Contributions to Mineralogy and Petrology

Contrib Mineral Petrol (1987) 97:43 50

9 Springer-Verlag1987

An experimental investigation of heulandite-laumontite equilibrium at 1000 to 2000 bar P f l u i d M. Cho*, S. Maruyama, and J.G. Liou Department of Geology, Stanford University, Stanford, CA 94305, USA

Abstract. The univariant reaction governing the upper stability of heulandite (CaA12SiTOls '6H20), heulandite = laumontite + 3 quartz + 2 H 2 0

CaA12Si7018 96 H 2 0 = CaA12Si4Oa2" 4 H 2 0 heulandite laumontite

(1),

has been bracketed through reversal experiments at: 155+6 ~ C, 1000 bar; 175+6 ~ C, 1500 bar; and 180+8 ~ C, 2000 bar. Reversals were established by determining the growth of one assemblage at the expense of the other, using both XRD and SEM studies. The standard molal entropy of heulandite is estimated to be 783.7+16 J mo1-1 K -1 from the experimental brackets. Predicted standard molal Gibbs free energy and enthalpy of formation of heulandite are - 9 7 2 2 . 3 + 6 . 3 kJ tool -1 and - 1 0 5 2 4 . 3 + 9 . 6 kJ mo1-1, respectively. The reaction (1), together with the reaction, stilbite = laumontite + 3 quartz + 3 H20, defines an invariant point at which a third reaction, stilbite = heulandite + H20 , meets. By combining the present experimental data with past work, this invariant point is located at approximately 600 bar and 140 ~ C. Heulandite, which is stable between the stability fields of stilbite and laumontite, can occur only at pressures higher than that of the invariant point, for Prho = Pro,a1. These results are consistent with natural parageneses in low-grade metamorphic rocks recrystallized in equilibrium with an aqueous phase in which an2o is very close to unity.

Introduction Heulandite occurs commonly as a vein or amygdaloidal mineral and as an alteration product of volcanic glass and calcic plagioclase in low-grade meta-volcanics and metaclastics. Heulandite + quartz ( • albite) constitute a mineral assemblage characteristic of the zeolite facies, together with laumontite-quartz and analcime-quartz (Coombs et al. 1959). A depth zonation of heulandite through laumontite to prehnite-pumpellyite facies assemblages in thick burial volcanigenic sequences has long been established (e.g., Co ombs 1954), and suggests that heulandite becomes unstable relative to laumontite with increasing temperature and pressure. Thus, the upper stability of heulandite may be governed by the simplified reaction, * Present address: Department of Earth and Space Sciences, UCLA, Los Angeles, CA 90024, USA Offprint requests to: M. Cho

+ 3 SiO2 + 2 H 2 0 quartz water

(1)

Natural heulandite, however, exhibits continuous solid solution between heulandite and clinoptilolite, (Na,K)6 (A16SiaoO72)-20HzO (Gottardi and Galli 1985). Thus, the Si/A1 and Ca/(Na + K) ratios and H20 contents of natural heulandite may vary significantly. It should be noted, therefore, that the reaction governing the heulandite to laumontite transition in natural parageneses may differ from reaction (1) by variable SiOz and HzO contents in heulandite, and by the participation of Na- or K-bearing phases. These zeolites, together with anorthite and common low-grade metamorphic C a - A 1 silicates (e.g., lawsonite, wairakite, yugawaralite, and stilbite), have CaO/AI203 = 1. Their compositions and abbreviations are listed in Table 1. Stability relations among these C a - A 1 silicates in the system CaAlzSi/O8- S i O 2 - H 2 0 have been extensively studied (Liou 1970, 1971 a, b; Zeng and Liou 1982). However, the stability of heulandite has not been experimentally investigated except for some reconnaissance experiments at 2 kbar Pn2o by Thompson (1971). Reaction (1) was experimentally located to establish the phase equilibria among heulandite, laumontite and stilbite. The entropy and volume changes of this reaction are very small, in contrast to common dehydration reactions (see below). This not only results in an extremely slow reaction rate, but also large uncertainties in the calculated thermody-

Table 1. Mineral compositions and abbreviations for phases considered in this study Mineral

Formula

Symbol

Anorthite Heulandite Kaolinite Laumontite Lawsonite Quartz Stilbite Wairakite Yugawaralite

CaA12SizO8 CaA12Si7018" 6 H20 A12Si205 (OH)4 CaA12Si4Olz "4H20 CaA12Si2Ov(OH)2HzO SiO2 CaA12SiTOI8-7H20 CaAlzSi4012" 2 H20 CaAla Si6016" 4 HzO

An Hu Ka Lm Lw Qz St Wr Yu

44

Table 2. Chemical compositions of natural laumontite and heulandite used as starting materials" Laumontite

Heulandite

SiOz A1203 FeO b BaO SrO CaO NazO K20

51.61 • 0.38 21.45 • 0.28 0.02 • 0.03 0.02 • 0.02 0.03 + 0.03 11.69 • 0.23 0.07+0.05 0.12 • 0.04

59.10 • 0.57 16.33 • 0.39 0.03 • 0.03 0.14 • 0.04 0.31 • 0.07 7.32 • 0.29 0.82_+0.21 0.68 • 0.21

