Geothermal mineral equilibria

Geothermal

mineral equilibria

WERNER F. GIGGENBACH’

Chemistry Division, Department of Scientific and Industrial Research, Private Bag Petone. New Zealand (Received 11 March

1980; accepred

in

revisedform 24

Ocfober

1980)

Abstract-The dominant reaction determining the chemistry of fluids in a geothermal system of the New Zealand type consists of the conversion of primary plagioclase by CO1 to calcite and clays with log Pco, = 15.26 - 7850/(r + 273.2), temperature t in “C. Subsequent reactions involving secondary minerals control relative CO,-H&contents. The distribution of mineral phases throughout a geothermal system reflects the stepwise conversion of thermodynamically unstable primary phases through a series of intermediate, metastable phases to a thermodynamicatiy stable, secondary assemblage. The relative stabilities of these phases was evaluated on the basis of their solubilities, the least soluble aluminiumsilicate representing the thermodynamically most stable phase under a given set of conditions. Observed assemblages of secondary minerals in geothermal systems represent indicators allowing mineral/fluidinteraction conditions to be evaluated on the basis of multi-component mineral stability diagrams.

THE THERMODYNAMIC STABILITY OF MINERALS IN GEOTHERMAL SYSTEMS

INTRODUCTION

RECENTadvancements in the understanding of geoIn the application of thermodynamic considerthermal rock-fluid interaction processes were discussed by ELLISand MAWON(1977). HELGES~N(1979) ations to the stability of mineral phases in geothermal and by ELLIS(1979) in reviews covering both theoretisystems, extensive use is made of logarithmic activity cal and field studies. It would be far beyond the scope diagrams, as e.g. by HU+SLEY and JONES(1964) in an of this introduction to attempt to present an objective early study into chemical aspects of hydrothermal assessment of the respective merits of the large alteration with emphasis on hydrogen metasomatism, number of investigations on this subject. Two major or by HELGFSON(1967) who extended such diagrams contributions, however. can be singled out as having to explore the chemistry of aqueous solutions associated with metamorphic reactions. ELLIS (1970) and had a major influence on the present study or having made it possible in the first place. BROWNEand ELLIS(1970) used this type of phase diagram to interpret the chemistry of waters discharged By applying thermodynamic equilibrium considerations to reactions linking geothermal solutions, gases from Wairakei and Broadlands geothermal wells in New Zealand. and mineral phases, ELLIS (1969) started to unravel While providing an elegant and convenient means the complex web of interactions thermodynamically connecting temperature, salinity, carbon dioxide of representing thermodynamic information, the stabpressure and occurrence of secondary minerals within ility diagrams now in use are limited to the presena geothermal system. The validity of the approach tation of mineral phase relationships in terms of only was checked by comparing analytical and theoretical two solution components at a time, or three if H+ is activities. a procedure requiring the evaluation of included. In order to be able to correlate the thermoequilibrium constants by use of sufficiently reliable dynamic stability of the wide range of phases making and accurate thermodynamic data for the reactants up a geothermal mineral assemblage generally, repinvolved. This task has recently been greatly facili- resentation in terms of two common parameters tated by HELGE~ONet al. (1978) who provided a critiwould be required. The bulk of these minerals are cally assessed set of such data for most of the reaction aluminium-silicates, their common constituents being participants encountered in geothermal investigations alumina and silica. The two most promising parand a computer program SUPCRT,expediting their ameters to be used in the representation of geotheruse. The present work forms largely an extension or mal solute-mineral relationships. therefore, would application of these techniques and of a previous appear to be based on the activities of their correstudy of geothermal gas equilibria (GIGGENBACH,sponding liquid phase components. By following 1980). The chemistry of ionic solute species present in established procedures, dissolved aluminium activities geothermal fluids will be discussed elsewhere. may then be represented in terms of the activity ratio aAll+Ia&, while silica contents are introduced in * Present address: International Atomic Energy Agency, terms of aHSiO,. In allusion to pH and E,,. this new A-1400 Vienna. Austria. central parameter governing aluminium-silicate equi393 ‘_‘_A

4’

3--H

W. F.GIGGENBACH

.-.-.-.-.-.

,._. -.-.-.

