed in seawater. Two oxygen isotope analyses were performed on the channel waters from the soil clinoptilolite. The S ‘*Ocwvalues obtained were similar for both samples ( - -8.9L) and probably represent the 6180 of the ambient water vapor that the samples were exposed to in the Chicago laboratory (Table 4). Laumontite Oxygen and hydrogen isotope ratios were obtained for one hydrothermal laumontite sample from Iceland (IMNH6 162). The results are shown in Table 4. The 6 “Or value for the framework oxygen was low, 610, which may indicate a high-temperature origin. Although 6 IsOr values for laumontite could be reproduced to only + 1‘%o,the value obtained here is considerably lower than that reported by STALLARDand BOLES ( 1989) on S-type laumontites from volcanogenic sediments of the Hokonui Hills, New Zealand. Their values ranged from 14.5 to 15%~.These two measurements on laumontite show that laumontite has at least a 9%0range in nature. At this point it is too early to speculate on the significance of this spread in terms of oxygen isotope geothermometry. The laumontite is associated with calcite whose 6’*0 is 3.57~ and b13C is -2.6%0. According to petrographic textures the two minerals either cocrystallized or the laumontite grew first, Assuming that the fluid composition was constant and the crystallization sequence resulted from cooling then these results indicate that laumontite has a greater tendency to enrich I80 than does calcite. This conclusion, however, is contrary to what is expected from mineral stoichiometry arguments, which suggest that laumontite should be similar to anorthite in terms of oxygen isotope fractionation and thus concentrate less I80 than calcite. One plausible explanation is that laumontite exchanges its oxygen isotopes more readily than calcite and hence has a lower closure temperature.

Both oxygen and hydrogen isotope ratios were obtained from the channel water of the Icelandic laumontite. Although the composition of this water (6 “0,: - 16.4%0;6D: -m-81.5% ) falls farther to the left of the meteoric water line than any channel water measured in other zeolites, it is still close enough to the MWL to indicate a related origin of the laumontite channel water (Table 4 and Fig. 5). Recent work by KARLSSON ( 1988) and KARLSSONet al. ( 1989) indicates that channel water can readily undergo oxygen isotope exchange with surrounding water vapor. Mordenite Oxygen isotope ratios were measured in one hydrothermal sample from Iceland (IMNH-7556). Table 4 displays the results. The sample yielded an intermediate d I80 f value of + 17.8% for framework oxygen. If this mordenite were formed at - 100°C then a fluid of 6 I80 = O%Owould be required, assuming equilibrium and a fractionation factor similar to calcite. A 6180, = -9.7k was obtained for the channel water. Exchange studies by KARLSSON( 1988) and KARLSSONet al. ( 1989) suggest that the channel water in the mordenite was derived from the Chicago water vapor, analogous to the channel waters in chabazite and clinoptilolite. Natrolite Oxygen and hydrogen isotopes were determined in a hydrothermal sample from Cape Split, Nova Scotia (FMNHM3340). Isotopic data are given in Table 4. Oxygen isotope analyses from two separate dehydrations gave an average 6 “0 f value of lS.OL, which may be indicative of low-temperature origin or a fluid rich in “0. LAMBERTet al. ( 1988) obtained a slightly lower value of 13.0% for a sample from south Texas, which they suggest formed at -67°C. Two channel waters were measured. For one sample both 6 180, and 6D values were obtained. Only 6 180, was analyzed in the other water sample (Table 4). Just as was observed for the other zeolites species, the isotopic composition of the natrolite channel water is close to the meteoric water line (Fig. 5). The excellent agreement between the d 180nv of the two waters run in November 1986, -7.7460, and May 1987, -7.9%+, signifies that channel water in natrohte does not reflect laboratory water vapor but rather that of the sample locality, like channel water in analcime. Framework Oxygen-General Ranges of 6 I80 r values in natural zeolites that were observed in this study, along with those of other workers, are summarized in Fig. 6. The lowest 6180r value is that of hydrothermal analcime, which is 4.3%~ The highest value obtamed for a zeolite to date is that of deep-sea phiIIipsite, which is 355I ( B~HLKE et al., 1984). This phillipsite value may even be an underestimate, since it is not clear whether the channel water had been eompletely removed.

Isotope geochemistry of zeolites

H

(2)

H

NJ

FIG. 6. Variations of 6 ‘*Orin natural zeolites. The figurecompiles the range of 6 ‘*Or values in natural zeohtes obtained in this study and by others. Zeohtes are ordered according to their Si/AI ratio from those with the lowest ratio at the bottom to those with highest ratio at the top. Individual analyses are shown as dots and ranges as horizontal bars. For each range the number of samples is indicated

by a number (N). Data obtained in the present work are unmarked; data from other sources are identified by the number in parentheses. Sources: ( 1) SAWN( I967 ): ( 2 ) STALLARD and BOLES ( 1986.1989 ): (3) SAW; (1973); (4) E%&L~E et al. (1984); (5) LEE(1987); (kj LAMBERTet al. ( 1988).

Three different factors could account for the variations in 6 “Or values shown in Fig. 6. These are 1) mineral chemistry, 2) isotopic composition of coexisting fluids, and 3) temperature. TAYLOR and EPSTEIN( 1962) were the first to suggest that there was a simple relation between mineral stoichiometry and the tendency to concentrate “0. They noted that silica-rich minerals were usually more enriched in “0 than silica-poor minerals and explained this orderly sequence of ‘*O enrichment in terms of the relative bond strength of oxygen atoms in the minerals (SiO > A10 > M”O > M’O). These observations have since then been borne out by laboratory experiments ( O’NEIL and CLAYTON, 1964; CLAYTON et al., 1972; BECKER and CLAYTON, 1976). O’NEIL and TAYLOR (1967) demonstrated that variations in the Si/Al ratio have a pronounced effect on oxygen isotope ratios in feldspars. Thus albite (Si/Al = 3) had a greater affinity for ‘*O than did anortbite (Si/Al = 1). In contrast to plagioclase, the alkali feldspar endmembers had similar oxygen isotope ratios, indicating that variations in the Na/K ratio had a negligible effect. However, the oxygen isotopic effect between albite and anorthite can be considerable, especially at low temperatures. CLAYTON ( I98 1) used the fractionation factors from MATSUHISAet al. ( 1979) to show that at 500°C albite is I .8%0more enriched in 6 “0 than anorthite. Since the zeolites are closely related to the feldspars both structurally and chemically (DEER et al., 1971), it is reasonable to expect similar

oxygen isotopic

effects in them.