Total

85.02 • 0.63

84.73 • 0.56

Si A1 X Fe Ba Sr Ca Na K Z"

O=12 4.025 1.972 5.997 0.001 0.001 0.001 0.977 0.011 0.012 1.003

O=18 6.824 2.230 9.054 0.003 0.006 0.021 0.906 0.092 0.050 1.078

ao bo co fl

14.733 • 0.012 13.067 • 0.011 7.571 + 0.004 112.08 • ~

17.708 • 0.004 17.889 • 0.004 7.420 • 0.002 116.41 +0.0l ~

" A few grains of laumontite and heulandite were analyzed by microprobe to check the chemical homogeneity. One standard deviation of 15 individual analyses are listed b All iron as ferrous iron

namic properties. Experiments with long d u r a t i o n up to a b o u t 10 m o n t h s were conducted to overcome the kinetic problems. The t h e r m o d y n a m i c p a r a m e t e r s o f heulandite are calculated from the experimental d a t a a n d c o m p a r e d to the recent calorimetric determination by Johnson et al. (1985). The present experimental results are c o m p a t i b l e with parageneses o f Ca-zeolites in low-grade m e t a m o r p h i c rocks, a n d can be used to estimate the P - T conditions for the f o r m a t i o n o f zeolite-facies mineral assemblages.

Experimental methods Reaction (1) was investigated in the presence of excess HzO employing conventional hydrothermal apparatus and procedures. Details of the experimental procedures are described by Liou (1971 b) and Zeng and Liou (1982). Natural heulandite and laumontite from Tanzawa Mountains, Japan and synthetic ~-quartz were used for starting materials. The compositions and cell parameters of heulandite and laumontite used as starting materials are listed in Table 2. Mineral compositions were determined by an automated JEOL 733 electron microprobe, following the methods described by Cho et al. (1986). Laumontite has nearly stoichiometric composition, whereas heulandite contains minor amounts of Na and K. The individual starting minerals were ground and examined by X-ray diffraction (XRD) to ensure their purity. The charges for each experiment consisted of mixtures of reactants and products in subequal proportion with excess water. They were sealed in Ag or AgvoPdao tube (0.3 mm outer diameter and 0.01 mm wall thickness) with 10-20 mg H20. Reaction (1) was investigated at Pfluia=1000, 1500 and

2000 bar, and temperatures of 150~ ~ C. The duration of experiment varied from 2 to 9.8 months. Fluid pressure was maintained within • 50 bar. The uncertainty in run temperatures is a combination of the variation of the day-to-day fluctuations in recorded temperatures ( • 0.5 ~ to • 2~ C), thermal gradient over the capsule length ( + 2 ~ C), and the uncertainty in thermocouple calibration ( • ~ C); the total uncertainty was ___4.5~ C to • ~ C. Experimental products were examined by an automated Rigaku X-ray diffractometer (XRD), scanning electron microscope (SEM) with energy dispersive X-ray analyzer (EDAX), and petrographic microscope. Reaction direction was deduced by observing the growth of one assemblage at the expense of the other. The reaction direction was determined by comparing the relative intensities of several XRD peaks (e.g. Huo2o, Lmllo, and Qz~ol) of run products with those of starting mix. Hence, reaction reversal was demonstrated. Most run products were also examined using SEM/EDAX which revealed growth or dissolution textures of each solid phases (e.g., Cho and Fawcett 1986). This corroborated the XRD determination of the reaction direction.

Experimental results The rate o f reaction (1) was extremely slow; the extent o f reaction, especially at P - T conditions near equilibrium, was very small even in the experiments with durations o f 2-3 months. Reaction textures

Representative S E M p h o t o m i c r o g r a p h s illustrating the growth or dissolution textures o f heulandite and laumontite are given in Fig. 1. W h e n heulandite and laumontite occurred as very fine grained mixtures, they were difficult to identify even with an E D A X analysis due to the similarity in their habit. Nevertheless, the growth and dissolution textures o f heulandite were useful for determining the reaction direction. Newly grown heulandite is euhedral, typically tabular parallel to (010) (Fig. l A). Overgrowths o f heulandite along its cleavage direction were rarely observed. The growth texture o f heulandite was c o m m o n l y not so conspicuous as its dissolution texture, p r o b a b l y due to the slow growth rate at low temperatures ( < 170 ~ C). The characteristic features o f heulandite dissolution include the f o r m a t i o n o f dissolution pits and steps as well as general rounding o f heulandite crystals (Fig. 1 B-D). Large crystals o f heulandite often recrystallize and form finer grains o f less than I g m size in length (Fig. 1 B and D). In the same experimental charge, M-14 o f Fig. 1 B, small crystals o f dissolving heulandite a p p e a r with r o u n d e d corners and irregular surface steps, lacking in the angular form of growing heulandite (Fig. 1 C). W h e n the dissolution is well advanced, linear etch pits develop along cleavage planes, resulting in a ragged surface m o r p h o l o g y (Fig. 1 D). These textures m a y suggest that the dissolution o f heulandite is controlled by surface reaction (Berner 1981). L a u m o n t i t e and quartz grains are less c o m m o n l y encountered than heulandite during the S E M examination o f the experimental charges due to their low abundance (laumontite) a n d / o r finer grain size (quartz) in the starting mix. N o extraneous phases were detected by S E M observation except very m i n o r calcite on the surface o f three experimental charges. It is inferred from the occurrence o f calcite that trace CO2 m a y have c o n t a m i n a t e d the experimental charge during sample p r e p a r a t i o n (see below for further discussion).

45

Fig. 1A-D. SEM photomicrographs showing the surface morphologies of heulandite and laumontite in the experimental charges. Scale bars represent 1 gin. A (Run No. M-t2) Growing crystals of euhedral heulandite with laumontite, which dissolved particularly along cleavage. B (Run No. M-14) Surface morphology of large seed crystal of heulandite dissolving and recrystallizing into smaller prismatic crystals. C (Run No. M-14) Dissolving grains of heulandite together with growing laumontite. Note the irregularities in the dissolution steps of heulandite indicated by arrow. D (Run No. M-3) Dissolution features of henlandite. Heulandite is rounded and etched along cleavage by dissolution, resulting in the steps on the (h01) plane. Note also the finer crystals of dissolving heulandite on the (010) plane When laumontite or quartz crystals were found, their morphologies always give a reaction direction consistent with that deduced from the heulandite textures. In Fig. 1 A, laumontite has dissolved along its cleavage, whereas heulandite exhibits euhedral growth morphology. On the other hand, Fig. 1 C shows the juxtaposition of a growing crystal of euhedral laumontite with dissolving heulandite. Thus, reaction reversal is proven from the reaction textures. Results

The experimental results for the reaction (1) are summarized in Table 3 and Fig. 2. Both XRD and SEM observations were used to locate the equilibrium at: 154.7 • 5.8 ~ C, 1000bar; 175.2•176 1500bar; and 179.7__7.6~ 2000 bar Pnuld (Fig. 2). The uncertainties in the bracketing data between two assemblages were calculated using the equation (6) of Demarest and Haselton (1981). Figure 2 also shows the dehydration curve of reaction (1) estim/tted by Thompson (1971). His calculated curve has a negative slope, based on his reconnaissance experiments at 2 kbar Pn2o- Thompson's curve, however, may be suspect because the short duration (two weeks) resulted

in only a tiny extent of reaction due to slow reaction rate. Furthermore, analcime was found in one of his runs and such occurrence indicates that his starting heulandite may contain significant amounts of sodium. However, Thompson (1971) gives no analytical data of his starting material. Heulandite used in this study contains small amounts of ( N a + K) and shows a smaller Si/A1 ratio than ideal CaA12SivO18-6H20. Breakdown of heulandite in the experiments was not strictly governed by the reaction (1), and may have involved other minor phases. In fact, small amounts of kaolinite were identified by XRD along with heulandite, laumontite, and/or quartz in some runs at both high and low temperatures (see Table 3). Therefore, the overall reaction defining the upper stability of natural heulandite may be approximately written from the starting compositions of heulandite and laumontite as: Hu = Lm + Qz + Ka + fluid. The occurrence of minor kaolinite may also be attributed to the possible incongruent dissolution of heulandite and/or laumontite. Incongruent dissolution of feldspars is well-documented from both field and experimental observations (e.g., Allen and Fawcett 1982; Anderson and Burn-

46 Table 3. Representative run data for the reaction (1), heulandite = laumontite + 3 quartz + 2H20 Run No.

P T (kbar) (~ C) 149.9__+5.2

264

M-17 b 1.0

159.5___4.7

244

170.0-1-4.6 179.8_+4.2

244 245

M-13 b 1.0

200.3•

245

S-9 M-1 S-1 M-22 ~ M-2 M-3

150+6.0 170.4__+4.2 180_+6.0 190.3• 209.8_+4.6 220.3___4.4

64 293 64 209 293 293

1.0 1.0

1.5 1.5 1.5 1.5 1.5 1.5

S-12 M-23 ~ M-11 S-11

2.0 2.0 2.0 2.0

150+__6.0 169.4_+4.7 189.9___4.4 200+5.5

M-19 S-IB S-10

2.0 2.0 2.0

210.8+4.3 220+6.0 250• 5.5

74 98 121 74 98 59 74

Hu(+), Lm(+), Ka(+) Hu(-), Lm(+), Ka(+) Hu(-), Lm(+), H u ( - ) , Lm(+), Ka(+) H u ( - ) , Lm(+), Hu(+), Lm(-), Hu(+), Lm(--), H u ( - ) , Lm(• Hu(-), Lm(+), H u ( - ) , Lm(§ H u ( - ) , Lm(+), Ka(+),

I

I

%o. o--\\ 9

Qz(-),

Qz(+) Qz(+), Qz(+) Qz(-) Qz(-) Qz(+) Qz(+) Qz(_) Qz(+),

Hu(+), Lm(-), Qz(+) Hu(+), Lm(-), Qz(• H u ( - ) , Lm(+), Qz(+) Hu(• Lm(+), Qz(• Ka(+) Hu(--), Lm(+), Qz(+) Hu(-), Lm(_+), Qz(+) Hu(-), Lm(• Qz(+)