Hence,

high-Si/Al

zeolites such as clinoptilolite (4.2 to 5.2) or mordenite (4.2 to 5.0) should concentrate “0 compared with low-Si/Al zeolites such as phillipsite ( 1.3 to 2.2) or analcime (2.0 to 2.7). Figure 6, however, reveals no such simple relationship

1379

between Si/Al ratios and 6180r in zeolites. In conclusion, although the Si/Al ratio varies significantly more in the zeolites ( 1.5 to 5.0) than in the feldspars ( 1 to 3), variations in 6 “Or among the zeolites are probably mostly controlled by fluid oxygen isotope composition and temperature. Fluid composition and temperature are the main variables that determine the oxygen isotope ratios in a given mineral (e.g., O’NEIL, 1977). The effects of these variables will be discussed below only with respect to analcime because it is the best constrained in terms of data. As is demonstrated on the basis ofthe experimental data discussed in our companion paper and the association between analcime and calcite, the fractionation of oxygen isotopes between analcime and water is similar to that between calcite and water. Hence, oxygen isotope variations in analcime can be explained by comparison with the calcite-water system. Below is a discussion of those samples that were best constrained by field data, namely the S-type analcime from the Aquarius Cliffs, Barstow, Big Sandy, and Green River formations, and the H-type analcime from Eyrarfjall and Oddskard, Iceland. S-type analcime Oxygen isotope results for the saline-alkaline lake deposits indicate that the formation conditions for the Green River analcime ( 16.6%0) differed from that of the Barstow and Big Sandy formations (22.5 to 24.5%0). These variations in 6’80r values of the authigenic analcime could arise from differences in the isotopic composition of the precursor material(s) or in the pore fluids and/or differences in temperature. Analcime in these deposits has never been found associated with relict volcanic glass but rather with other alkali-rich zeolites, such as clinoptilolite and phillipsite. Hence, it has been suggested that analcime did not form as a result of reactions between the pore fluids and glass-the main source material in these formations-but from preexisting alkali-rich zeolites ( IIJIMA and HAY, 1968; SHEPPARDand GUDE, 1969, 1973). If precursor material was the principal factor governing the 6 “0 f values, then the lower value for the Green River analcime would imply either a greater contribution from sources with low 6 “0 values, such as volcanic glass, detrital igneous quartz, or feldspar, or conversely a smaller contribution from sources with high 6 I80 values, such as preexisting authigenic minerals such as clays or zeolites, for example. The combined effects of changes in isotopic composition of the pore fluids and in the temperature can be estimated from Fig. 7. Formation temperatures in these deposits are estimated to have ranged from -40 to 70°C (IIJIMA and HAY, 1968; R. L. HAY, pers. comm., 1985). Given this temperature range and the 6 “Or values, it is possible to outline fields for the Green River (CR) and the Barstow-Big SandyAquarius Cliffs (BSA) analcime in Fig. 7. According to this figure the pore fluids in equilibrium with the analcime at Green River Formation (-8 to -3%) could have been as much as 13%~’lower in 6 I80 than those in either the Barstow, Big Sandy, or Aquarius Cliffs formations (-3 to 5%0). These differences could reflect variations in pore fluid salinities and/ or input meteoric waters. More data are needed to resolve this question.

H. R. Karlsson and R. N. Clayton

1380

-250°C. Oxygen and hydrogen isotope evidence from the Iceland Research Drilling Project (IRDP) suggests that the fluid at Oddskard was probably similar to that of presentday meteoric waters. HATTORI and MUEHLENRA~HS( 1982) estimate that the 6180 value ofthe original fluid in the IRDP core was - 1I X0. Origin of Channel Water

0

100

200

300

Temperature (“C) FIG. 7. Plausible formation conditions of sedimentary analcime in saline-alkaline lake deposits of the Aquarius Cliffs, Barstow, Big Sandy, and Green River formations. The data in the plot indicate that the fluid composition in the Green River (CR) formation was lower in 6”O than those of the Aquarius Cliffs, Barstow, and Big Sandy formations (BSA). In the latter three, however, the fluid 6 “0 values were similar.

ff-type analcime

It is possible to constrain temperature of formation for some of the hydrothermal analcimes listed in Table 1 by assuming that the fluid composition was that of present-day local meteoric water. These estimates will yield a minimum temperature, since oxygen isotope exchange between rock and fluid may increase the 6 I80 of the fluid (CRAIG, 1963; CLAYTONet al., 1968). On this basis temperatures were estimated for two Icelandic samples: Eyrarfjall (HRK-82-57 ) and Oddskard (IMNH-1556). The Eyrarfjall analcime occurred in the Quaternary basalts of western Iceland. It was taken from a sea-cliff outcrop and was located approximately 5 m from a basalt dike. The Oddskard analcime occurred -620 m above sea level in one of the classic zeolite zones in the Tertiary basalts of eastern Iceland (analcime zone; WALKER, 1960). The 6D of the waters at the Eyrarljall and Oddskard localities are approximately -54 and -686, respectively ( ‘ARNASON,1976). Given these 6D values, we can calculate the corresponding 6 I80 values from MWL relation (6D = 86 I80 + 10). The calculation gives -8% for Eyrarljall and -10% for Oddskard. Assuming equilibrium and the same fractionation as for the calcite-water system, the estimated temperature of formation for these samples was -80°C. This temperature seems reasonable in the case of Oddskard but seems low for the Eyrarljall, considering its close proximity to the dike mentioned above. The analcime may be younger than the dike and thus have formed at a lower temperature. Alternatively, the 6 “0 value of the fluid at Eyratijall may have been higher than that of present-day precipitation. The location next to the sea indicates that the fluid may have been seawater, similar to fluids in the active Reykjanes geothermal system, Iceland ( SVEINBJ~RNSD~TTIR et al., 1986). In that case the temperature would have been