Growth or diminution of a phase is indicated by (+) or ( - ) , respectively. Note that some experimental charges contain kaolinite (see text for further discussion) b Runs in which minor calcite grains are observed with the SEM Reground experimental charges of M-14 and -17 were used as starting material for runs M-22 and -23, respectively

ham 1983). Although no solubility measurements ofheulandite and laumontite are available, it is likely from the structural similarities between feldspars and zeolites that heulandite and/or laumontite may also dissolve incongruently. Thus, Ca and minor alkali cations of heulandite or laumontite may have dissolved faster than the framework A1 and Si cations, leading to the crystallization of kaolinite. The occurrence of minor calcite crystals on the surface of some experimental charges further supports this interpretation: Calcite may have precipitated from a Ca2+-rich fluid. It is beyond the scope of this study to determine the mechanism responsible for the formation of kaolinite. In the following thermodynamic analysis, kaolinite was ignored because it was observed in negligible amounts only in several experimental products. Discussion

Thermodynamic considerations The entropy of heulandite was calculated using the method described by Zen (1971) from three reversal brackets: (1) 154.7• ~ C, 1000 bar; (2) 175.2_+6.5 ~ C, 1500 bar; and (3) 179.7• ~ C, 2000 bar Prlu~d (Fig. 2). All thermodynamic data employed in this study are listed in Table 4. Three independent estimates of the standard molal entropy of formation of heulandite (A ~,n,) are obtained from each pair of equilibrium P - - T brackets: (1, 2); (1, 3); and (2,

o-

/

x\4 Lm+Qz+H20

n-

Qz(+),

I

This Study

t Remarks" (days)

M-12 b 1.0

M-16 9 M-14

2.5

eD

, ,/o'o, Hu

n

1.o 0.5

1O0

/

oo \NN\

7000 i

150

i

200

i 250

TEMPERATURE,~ Fig. 2. Pfluid- T diagram for the reaction (i): heulandite = laumontire + 3 quartz + 2H20. Solid circles refer to the growth of heulandite, and open circles to the growth of high-temperature assemblage. Also plotted in dashed curve are the equilibrium conditions of the reaction (1) determined by Thompson (1971) 3). The respective ranges are -2682.2 to -2698.0; -2671.0 to -2693.8; and -2644.0 to -2702.7 J mo1-1 K - 1. These values overlap each other. The large difference between the last two values is apparently due to the similar equilibrium temperatures at 1500 and 2000 bar (Zen 1985). A mean of the first two entropy values, -2690.1 _+16 (2a) J moV 1 K - a , is adopted in this study. The standard molal entropy of heulandite (S~,) is calculated as 783.7• tool 1 K-1, using the entropy data for Ca, A1, Si, 02 and H2 compiled by Helgeson et al. (1978). This entropy value is 20.5 J mol 1 K - 1 larger than S~u ( = 763.2 J mol 1 K - 1) estimated from the oxide sum method by Helgeson et al. (1978), and 16.5 J mol-a K - 1 larger than S~, ( = 767.18 • 0.77 J tool 1 K - 1) calorimetrically determined by Johnson et al. (1985). From the estimated value of A S~, H, in conjunction with Zen's (1971) method, internally-consistent equilibrium P - - T conditions for reaction (1) were calculated: 159•176 1000bar; (2) 172•176 1500bar; and (3) 187• 8~ C, 2000 bar. The calculated equilibrium curve is shown in Fig. 3; its concavity is unlike that of common dehydration reactions. However, the heat capacity functions or compressibilities of zeolites, which are not taken into account in this calculation, can greatly affect the shape of the equilibrium curve (see below). The standard molal Gibbs free energy of formation of heulandite, AG~,.u, is calculated from the method outlined by Fisher and Zen (1971). Three calculated P - Tdata given above yield a consistent value of AG~,Hu= -9722.3 _+6.3 kJ tool-1. From the relation, aG~ = 3 w e - rAS~, the standard molal enthalpy of formation of heulandite (AH~,nu) is estimated to be --10 524.3_ 9.6 kJ mol-1. These thermodynamic parameters of heulandite described above assume that a(AS)/aT and 6(3 V)/OP among solid phases are equal to zero. Further attempts have been made to retrieve the thermodynamic parameters of heulandite by incorporating the heat capacity function of each

47 Table 4. Thermodynamic data for laumontite, quartz, stilbite, and heulandite Laumontite

Quartz

Formula CaA12Si4012.4H20 S i O 2 2.2688 V, J mol-1 bar-1 20.755 AS~, J tool -1 K -1 --1 850.2 182.51 41.34 S~, J m o l - l K -1 485.8 AG~, kJ mol -I 6682.03 856.239 AH~f, kJ mol -I -7233.65 -910.648

--

Ca, coefficients a b c Source

515.47 186.06 68.74

Stilbite

Heulandite

CaA12SivO18-7H20 32.869 b 31.637 --2710.7 828.3 c 763.2

CaA12SiTO18"6HzO 31.06 31.696 --2734.9 -2689.9 767.18 783.7 --9675.6 -9722.3 -10491.0 --10524.3

46.94 34.31 11.30

Helgeson et al. (1978)

Helgeson et at. (1978)

751.70 288.99 102.63 This study

Helgeson et al. (1978)

745.55 651.33 141.08 Johnson et al. (1985)

This study

" Cp(Jmo1-1K 1 ) = a + b x l 0 - 3 T - c x l 0 - S T - 2 b Calculated from cell parameters listed in Liou (1971 a) ~ Calculated using epistilbite as a structural analog from the Eq. (75) of Helgeson et al. (1978)

2.5

I

2.0

I

/ Lm+Qz+H2O ./

I

/

,..Q

i 1.5

1.0 /~-/ Hu //~ 0.5

I 100

[ 150 200 TEMPERATURE,~

Fig. 3. Calculated dehydration curves for the reaction (1). Open boxes represent uncertainty ranges in the equilibrium pressures and temperatures of the reaction (1) determined in this study. Dashed curve corresponds to the calculated equilibrium condition employing the method described in Zen (1971). Solid curves represent the calculated P - T curves using the SUPCRT computer program and its database (Helgeson et al. 1978) by employing 2000 bar (A), 1500 bar (B), 1000 bar brackets (C), respectively phase. However, neither regression nor linear programming technique (see Berman et al. 1986, for review) yielded a data set consistent with the experimental results. F o r example, AG~,nu was calculated from three P - T data points to be - 9725.2 kJ mol ~ employing the S U P C R T computer program and its data base (Helgeson and Kirkham 1974; Helgeson et al. 1978). However, when the experimental brackets and retrieved thermodynamic parameters of heulandite are used to generate the univariant curve for reaction (1), the individual bracketing data do not give a curve that passes through the other brackets (Fig. 3). On the other hand, when the linear programming technique is employed in conjunction with the thermodynamic data listed in Helge-

son et al. (1978) or Berman et al. 0985), the reversal data at three different pressures did not yield an overlap in the entropy vs. enthalpy space (Day et al. 1985). These inconsistencies between the experimental data and the available thermodynamic data may be attributed to (a) the very small values of A S and A V for the reaction (1) (they are in the order of 5 J mol 1 K - 1 and 0.3 J tool- 1 bar 1, respectively, at 175~ and 1500 bar); (b) the large uncertainties in the entropy and heat capacity data of heulandite and laumontite; and (c) the unknown pressure dependency of volume and/or composition of zeolites. Johnson etal. (1985) recently calculated AG~,Hu ( = -- 9722.2 kJ m o l - 1) from our preliminary experimental data (Maruyama et al. 1983) by combining their own calorimetric data of heulandite with the auxiliary data for laumontite and quartz from Helgeson et al. (1978) and H 2 0 from Helgeson and Kirkham (1974). However, the calorimetric data for S~u by Johnson et al. (1985) are greater than the entropy sum of the assemblage, L m + 3 Q z + 2 H 2 0 , at temperatures above 25 ~ C. Therefore, heulandite must be a high-temperature mineral with respect to the assemblage, Lm + Qz + H 2 0 , which contradicts both the experimental results and natural paragenetic relations. It is thus concluded that the calorimetric data o f ~ or Cp,Hu determined by Johnson et al. (1985) cannot be combined with the data set listed in Helgeson et al. (1978). This apparent discrepancy between the two data sets may indicate that the entropy or heat capacity function o f laumontite in Helgeson et al. (1978) needs to be revised by future calorimetric measurements. P - T stability relations among Ca-zeolites In the system C a A I ~ S i 2 O s - S i O 2 - - H 2 0 , heulandite, laumontite and stilbite together with quartz and H 2 0 define an invariant point (I0, from which the following four univariant reactions radiate: (St) (Hu)

heulandite = laumontite + 3 quartz + 2 HzO

(1)

stilbite = laumontite + 3 quartz + 3 H 2 0

(2)

48 A

I

2

I

I

I

/

@ ..Q

/ 1,2

u3

rr

/

cO [J.I

09 LU

,w

no. 1

--

2

0-

3

1

0 100

I 150

I 200

I 250

I 300

2

100

]

I

I

200

I

300

TEMPERATURE, ~ B

I

3

I

I

I

I 200

I 250

I 300

Fig. 5. A P n u i a - - T diagram showing the relationship among four univariant reactions (1) to (4), intersecting at an invariant point I1. Open boxes represent the uncertainty in the equilibrium pressures and temperatures for reaction (1) determined in this study. Brackets denote the experimental results of the stilbite/laumontite reaction (2) determined by Liou (1971 a). Note that all of his brackets represent the metastable equilibrium. The equilibrium dehydration curves relating laumontite, wairakite, yugawaralite and lawsonite are also shown (from Liou 1971a, b; and Zeng and Liou 1982). 12 refers to the invariant point where heulandite, laumontite, lawsonite, quartz and fluid coexist

1 Hu

2 St

E

i 2 0 100

/,/i I 150

TEMPERATURE, oC

Fig. 4A, B. Possible phase relations among three dehydration reactions (1) to (3), governing the stabilities of stilbite, heulandite and laumontite in the presence of excess quartz and H20 (modified from Liou 1983). A combination of the experimental results of Liou (1971 a) and Thompson (1971) gives (A). However, an alternative topology (B) is preferred in this study, based on the experimental results for the reaction (1). Note that heulandite is stable only at low pressures in A, and at high pressures in B. The univariant curve for the reaction, laumontite = wairakite + 2H20, is also plotted for reference (from Liou 1971 b)

(Lm), (Qz) stilbite = heulandite + H 2 0 (H20)

3 heulandite = 2 stilbite + laumontite + 3 quartz.