The oxygen and hydrogen isotopic compositions of the channel waters indicate that the channel water in analcime and natrolite may be related to meteoric waters at the sample site. However, isotope compositions of channel waters in chabazite, clinoptilolite, laumontite, and mordenite most likely reflect those of ambient water vapor in the Chicago laboratory. The isotopic composition of channel water in latter zeolites is therefore of no geologic value. This subject is discussed further by KARLSSON ( 1988) and KARLSSON et al. (1989). Framework-Channel Water Thermometer HAMZA and EPSTEIN ( 1980) originally suggested that it might be possible to calibrate a single mineral geothermometer based on the fractionation of oxygen isotopes between different sites in a mineral. Since then there have been numerous attempts to calibrate such internal thermometers ( PICKTHORNand O’NEIL, 1985; ALPERSet al., 1988 ). A natural outgrowth of our study is an investigation of the oxygen isotope fractionation between the analcime framework and the channel water. Tables 1 and 2 show that framework and channel water may differ by as much as 40% in 6’“0, implying that an internal geothermometer is conceivable. This difference is, however, misleading because the channel water is much more readily exchanged than the framework ( KARLSSONand CLAYTON, 1990b) such that high-temperature equilibrium fractionations (> 100°C) are not preserved except for active systems such as Surtsey. Role of Zeolites in Water-Rock Interaction Zeolites are common alteration minerals in diagenetic and hydrothermal environments and must therefore be important in rock-water interaction. Due to lack of oxygen isotope data, however, little is known about their role in this process. With the new information obtained in the present study it is possible to learn more about their function. Pore,fluids in deep-sea sediments

Systematic chemical and isotopic changes have been documented in pore fluids squeezed from sediments in DSDP (Deep Sea Drilling Project) cores (LAWRENCEet al., 1975, 1976, 1979; FRIEDMANand HARDCASTLE,1973, 1974; ELDERL~ELDet al., 1982; MATSUHISA, 1985). The basic trend is for Mg 2+, K’, and 6t80 to decrease and for CaZt to increase with depth. These chemical and isotopic variations have generally been attributed either to alteration of the underlying basalts in Layer II of the oceanic floor or to alteration of volcanic ash in the sediments (LAWRENCE et al., 1975; GIESKES and LAWRENCE,198 1).

Isotope geochemistry of zeolites

Phillipsite and clinoptilolite are the most abundant zeolites in deepsea sediments (BOLES, 1977; KASTNER, 1979 ) . They can constitute up to 80% by weight of the sediment ( CZYNCINSIU, 1973). As these zeohtes are potassium-rich and calcium-poor species, their formation in deep-sea sediments could account at least partially for some of the chemical changes discussed above. It has been shown in this work and by SAWN ( 1967, 1973) that clinoptilolite and phillipsite concentrate “0 relative to fluids at low temperatures. Formation of these minerals in deep-sea sediments could therefore contribute to the lowering of 6 I80 values observed in the pore fluids. Basalt-seawater interaction

MUEHLENBACHSand CLAYTON ( 1972) found a strong positive correlation between the 6 “0 and the water content of some altered submarine basalts. They attributed this observation to the formation of water-rich alteration minerals with high 6 “0 values from primary igneous minerals with low 6 “0 values. Hence, the positive correlation between the 6 ‘*O and the water content could be explained as a mixing curve having fresh basalt as one endmember and hydrous alteration mineral(s), formed at low temperatures, as the other endmember. MUEHLENBACHS( 1980) measured oxygen isotope ratios in a large number of whole rocks from DSDP Site 417 and found a positive correlation between 6 “0 values and water content (H*O+). The 6 “0 enrichment was particularly striking in the basalts from Hole 4 17A. Basalts in this hole ranged in 6’*0 values from 6.7 to 26.4%-the latter being the highest 6 “0 value recorded for a deepsea basalt. MUEHLENBACHS ( 1980) explained the 6 “0 enrichment seen in Hole 417A in terms of a fresh basalt-hydrous mineral mixing model. Two analcime samples analyzed in the present study yield 6’*Or of 24.1 and 24.2% (Table 1), and plot close to a basalt-smectite mixing line for the basalts. Although these analcime 6 “0 r values are too low to account for the highest d’*O basalts at -26%0, they can explain the lower 6 “0 basalts. Acknowledgmenfs-This work was a part of the senior author’s Ph.D. dissertation and was funded by NSF grant EAR-8616255 to R. N. Clayton. Field work in Iceland was supported by a grant from the Oherine. Fund of the Department of the Geouhvsical Sciences at the Univemity ofchicago. This work would not have been possible without the expertise, technical advice, and assistance of Toshiko K. Mayeda. We are deeply grateful to James R. O’Neil for generously allowing H.R.K. to work in his laboratory at the USGS in Menlo Park. We are indebted to Pete Dunn, Sveinn P. Jakobsson, Edward Olsen, Alan Edgar, Thomas Pearce, Jeffery Alt, Richard Hay, Richard Sheppard, Paul Moore, Michael Berry, Douglas Ming and James Luhr for donating excellent samples. We thank James Eason for doing a superb job in proofreading the manuscript. Helpful reviews were provided by Drs. James Boles, Michael Bird, and Samuel Savin. Editorial handling: J. R. O’Neil REFERENCES ALPERS C. N., NORDSTROMD. K., and WHITE D. (1988) Solid

solution properties and deuterium fractionation factors for hydronium-bearing jarosites from acid mine waters (abstr.). Eos 69, 1480.