(3) (4)

The phases in th e parentheses are the absent phases following Schreinemakers' notation. Two possible geometric relations a m o n g reactions (1) to (3), assuming excess quartz and H 2 0 , have been qualitatively determined by Liou (1983) (Fig. 4). Figure 4 A indicates that the transition from stilbite to laumontite would take place at higher pressure than that from stilbite through heulandite to laumontite. However, such relations are inconsistent with occurrences o f zeolite minerals recorded in the literature (e.g., C o o m b s 1954; Seki et al. 1969), and the converse (Fig. 4B) is true as suggested by Liou (1983). The present experimental data support the relationships in Fig. 4 B, as discussed below. Our experimental results on reaction (1) were combined with experimental data on reaction (2) determined by Liou

(1971a) to establish the P - T relations among the above four univariant reactions (Fig. 5). The equilibrium temperature o f reaction (1) is about 15 ~ C higher than the metastable equilibrium temperature of reaction (2) at 2000 bar. The invariant point /1, where laumontite, stilbite, heulandite, quartz and fluid coexist, is defined by an intersection o f the experimentally determined curves for reactions (1) and (2), and is located at approximately 600 bar and 140 ~ C. Thermodynamic data for heulandite listed in Table 4 were used to calculate a univariant curve for the H20-conservative reaction (4). The calculated slope ranges from /4 to 47 bar/K, assuming that volume and entropy of the solidsolid reaction (4) are independent o f pressure and temperature, respectively. A value of 14 bar/K, based on the thermodynamic data of heulandite by Helgeson et al. (1978), is adopted in Fig. 5 in order to satisfy the topological constraints. Similarly, a steeper positive slope of the reaction (3) is obtained from the same data set together with the H 2 0 properties at 1000 bar and 150 ~ C. The P - - Trelations for the four univariant reactions shown in Fig. 5 indicate that at the condition of Pn2o = P t o t a l , heulandite is stable only at pressures higher than 600 bar and temperatures between the stability fields of stilbite and laumontite. With decreasing PH20/Ptotal, the invariant point shifts toward lower pressures and lower temperatures along the H20-conservative reaction (4). Thus, heulandite may occur stably at pressures less than 600 bar, when Pn2o < Pto,a~ (e.g., in fracture systems or in environments where the other components are present in the fluid). On the other hand, the invariant point will migrate along the univariant curve for reaction (3) with decreasing asio2. Thus, the occurrence of heulandite may be restricted to pressures higher than 600 bar at asao~< 1. The effect of other components on phase relations can be deduced from the chemistry and occurrence of natural Ca-zeolites. Natural heulandite and stilbite deviate more

49 from their ideal compositions than does laumontite. Thus, when variations in K and Na together with the variations in Si/A1 ratios are taken into account, (a) the system has more than 3 components; (b) invariant point I1 and the univariant curves of Fig. 5 become multivariant; and (c) the heulandite stability field will shift toward higher temperature and possibly toward lower pressure. The latter displacement may be necessary to account for the common occurrence of non-stoichiometric heulandite or clinoptilolite at very low pressures in ocean-floor metabasalts or deep sea sediments (e.g., Iijima 1978; Klein and Lee 1984). The stability relations of stilbite/laumontite (Liou 1971 a), laumontite/wairakite/lawsonite (Liou 1971 b), heulandite/lawsonite (Nitsch 1968), and yugawaralite/laumontite/wairakite (Zeng and Liou 1982) have been previously studied. The results are summarized in Fig. 5. The extrapolation of the equilibrium reaction (1) towards higher P and T defines another invariant point (I2) for the coexistence of heulandite, laumontite and lawsonite in the presence of quartz and fluid. It occurs at about 200 ~ C and 3000 bar, where the lawsonite/laumontite curve of Liou (1971 b) and the heulandite/lawsonite curve of Nitsch (1968) intersect with the reaction (1). Again, both laumontite and lawsonite are known to have nearly stoichiometric compositions. Substitutions of N a and/or K for Ca in natural heulandite would expand its stability field toward higher temperatures at Prho = P t o t a l .

found in the Reydarfjordur drill core sample, Eastern Iceland (Mehegan et al. 1982). Natural heulandite commonly contains Na or K or both substituted for Ca, and its chemical composition may extend to those of clinoptilolite and alkali-clinoptilolite (Boles 1972; Gottardi and Galli 1985). In the silicic tuff sequences of Niigata oil field, Japan, Na- or K-rich clinoptilolite is replaced by Ca-clinoptilolite or heulandite with a transition zone, in which these heulandites occur together (Iijima and Utada 1972; Iijima 1978). With increasing depth, laumonrite (+analcime) replaces heulandite. Such a systematic change of composition and the existence of a transitional zone suggests a continuous reaction relationship between clinoptilolite and heulandite. Iijima (1978) has estimated from the measured temperatures of deep wells that the transition between Ca-rich heulandite and laumontite occurs at about 100 ~ C and at depths ranging from 2000 to 3000 m. This temperature estimate is approximately 40~ lower than the experimentally-determined temperature of the present study. The lower temperature of reaction may be attributed to PH20
Geological applications

Conclusions

The regional occurrences of zeolites in low-grade metamorphic rocks have been extensively described, particularly from the Circum-Pacific orogenic belts after the establishment of the zeolite facies by Fyfe et al. (1958) and Coombs et al. (1959). Heulandite is one of the common Ca-zeolites in the Taringatura and Hokonui Hills, New Zealand, and is replaced by laumontite plus quartz with increasing burial depth (Coombs 1954; Boles and Coombs 1975). The depth zonation of heulandite~laumontite~pumpellyite and prehnite assemblages has been documented in burial sequence of andesitic tuffs. On the other hand, a distinct zonal distribution of stilbite ( _ clinoptilolite or heulandite)~laumontite (+wairakite or yugawaralite)--,prehnite~actinolite~ hornblende has been well established in similar rocks (e.g., Green Tuff formation) with increasing proximity to a hypabyssal intrusive in the Tanzawa Mountains, Japan (Seki et al. 1969). Such occurrences support the topologic relations among stilbite, heulandite and laumontite as shown in Figs. 4B and 5 in which the zonations observed in Tanzawa aureole can be explained by increasing temperature at very low pressure. Ishizuka (1985) has recently described the progressive mineral assemblages ranging from the zeolite facies to the granulite facies in the Horokanai ophiolite, Japan. The zeolite facies rocks in the upper sequence are subdivided into three zones by the sequential appearance of chabazite + stilbite ( + analcime), laumontite ( + albite), and wairakite (+albite). From the P - T relations shown in Fig. 5, the direct transition from stilbite to laumontite without a heulandite zone further supports the low-pressure (ocean-floor) metamorphism of the Horokanai ophiolite. Similar depth sequence of stilbite--,laumontite--~wairakite has been documented in active Iceland geothermal system (e.g., Kristmannsdottir 1982), although rare heulandite grains are

P - T conditions of equilibrium for the reaction: heulandite = laumontite + 3 quartz + 2 H 2 0

(1),

were experimentally determined and govern the upper stability of heulandite with a composition (Cao.9o6Bao.oo6 Sro.o21Feo.oo3Nao.o92Ko.oso) A12.23oSi6.83sO18"6H20. Attempts to generate an internally-consistent thermodynamic dataset of heulandite based on experimental brackets were not successful primarily because of the large uncertainties introduced by the small entropy and volume change of the reaction (1). These results, together with previous experimental studies on Ca-zeolites, are used to establish phase equilibria in the system CaA12Si2Os-SiO2- H20. Phase relations of Ca-zeolites indicate that Ca-rich heulandite is stable only at pressures higher than 600 bar and temperatures between the stability fields of stilbite and laumontite. However, substitution of Na and/or K for Ca, and variations in SIO2, H 2 0 contents, and PH2o/Ptota I ratio may significantly affect the experimentally-determined P - - T stability of endmember heulandite. Nevertheless, the experimentally determined P - T relations are consistent with those deduced from natural parageneses. The direct transition from stilbite to laumontite without the heulandite zone may be favored at low pressures, such as Tanzawa Mountains, Japan, active geothermal systems, and ocean-floor metamorphism. On the other hand, the transition from heulandite to laumontite such as in New Zealand may be typical in burial sequences at higher pressures.

Acknowledgements. This research was supported by NSF EAR 82-04298 and 85-07988. We wish to thank Drs. Rona Donahoe and Yotaro Seki for their helpful comments. Thoughtful and constructive reviews by Drs. D.K. Bird, J.V. Chernosky, H.W. Day,

50 J.M. Ferry and M.F. Hochella, Jr. have greatly improved an earlier draft. Thanks are also due to Dr. W.S. Wise for generously providing a preprint prior to its publication.