1381

‘ARNASON B. ( 1976) Ground Water Systems in Iceland Traced by Deuterium. Societas Scientiarum Islandica. BECKER R. H. and CLAYTON R. N. ( 1976) Oxygen isotope study of a Precambrian banded iron formation, Hamersley Range, Western Australia. Geochim. Cosrnochim.Acta 40, 1153-l 165. ( 1952) Conversion of hydrogenic materials to hydrogen for isotopic analysis. Phys. Chem. Minerals 24, 1356-I 357. BOHLKEJ. K., ALTJ. C., and MUEHLENBACHS K. ( 1984) Oxygen isotope-waterrelations in altered deep-sea basal& Low-temperature mineralogical controls. Phys. Chem. Minerals 21,67-77. BOLESJ. R. (1977) Zeolites in deep-sea sediments. In Mineralogy and Geologyof Natural Zeolites(ed. F. A. SAND), Vol. 4, Chap. 7, pp. 137-163. Mineralogical Society of America. BRECKD. W. ( 1974) ZeoliteMolecular Sieves. Wiley. CLAYTONR. N. ( 1981) Isotopic thermometry. In Thermodynamics of Minerals and Melts (eds. R. C. NEWTON,A. NAVROTSKY,and B. J. WOOD), Vol. 6, Chap. 5, pp. 85-109. Springer-Verlag. CLAYTONR. N. and MAYEDAT. K. ( 1963) The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta 21,43-52. CLAYTONR. N., MUFFLERL. J. P., and WHITED. E. ( 1968) Oxygen isotope study of calcite and silicates of the River Ranch No. 1 well, Salton Sea geothermal field, California. Amer. .I. Sci. 266, 968979. CLAYTON R. N., O’NEIL J. R., and MAYEDAT. K. ( 1972) Oxygen isotope exchange between quartz and water. J. Geophys.Res. 77, 3057-3067. CRAIGH. ( 1957) Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta 12, 133-149. CRAIGH. ( 1961) Isotopic variations in meteoric waters. Phys. Chem. Minerals 133, 1702- 1703. CRAIGH. ( 1963) The isotopic geochemistry of water and carbon in geothermal areas. In Nuclear Geology on GeothermalAreas (ed. E. TONGIORGI),pp. 17-53. Cons&ho Nazionale della Richerche. CZYNCINSKIK. ( 1973) Authigenic phillipsite formation rates in the central Indian Ocean and equatorial Pacific Ocean. Deep-Sea Res. 20, 555-559. DEERW. A., HOWIER. A., and ZLJ~SMAN J. ( 1971) An Introduction to the Rock-Forming Minerals. Longman Group Ltd. ELDERFIELDH., GIE&S J. M., BA&R P. A., GLDF~ELDR. K., HAWKESWORTH C. J.. and MILLERR. C1982) *‘Sr/%r and ‘*O/ I60 ratios, interstitial ‘water chemistry and diagenesis in deepsea carbonate sediments of the Ontong Java Plateau. Geochim. Cosmochim. Acta 46,2259-2268. FAUREG. ( 1986) Principlesof Isotope Geology.Wiley. FERGLJSONL. J. ( 1977) Petrogenesis of the analcime-bearing and associated volcanic rocks of the Crowsnest Pass area, Alberta. MSc. dissertation, Univ. Western Ontario. FRIEDMANI. and HARDCASTLE K. ( 1973) Interstitial water studies, Leg 15-isotopic composition of water. Init. Rept. Deep Sea Drill. Proj. 20,901-903. FRIEDMANI. and HARDCASTLE K. ( 1974) Deuterium in interstitial waters from Red Sea Cores. Init. Rept. Deep Sea Drill. Proj. 23, 969-970. GARLICKG. D. ( 1965) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rocks. Ph.D. dissertation, California Institute Technology. GIESKESJ. M. and LAWRENCEJ. R. ( 198 1) Alteration of volcanic matter in deep-sea sediments: Evidence from the chemical composition of interstitial waters from deepsea drilling cores. Geochim. Cosmochim. Acta 45, 1687-1703. GOI-~ARDIG. and GALLIE. ( 1985) NaturalZeolites.Springer-Verlag. GREGORYR. T. and CRISSR. E. ( 1986) Isotopic exchange in open and closed systems. In Stable Isotopesin High Temperature Geological Processes (eds. J. W. VALLEY,H. P. TAYLOR,JR., and J. R. O’NEIL), Vol. 16, Chap. 1, pp. 91-127. Mineralogical Society of America. HAIMSONM. and KNAUTH L. P. ( 1983) Stepwise fluorination-a useful approach for the isotopic analysis of hydrous minerals. Getchim. Cosmochim. Acta 47, 1589-1595. BIGELEISEN J., PERLMANM. L., and FRASER H. C.

H. R. Karlsson and R. N. Clayton

I382

M. S. and EPSTEINS. ( 1980) Oxygen isotopic fractionation between oxygen ofdifferentsites in hydroxyl-bearing silicate minerals. Geochim. Cosmochim. Acra 44, 113- 182. HATTORIK. and MUEHLENBACHS K. ( 1982) Oxygen isotope ratios of the Icelandic crust. J. Geophys. Rex 87, 6559-6565. HOEF~J. ( 1987) Stable Isotope Geochemistry. Springer-Vedag. HONNOREZJ., EMMERMANNR.. HUBBERTENH. W.. LAWRENCE C.. and MUEHLENBACHS K. ( 1983) Alteration processes in layer 2 basalts, DSDP Hole 504B, Costa Rica Rift. Inil. Repr. Deep Sea HAMZA

Drill. Proj. 69, 509-546.

IIJIMAA. and HAY R. L. ( 1968) Analcime composition in the Green River Formation of Wyoming. Amer. Mineral. 53, 184-200. JAKOBSON S. P. ( 1978) Environmental factors controlling the palagonitization of the Surtsey tephra, Iceland. Bull. Geol. Sot. Denmark 27, 91-105.

.4mer. 97, 648-659.

H. R. ( 1988) Oxygen and hydrogen isotope geochemistry of zeolites. Ph.D. dissertation, Univ. Chicago. KARL~SONH. R. and CLAYTONR. N. ( 1990a) Analcime phenocrysts in igneous rocks: Primary or secondary? Amer. Mineral. (submitted). KARLSSON H. R. and CLAYTONR. N. ( 1990b) Oxygen isotope fractionation between analcime and water: An experimental study. Geochim. Cosmochim. Acta 54, 1359- 1368 (this issue). KARLSSONH. R., CLAYTONR. N., and MAYEDAT. K. ( 1985) Analtime: A potential geothermometer (abstr.). Geol. Sot. Amer. 17, 622-633. KARLSSON H. R., CLAYTONR. N.. and MAYEDAT. K. ( 1988) Fractionation and kinetics of oxygen isotope exchange between analtime and water (abstr.). Eos 69, 527. KARLSSON H. R., CLAYTONR. N.. and MAYEDAT. K. ( 1989) Water in zeolites: Fractionation of oxygen and hydrogen isotopes during exchange and dehydration (abstr.). Geol. Sot. Amer. 21, A 12. KASTNERM. ( 1979) Zeolites. In Marine Minerals (ed. R. G. BURNS), Vol. 6. Chap. 4, pp. I 1I - 122. Mineralogical Society of America. KISHIMAN. and SAKAIH. (1980) Oxygen-18 and deuterium determination on a single water sample of a few milligrams. .4nal. C’hcmistry 52, 356-358. LAMBERI’D. D., MALEKD. J., and DAHL D. A. ( 1988) Rb-Sr and oxygen isotopic study of alkalic rocks from the Trans-Pecos magmatic province, Texas: Implications for the petrogenesis and hydrothermal alteration of continental alkalic rocks. Geochim CosKARLSSON

.4na 52, 2357-2367.

LAWKENC‘E J. R., GIESKESJ. M., and BROECKER W. S. ( 1975) Oxygen isotope and cation composition of DSDP pore waters and the alteration of Layer II basalts. Earfh Planet. Sci. Letr. 27, l-10. LAWRENCE J. R., GIFXES J. M., and ANDERSON T. F. ( 1976) Oxygen isotope material balance calculations. Leg 35. Init. Rep/. Deep Sea Drill. Proj. 35, 507-5

mochim.

Acta 44, 381-402.

MCCREAJ. M. ( 1950) On the isotopic chemistry of carbonates and the paleotemperature scale. J. Chem. Physics 18, 849-857. MING D. W. and DIXONJ. B. ( 1987) Technique for the separation of clinoptilolite from soils, Clays Clay Minerals 35, 469-472. MUEHLENBACHS K. ( 1980) The alteration and aging of the basaltic layer of the sea floor: oxygen isotope evidence from DSDP/IPOD Legs 5 I, 52, and 53. Init. Rep1 Deep Sea Drill. Proj 51, 52, 53, 1159-l 167.

MUEHLENBACHS K. and CLAYTUNR. N. ( 1972) Oxygen isotope studies of fresh and weathered submarine basalts. Canadian .I Earth Sci. 9, I72- 184.

O’NEIL J. R. ( 1977) Stable isotopes in mineralogy. P~KY. Chem. Mineral. 2, 105-l 23.

JAKOBSSON S. P. and MOOREJ. G. ( 1986) Hydrothermal minerals and alteration rates at Surtsey volcano, Iceland. Bull. Geol. Sot.

mochinl.

metamorphism: A study of albite crystallization. Geochinz. (‘OS-

12.

LAWRENCEJ. R.. DREVERJ. I., ANDERSONT. F., and BRUECKNER H. K. ( 1979) Importance of alteration of volcanic material in the sediments of Deep Sea Drilling Site 323: Chemistry, ‘*O/ I60 and “7Sr/XhSr.Geochim. Cosmochim. Acta 43, 573-588. LEE Y. I. ( 1987) Isotopic aspects of thermal and burial diagenesis of sandstones at DSDP site 445, Daito Ridge, northwest Pacific Ocean. Chcm Geology (Isotope Geosci. Sect.) 65, 95- 102. Luttn J. F. ( 1980) The Colima volcanic complex, Mexico. I. Postcaldera andesites from Volcan Colima. Ph.D. dissertation. Univ. California. Berkeley. LUHR J. F. and CARMICHAELI. S. E. ( 1981) The Colima volcanic complex, Mexico: Part II. Late-Quaternary cinder cones. Contrih. .Vinerul. Petrol. 76, I II- 147. LUHR J. F. and KYSERT. K. ( 1989) Primary igneous analcime: The C‘olima minettes. Amer. Mineral. 74, 216-223. MA’I‘SIII-~ISA Y. ( 1985) Oxygen isotope ratios of interstitial waters from the Nankai Trough and the Japan Trench, Leg 87. Inil. Rcpt. &~c,p.Seu Drill. Proj 87, 853-856. MATSUHISAY.. GOLDSMITHJ. R.. and CLAYTONR. N. ( 1979) Oxygen isotopic fractionation in the system quartz-albite-anorthitewater. Geochim. Cosmochim. Acta 43, 1131-l 140. MATTHEWSA. ( 1980) Influences of kinetics and mechanism in

O’NEIL J. R. and CLAYTONR. N. ( 1964) Oxygen isotope thermometry. In Isotopic and Cosmic Chemistry (eds. H. CRAIG,S. MILLER, and G. WASSERBURG),pp. 157-168. North-Holland. O’NEIL J. R. and EPSTEINS. ( 1966) A method for oxygen isotope analysis of milligram quantities of water and some of its applications. J. Geophys. Res. 71,4956-4961. O’NEIL J. R. and TAYLORH. P., JR. ( 1967) The oxygen isotope and cation exchange chemistry of feldspar. Amer. Mineral. 52, 14141437. O’NEIL J. R.. CLAYTONR. N., and MAYEDAT. K. ( 1969) Oxygen isotope fractionation in divalent metal carbonates. J Chum. Ph.ysics 51, 5547-5558.

PICKTHORNW. J. and O’NEIL J. R. ( 1985) “0 relations in alunite minerals: potential single-mineral thermometer (abstr.). Gee/. Sot. .Imer. 17, 689. VAN REEUWIJKL. P. ( 1974) The ‘Thermal Dehydratmn of Natural Zeo1ite.s. Medelelingen Landbouwhogeschool Wageningen, 74-9. Ross C. S. and SHANNONE. V. (1925) Mordenite and associated minerals from near Challis, Custer County, Idaho. Proc US Natl. Museum 64, 1-2 1. SAVIN S. M. (1967) Oxygen and hydrogen isotope ratios in sedimentary rocks and minerals. Ph.D. dissertation, California Institute Technology. SAVIN S. M. ( 1973) Oxygen isotope studies of minerals in ocean sediments. Proc. Intl. Symp. Hydrogeochem. Biogeochem., 372391.

SAVINS. M. and EPSTEINS. ( 1970a) The oxygen and hydrogen isotope geochemistry of clay minerals. Geochim. Cosmochim. Actu 34,25-42.

SAVIN S. M. and EPSTEINS. ( 1970b) The oxygen and hydrogen geochemistry of ocean sediments and shales. Geochim. Cormochim. 4cta 34,43-63.

SHEPPARDS. M. F. ( 1986) Characterization and isotopic variations in natural waters. In Stable Isotopes in High Temperature Geological Processes (eds. J. W. VALLEY,H. P. TAYLOR,JR., and J. R. O’NEIL), Vol. 16. Chap. 6. pp. 165-185. Mineralogical Society of America. SHEPPARDR. A. and GUDE A. J., III (1969) Diagenesis of tuffs in the Barstow Formation, Mud Hills, San Bernardino County. California. CrS Geol. Surv. ProJ: Paper 634, l-35. SHEPPARDR. A. and GUDE A. J., III ( 1973) Zeolites and associated authigenic silicate minerals in tuffaceous rocks of the Big Sandy Formation, Mohave County. Arizona. LIS Geol. Surv. Pro/: Paper 830, l-36.