References

Allen JM, Fawcett JJ (1982) Zoisite - anorthite - calcite stability relations in H20 - COa fluids at 500 bars: An experimental and SEM study. J Petrol 23:215-239 Anderson GM, Burnham CW (1983) Feldspar solubility and the transport of aluminum under metamorphic conditions. Am J Sci 283-A: 283-297 Berman RG, Brown TH, Greenwood HJ (1985) An internally consistent thermodynamic data base for mineral in the system Na20 - K 2 0 - C a O - MgO - FeO - F e 2 0 3 - - A 1 2 0 3 - S i O 2 T i O 2 - H 2 0 - C O 2 : representation, estimation, and high temperature extrapolation. Atomic Energy Canada Ltd. Tech Report TR-377, 62 p Berman RG, Engi M, Greenwood HJ, Brown TH (1986) Derivation of internally-consistent thermodynamic data by the technique of mathematical programming: a review with application to the system M g O - S i O 2 - H 2 0 . J Petrol 27:1331-1364 Berner RA (1981) Kinetics of weathering and diagenesis. In: Lasaga AC, Kirkpatrick RJ (eds) Kinetics of geochemical processes. Rev Mineral 8 : 1 1 1 - 1 3 4 Boles JB (1972) Composition, optical properties, cell dimensions, and thermal stability of some heulandite group zeolites. Am Mineral 57:1463-1493 Boles JB (1977) Zeolites in low-grade metamorphic rocks. In: Mumpton FA (ed) Mineralogy and geology of natural zeolites. Rev Mineral 4:103-136 Boles JR, Coombs DS (1975) Mineral reactions in zeolitic Triassic tuff, Hokonui Hills, New Zealand. Geol Soc Am Bull 86:163-173 Cho M, Fawcett JJ (1986) Morphologies and growth mechanisms of synthetic Mg-chlorite and cordierite. Am Mineral 71:78-84 Cho M, Liou JG, Maruyama S (1986) Transition from the zeolite to prehuite-pumpellyite facies in the Karmutsen metabasites, Vancouver Island, British Columbia. J Petrol 27:467-494 Coombs DS (1954) The nature and alteration of some Triassic sediments from southland, New Zealand. Trans R Soc NZ 82 : 65-109 Coombs DS, Ellis AJ, Fyfe WS, Taylor AM (1959) The zeolite facies, with comments on the interpretation of hydrothermal syntheses. Geochim Cosmochim Acta 17 : 53-107 Day HW, Chernosky JV, Kumin HJ (1985) Equilibria in the system M g O - S i O z - H 2 0 : a thermodynamic analysis. Am Mineral 70 : 237-248 Demarest HH, Jr., Haselton HT, Jr. (1981) Error analysis for bracketed phase equilibrium data. Geochim Cosmochim Acta 45: 217-224 Fisher JR, Zen E (1971) Thermochemical calculations from hydrothermal phase equilibrium data and the free energy of H20. Am J Sci 270:297-314 Fyfe WS, Turner FJ, Verhoogen J (1958) Metamorphic reactions and metamorphic facies. Geol Soc Am Mem 73:259 Gottardi G, Galli E (1985) Natural zeolites. Springer-Verlag, New York Berlin Heidelberg, 409 p

Helgeson HC, Kirkham DH (1974) Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures. I. Summary of the thermodynamic/ electrostatic properties of the solvent. Am J Sci 274:1089-I 198 Helgeson HC, Delany JM, Nesbitt HW, Bird DK (1978) Summary and critique of the thermodynamic properties of rock-forming minerals. Am J Sci 278-A: 1-229 Iijima A (1978) Geological occurrences of zeolite in marine environments. In Sand LB, Mumpton FA (eds) Natural zeolites: occurrences, properties, use. Pergamon Press, New York, pp 175-198 Iijima A, Utada M (I972) A critical review on the occurrence of zeolites in sedimentary rocks in Japan. Jpn J Geol Geogr 42: 61-83 Ishizuka H (1985) Prograde metamorphism of the Horokanai ophiolite in the Kamuikotan zone, Hokkaido, Japan. J Petrol 26:391-417 Johnson GK, Flotow HE, O'Hare PAG, Wise WS (1985) Thermodynamic studies of zeolites: heulandite. Am Mineral 70:1065-1071 Klein GdeV, Lee YI (1984) A preliminary assessment of geodynamic controls on depositional systems and sandstone diagenesis in back-arc basins, Western Pacific Ocean. Tectonophysics 102:119-152 Kristmannsdottir H (1982) Alteration in the IRDP drill hole compared with other drill holes in Iceland. J Geophys Res 87 : 6525-6531 Liou JG (1970) Synthesis and stability relations of wairakite, CaA12Si4012' 2H20. Contrib Mineral Petrol 27:259-282 Liou JG (1971 a) Stilbite-laumontite equilibrium. Contrib Mineral Petrol 31 : 171-177 Liou JG (1971 b) P - T stabilities of laumontite, wairakite, lawsonite and related minerals in the system CaA12Si2Os-SiO2H20. J Petrol 12:379-411 Liou JG (1983) Occurrence, compositions and stabilities of some C a - A I hydrous silicates in low-grade metamorphic rocks. Geol Soc China Mem 5:47-66 Maruyama S, Cho M, Liou JG (1983) Experimental investigation of heulandite -laumontite equilibrium. EOS 64:897 Mehegan JM, Robinson PT, Delaney JR (1982) Secondary mineralization and hydrothermal alteration in the Reydarfjordur drill core, Eastern Iceland. J Geophys Res 87:6511-6524 Nitsch KH (1968) Die Stabilit/it yon Lawsouit. Naturwissenschaften 55 : 388 Seki Y, Oki Y, Matsuda T, Mikami K, Okumura K (1969) Metamorphism in the Tanzawa Mountains. J Jpn Assoc Min Petrol Econ Geol 61 : 1-25, 50-75 Thompson AB (1971) Analcite-albite equilibria at low temperature. Am J Sci 271:79-92 Zen E (1971) Comments on the thermodynamic constants and hydrothermal stability relations of anthophyllite. Am J Sci 270:136-150 Zen E (1985) An oxygen buffer for some peraluminous granites and metamorphic rocks. Am Mineral 70: 65-73 Zeng Y, Liou JG (1982) Experimental investigation of yugawaralite - wairakite equilibrium. Am Mineral 67: 937-943 Received October 14, 1986 / Accepted April 13, 1987