STALLARDM. and BOLESJ. R. ( 1986) Stable oxygen isotopes in laumontite, stilbite, and heulandite (abstr.). Ann. Meet. Clays Clay, Mineral Sot.

STALLARDM. and BOLESJ. R. ( 1989) Oxygen isotope measurements of albite-quartz-zeolite mineral assemblages, Hokonui Hills, Southland. New Zealand. C/ays Clay Minerals 37, 409-4 18. STAUDIGELH., MUEHLENBACHS K.. RICHARDSON S. H.. and HART S. R. ( 1981) Agents of low temperature ocean crust alteration. Cbntrih. ,Mineral. Petrol. 71, 150- 157. SVEINBJ~RNSD~TTIRA. E., COLEMAN M. L., and YARDLEY B. W. D. ( 1986) Origin and history of hydrothermal fluids of the Reykjanes and Krafla geothermal fields, Iceland. Contrih. Mineral. Petrol. 94, 99-109.

Isotope geochemistry of zeolites TAYLOR H. P., JR., and EPSTEINS. ( 1962) Relationship between

01*/Oi6 ratios in coexisting minerals of igneous and metamorphic rocks. Part 2. Application to petrologic problems. Bull. Geol. Sot. Amer. 13,6X-694. WALKER G. P. L. ( 1960) Zeolite zones and dike distribution in relation to the structure of the basalts in eastern Iceland. J. Geology 68,5 15-528. YAPP C. J. ( 1987) Oxygen and hydrogen isotope variations among goethites (c~-FeOOH) and the determination of paleotemperatures. Geochim. Cosmochim. Acta 51,355-364. YURTSEVERY. and GAT J. R. ( 198 1) Atmospheric waters. In Stable IsotopeHydrology:Deuterium and Oxygen-18 in the WaterCycle (eds. J. R. GAT and R. GONFL~NTINI);Tech. Rept. Ser. 210, Chap. 6, pp. 103-142. Inn. Atomic Energy Agency.

a core which is mantled by chabazite. The hugest analcime crystal is 1 cm in diameter. The chabazite forms cubes that reach 1 cm in width and appears to have crystallized before analcime (H). FMNH-9255

Locality: Brevik, Norway. Description:Vug or vein in a scoriacious basalt occupied by analcime and natrolite. Analcime forms large well-shaped crystals that reach 1.5cm in diameter (H).

FMNH-M9257

Locality: Edgewater, New Jersey, USA. Description: The sample (5 X 4 X 3.5 cm) contains at least three different mineral species. Most abundant is a white fibrous natrolite. The natrolite needles reach a length of 3.5 cm. In one place the natrolite needles are intergrown with calcite. The third most common mineral forms stout gray crystals with a tabular habit that are up to 1 mm long (natrolite?). Intergrown with this tabular mineral are analcime icosahedra that reach 3 mm in diameter (H).

Green River

Locality: Sand Butte, Laney Shale Member, Green River, Wyoming, USA. Description:Light yellow to brown sandstone made almost entirely of euhedral to subhedral analcime icosahedra (S). Source: Richard Hay, University of Illinois.

IMNH-1556

Locality: Oddskard (mt. pass), 620 m.a.s.l., on the Nordtjardar side, S. MGlasysla, E. Iceland. Description:Vug or vein. Four large euhedral crystals (H).

IMNH-7522

Locality:Melshom (mountain), Berutjordur, S. Mulasysla, S.A. Iceland. Description: Vug or vein. A large euhedral crystal -2 cm in diameter (H).

HRK-82-57

Locality: Eyrartjall Mt., West Iceland. Description:Amygdules in a Quatemary basalt. Sampled from a seacliff approximately 5 m from a dike. The analcime occurs as clear icosahedra whose diameter reaches 2 mm (H). Source: Haraldur R. Karlsson.

Keweenaw

Locality:Copper Falls, Keweenaw Peninsula, Northern Michigan, USA. Description:A fist-size nugget collected at a mine dump site. Three minerals occur in the sample: massive, light-pink datolite (most abundant), euhedral, reddish analtime, and subhedral, gray calcite rhombs. The analcime and calcite are embedded in the datolite. Individual analcime crystals reach 1.5cm in diameter. The crystallization sequence is analcime r calcite > datolite (H). Source: Michael Berry, Copper Harbor, northern Michigan.

M3-105

Locality: Big Sandy, California, USA. Description:Analcimic sandstone (tutl), consisting almost entirely of analcime. Some quartz and feldspar present. (Hay field No. 65-4-5C) (S). Source: Richard Hay, University of Illinois.

PN-39E

Locality: Crowsnest, Alberta, Canada. Description:Large brown euhedral analcime

APPENDIX DESCRIPTIONS

OF SAMPLES

Below is a brief description of samples studied in this work. A more detailed description can be obtained from the senior author. Most of the samples were obtained from museums or from individuals. All other samples are referred to by the original museum or collector’s number. Abbreviations in front of the sample numbers that refer to particular museums or collections are as follows: FMNH

Field Museum of Natural History, Chicago. Source: Edward Olsen.

IMNH

Icelandic Museum of Natural History, Reykjavik, Iceland. Source: Sveinn P. Jakobsson.

UC

University of Chicago rock and mineral collection, Chicago. Source: Paul B. Moore.

USNM

US National Museum (i.e., The Smithsonian Institution), Washington. Source: Peter Dunn.

The letter in parentheses following each sample description refers to occurrence type: S, sedimentary; H, hydrothermal; I, igneous. Analcime Big Sandy

Locality: Big Sandy, California, USA. Description: Fine-grained sandstone (tutT) similar to M3-105 (S). Source: Richard Hay, University of Illinois.

DSDP-4 I7A 38-2

Locality: Deep sea drill core DSDP site 4 17, Bermuda Rise, 25”6’N, 68”2W, North Atlantic Ocean. Description:A slab ofbasalt, 5.2 cm wide, 7.5 to 9.5 cm long, and 1cm thick, cut lengthwise through the core. Two hydrothermal alteration veins are exposed iu the basalt. The first of these is in the lower half of the slab and is -0.5 to 1.5 cm wide. This vein consists of euhedral analcime icosahedra which are up to 1 mm in diameter. When a piece of this analcime was broken off in one place, a core of caicite was revealed. The second vein is fully exposed at the top edge of the slab. It consists of a small analcime iscosahedron (5 1 mm in diameter) imbedded in gray calcite (H). Source; Jeffrey Ah, Washington University.

FMNH-M6743

Locality:Golden, Colorado, USA. Description: An amygdule or vein in pyroelastic rock (basalt?) consisting of analcime and chabazite. Analcime icosahedra form

1383

1384

H. R. Karlsson and R. N. Clayton phenocrysts in an analcime phonolite. This sample is described in detail by Pearce (1970) (I). Source; T.H. Pearce, University of Western Ontario.

SAY-104

calcite crystals, 3 X 3 mm, occur on both analcime and green base (H). USNM-8363

tals on a yellow calcite rhomb. The calcite is - 1.5 cm in diameter and the coating. which is -2 mm thick, consists of a monolayer of analcime trapezohedra (H).

Locality: Colima Volcanic Complex, Jalisco,

Mexico. Description: Volcanic rock (minette)

containing olivine and phlogopite phenocrysts. Analcime occurs as euhedral microphenocrysts (520 pm in diameter) in the glassy matrix of this rock. A further description ofthis sample is given by LUHR ( 1980) and LUHR and CARMICHAEL(1981) (I). Source: James F. Luhr, Washington University. surtsey

Locality: Surbey Island, south of Iceland. Description: Analcime from a hydrothermal vein at 90.05 m depth in a borehole (JA-

SW-3-2

Locality: Big Sandy, California, USA. Descripfion: Sandstone (tuff) consisting en-

Mill Dump, Baltimore Tunnel. Maryland USA. Description: Fragments 1.5 cm across made of analcime and calcite. Analcime is most abundant as white massive crystals except for one fragment that contains small colorless trapezohedra. The calcite is mostly white to colorless Iceland Spar (H).

USNM-I 2792

Locality: Iganak, Greenland. Description: Fragmented clusters of white

analcime crystals. Chabazite and an unidentified fibrous zeolite are at the triple junctions between the analcime crystals (H). USNM-83258

Locdify

USNM-95982

Loditp

Aquarius Cliffs, Yavapai County, Arizona, USA. Description: Green sandstone of small round analcime grains. Individual analcime grains are less than 0.1 mm across (Sf.

USNM- I 15625

Locuhfy: Near Umalac, Guam, Pacific Ocean,

tirely of fine-wined analcime. For further description of this sample see SHEPPARD and GUDE (1973) (S). Source: Richard Sheppard, USGS. UC- I547

Localify: Nova Scotia, Canada. Description: Cavity filling. Analcime occurs

Locality: Chechy, Czechoslovakia. Descripfion: White massive analcime in places

USA. filling. One layer is of a brownish material of thickness of -5 mm. On top of the brown layer is a second layer of mostly grayish well-shaped analcime trapezohedra. The analcime crystals range from -2 mm to 1 cm in diameter(H). USNM-125977

Loculify: Giant’s Causeway, Antrim, Ireland. Description: Analcime on a basalt. Clear tra-

USNM-132507

USNM-7099

Locdify:

Quebec, Canada. crystal about 2.5 X 2.5 X 3.0 cm (H).

Two Islands, Minas Basin, Nova Scotia, Canada. Descripfion: An - I cm thick crust of milkwhite stubby analcime crystals blanketing a vug or a vein in a basalt. A large, 1.7 X I .7 cm, yellow transparent calcite (Iceland Spar) protrudes out of the analcime. The analcime and calcite appear to have grown contemporaneously (H).

USNM-137316

Locality:

USNM-145340

Faeroe Islands, North Atlantic

Ocean.

Locality: De Mix Quarry, Mt. St. Hi&e, Description: A large single euhedral analcime

pezohedral crystals up to 5 mm in diameter in a vein or vug (H). USNM-42 10

Loculify: Sicily. Descripfion: A basalt crusted with analcime

and calcite. Clear well-shaped analcime crystals are up to 4 mm in diameter. Yellow to white massive calcite fills in the interstices between the analcime crystals. Analcime appears to have crystallized before the calcite (H).

overgrown with an unidentified pink platy mineral. A vein of clear calcite cross-cuts the pink mineral but not the analcime. The crystallization order is analcime > pink mineral - calcite (H). USNM-354 I

South Table Mountain, Golden. Colorado, USA. Description: Fragments of white granular anal&me and yellow calcite (H).

Description: A piece of layered vein or vug

as large white euhedral crystals up to 2.5 cm in diameter. One such large analcime crystal was cut into 4 slabs of similar thickness. Samples taken for oxygen isotope analysis were taken from slabs A 1 and A2. A traverse of three samples was taken from core (Al-l) to edge (Al-5) (H). USNM-3536

Locality.- Powder

USNM- I 1884

KOBSSON

and MOORE, 1986). The analtime occurs as aggregates of euhedral clear crystals that together form an empty shelllike structure around what used to be calcite (pseudomorph) (H). Source: S. P. Jakobsson, Icelandic Museum of Natural History.

Locdify. Bergen Hill, New Jersey, USA. Description: Coating of white analcime crys-

Locality: Challis, Custer County, Idaho, USA. Description: Well-shaped clear analcime

crystals on top of a white to gray substratum of granular and layered calcite. The analtime falls into two size groups: two hemispherical crystals, 3 mm in diameter, and small globules,
sylvania, USA.

Description: Massive, white analcime resting

Description: A vug or vein filling consisting

on green base (pyroxene?). A couple of clear

of at least three different minerals. Anal-

Isotope geochemistry of zeolites time is the second most abun~t. It forms stout rn~~-whi~ well-shaped trapezohe&a. They range in diameter from f mm to 1 cm (H). USNM-151697

USNM-17271

USNM-155369

USNM-B I 7295

USNM-C3553

Locality:Prospect Park, New Jersey, USA. Descripfion:Vug or vein filling in a dark igneous rock (basalt) containing analcime, calcite, and quartz. Analcime is most abundant and consists of two white subhedral crystals. In contact with the analcime is smaller colorless calcite. Both analcime and calcite rest on drusy quartz and are intergrown with it at their bases (H). Lo&&y: Norheim near Kreuznach, Rheinland-P&z, Germany. Descri~~~n: Large inte~~king subh~ml anakime crystal% up to 1.5cm across, on a base of smaller euhedmi chabazite (H). Locality:Flinders, Victoria, Australia. Descripfion: A mineral lining in a cavity, 6 X 4.5 X 3 cm, of a dense greenish rock (basalt?) made up of analcime, calcite, a Iibrous mineral and a gray c~pt~~s~lline material. Clear analcime trapezohedra reach a diameter of 2 mm. In one pIace a large subhedrai “dog-tooth’ calcite crystal overlaps some of the analcime crystah (H). Locality: Antrim, Ireland. Descr~~rion:A vug or vein tilling consisting of analcime and calcite. The analcime forms subhedral to euhedral interi~~~g trapezohedra ranging in diameter from 1 mm to 1cm. Interstices between analcime crystals are partially filled by calcite (H). Loculify: Great Notch, New Jersey, USA. Description:A cavity or vein filling in altered basalt or diabase. The sampIe contains at least five different mineral species. Most abundant are white coiumnar crystals that tend to form clusters in the shape of cornbundles. Upon the columnar mineral (natrohte?) he analcime, calcite, a platy green mineral, and chabazite. The analcime forms white euhedral trapezohedra up to 3 mm in diameter. The calcite occurs as spherules -3 mm across (H).

USNM-R7 100

Locality:Adrendal, Aust Agder, Norway. Description:Vein(?) composed mainly of calcite and analcime. Calcite is more abundant than analcime and has two different habits: massive and scaienohedml. The analcime is clear trapezohedra up to 2 cm wide. Inferred ~s~ii~tion order is: massive calcite 2 analcime 2 scaienohedral calcite (H}.

USNM-R9674

Locality: Old Kilpatrick, Dumbarton, SGotland. Descr~p~~on: Vein or vug? Fmgmen~ ofwhite massive analcime and rhom~hedml calcite (Iceland Spar) (H).

USNM-RI 1059

Locality: Springfield, Lane County, Oregon, USA. Description:Vein or vug filling in an altered basalt composed of analcime and a fibrous zeoiite. Analcime forms intergrown euhe&al trapezohedra ranging in diameter Born 2 to 8 mm (H).

1385

USNM-RI IO62

Locality: Horni, Zaiezky, Chechy, Czechoslovakia. ~s~ipti~: Vug or vein in a basalt consisting of anaicime, calcite, and a Fibrous zeohte. Tbe base is made ofdeareuhedral analcime trapezohedra up to 2 mm across. Upon the analcime rest a few cm-sized yellow subhedral calcite rhombohedm (H).

USNM-134134

Locality: Spodumene mine (Foote Mineral Comp.) near Rings Mt., North Carolina, USA. Description: Vein or vug filling in dark igneous rock. The rock is coated on one side with a -3 mm thick layer ofwhite euhedml anaIcime trapezohedra. The largest analtime is -2 mm in diameter (H).

USNM- 120762

t;oc&y: Unnamed lacustrine deposit near Wikiuo. Mohave Countv. Arizona, USA. Descripti’dn:A green sandstone consisting nearly entirely of analcime crystals smaller than 0.1 mm in diameter (H).

2a

Locality:Crowsnest, Alberta, Canada, Description: Small round orange analtime phenocrysts from a blairmorite. A detailed description is given by BRGUSON (1977) (I). Source: Alan D. Edgar, Univenity of Western Ontario.

Chabazite FMNH-M6743

Locality:Golden, Colorado, USA. Description:See analcime FMNH-M6743.

FMNH-M9250

Locality:Bayof Fundy, Nova Scotia, Canada. Description:A cavity lined with pink chabazite and heulandite. The chabazite forms pink cubes up to 8 mm in width, which o&en penetrate one another. The host rock is a tuff or a basalt (H).

IMNH-2456

Locality: ‘OshSd, near K&It%River, W. fsafjardamj&a, N.W. Iceland. ~e~cr~~f~o~;Vug or vein. White crystalline aggtegate consisting of intergrown, euhedral cbabazite cubes. The largest ctystah are -4 mm in diameter (H).

INMH-7924

Locality: Mjbidalur (valley), 200 m,a.s.l., Isafrordur. N. isaIiardars~sla. N.W. Iceland. Deshiptiok Amygdule id b&t. The specimen, 2 X I.8 X 1 cm, is a portion of large euhedral single cry@. The hugest chabazite crystals found to date in Iceland, up to 2.7 cm across, are from this iocahty (S.P. JAKOBSON, pers.comm., 1988) (H).

SD+2

Locality: Durkee, Oregon, USA. Descripfkm: Light-yellow sandstone (t&f). Fine-Gina with small dark impurities. Sample was sieved ~thout crushing and the I50 to 200 mesh portion analyzed (S). Source: Richard Sheppard, USGS.

CBIoptIlotIte Chno-BCk

Locality:Webb County, Texas, USA. Descri@ion: Soil. A chnoptilohte mineral separate from a depth of 66 to 89 cm (ZEO8403TXBCk) by D.W. Ming. The soil also

I386

H. R. K&son contains quartz, feldspar. calcite smectite (S). Source: Douglas W. Ming, NASA.

Clino-CBk2

and R. N. Clayton places the laumontite is intergrown with calcite. Crystallization order: laumontite 2 calcite. (H).

and

Locality: Webb County, Texas, USA. Description: Soil. A clinoptilolite mineral separate from a depth of 104 to 137 cm (ZEO-8403TX-CBk2) by D.W. Ming. See description for Clino-BCk (S). Source: Douglas W. Ming, NASA.

Mordenite IMNH-7556

Localby: Thorpsfjara (beach), SttivartjGrdur, S. Mlilas+sla, SW Iceland. Description: Microcrystalline, white fibrous substance consisting of minute needles (thickness 5 30 pm) (H).

Laumontite IMNH-6162

Localitv: Stein&m, Hnaukar. Alftafiiirdur. S. L&las$sla, SE Iceland. Description; Vug or vein of crumbly euhedral crystal laths, up to -5 X I X 1 mm. ln

Natrolite FMNH-9257

Locality: Edgewater, New Jersey, USA. Description: See description for analcime FMNH-M9257 (H).