395

Geothermal mineral equilibria

libria, log (aA,,+/a&), is abbreviated Al”, the subscript H reminding the user of the cube of the hydrogen ion activity in the denominator. The stability ofgeothermal minerals in terms of Al, Construction of mineral phase diagrams is most readily carried out by use of equilibrium expressions describing the hydrolysis of a mineral into its solution components. Equilibrium constants for a set of such hydrolysis reactions were reported e.g. by HUGEXIN (1%9). More up to date values used in this investigation were obtained by use of a 1977 version of the program SUPCRTdeveloped by HELGES~Net al. (1978). In spite of the rapid rate at which new, and often probably improved information on the thermodynamic stability of many of the mineral phases referred to in this study becomes available, no attempt was made to update the 1977 data set as a mineral by mineral assessment of the internal consistency and compatibility of these new data with those given in SUPCRTwould have been beyond the scope of this investigation. As updated versions of the SUPCRTdata file become available, it should be a trivial task to revise and to possibly complete the stability diagrams given below. Most of these revisions, however, are likely to be minor and should not affect the major conclusions reached in this attempt to evaluate general processes controlling the distribution of minerals in geothermal systems. In Fig. 1 mineral phase stability diagrams in terms of Al” and the activity of dissolved silica are given for 50, 100, 150, 200, 250 and 300°C at the activity of saturated water vapor. They allow the immediate assessment of the thermodynamic stability of a given mineral assemblage or the composition of an assemblage made up of a wide variety of components to be predicted for a given set of conditions. For reasons of clarity, Fig. 1 is limited to the system Na20, KzO, MgO, CaO, Alr03, SiOl, CO1 and Hz0 but may be extended to include FeO, Hz and HIS. The diagrams are constructed by assuming aRBio, to correspond to the activity of non-specific SiO&q) listed by HEG~ON et al. (1978) and on the basis of idealised phases neglecting the effects of compositional variations and substitutional and structural disorder. An indication of the effects of substitutional disorder is provided by considering variations in Alu resulting from the use of thermodynamic data for microcline and sanidine in the evaluation of the K-feldspar/muscovite boundary and those of low and high albite in the evaluation of the Na-feldspar/paragonite boundary as shown in Fig. 1 for 250°C. For both the disordered potymorphs the coexistence boundary with respect to their corresponding mica shifts to lower Al”-values implying that the disordered feldspar releases alkali ions more readily allowing the corresponding mica to form at lower Al,,-values as compared to the more stable, ordered feldspar. The parameter Al” can, therefore, be considered to present a measure for the relative tendency

of a phase to retain alkali or, in the case of Mg- and Ca-aluminium silicates. earth-alkali ions. Thii correlation between the activities of dissolved aluminium and that of the cation Y is a direct consequence of the relationship nYH + mAlR = Li.q,v

(1)

the logarithmic form of the ‘solubility product’ of aluminium silicates, where Y” = log (ai/&+)

(2)

with Ui representing the activity of the cation Y and zi its charge, and Li,4Vvthe logarithm of the hydrolysis constant of a Y,Al,-silicate for quartz (q) saturated solutions at the activity of saturated water vapor (w). Figure 1 displays stability areas of most of the common geothermal sodium, potassium, magnesium and calcium aluminium-silicate phases. Much of the space is occupied by feldspars, zeolites and micas, the omission of one of the most important groups of geothermal alteration products, the montmorillonites (BROWNE,1978) is due to the fact that thermodynamic data available for Mg- and Ca-montmorillonites (H~zLGE~~N,1969) give rise to montmorillonite/mica coexistence boundaries slightly beyond, and Na- and K-montmorillonite/mica coexistence lines well beyond the ‘kaolinite barrier’ as shown in Fig. 1 for 250°C. This of course, represents a thermodynamically impossible situation implying alkali-montmorillonites to form at AIKvalues above the limit hnposed by the solubility of kaolinite. In the pmsenoz of quartz maximum equilibrium AIR-values in a geothermal system should, for temperatures to around 300”C, be uniquely fixed by the formation of kaolinite, at higher temperatures or higher dissolved silica activities pyrophyllite is likely to limit Al”. A similar situation arises from the use of thermodynamic data quoted for illite (HWXWN, 1969) where muscovite, chlorite/illite-boundaries close to those of the alkalimontmorillonites shown in Fig. 1 are obtained. The data reported for these silicates would imply that montmorillonites are able to form only in association with ahnninium silicates more soluble than kaolinite such as dickite or hahoysite or that the data provided by the two sources (HEU;ESON,1%9 and SUPCXT)are incompatible. presentation of thermodynamic informatron m terms of AlR/log au,sR-,4-diagrams, therefore, may provide a convenient means of weeding out unrealistic and conflicting data, or to derive, at least, estimates of more consistent thermodynamic data by plotting observed stability ranges of undocumented phases in relation to thermodynamically well-established phases. Based on Fig. 1, a series of major boundaries subdividing the Al,+cale along the quartz saturation lines can be distinguished: at low Al,,-values, the clinozoisite/grossular-boundary followed with increasing Alu by the muscovite/microcline coexistence line. At all temperatures paragonite/low albite-and clinochlore/phlogopite-boundaries cross the quartx

W.

396

F. GIGGENBACH

saturation lines at close to common AI”-values. The upper limits of Al” are given by the kaolinite-pyrophyllite (-montmorillonite)-barrier. The stability of laumontite is restricted to temperatures below 260°C and dissolved silica activities exceeding those of quartz dissolution to be replaced at higher temperatures by wairakite as the more stable zeolite. Analtime and mar&e are thermodynamically only stable at SiOz-activities below those of quartz stability. Of the large number of other mineral species within the Na*O, KzO, MgO, CaO, A1203, SiOz, CO*, H20-system listed in SUPCRT none was found to be stable over the range of temperatures, silica activities, and Al+~alues depicted in Fig. 1. Figure 1 also contains contour lines delineating variations in Pco, for assemblages including calcite. Along the quartz saturation line equilibrium Al”-values increase with increasing Pco,. For instance at 250°C grossular becomes unstable at Pco, approaching 0.2 bar to be replaced by clinozoisite. The muscovite/microcline-boundary corresponds to a CO*-pressure of around 1 bar. At CO*-pressures exceeding 2 bar both albite and phlogopite will tel ’ to disappear while Pco, above 10 bar can be expects, to lead to complete cation leaching of the rock matrix with only kaolinite. calcite and quartz remaining. Because of its central role in geothermal mineral phase relationships. direct evaluation of Al” from analytical data would provide an important parameter allowing equilibrium mineral assemblages to be related directly to thermal water compositions (GOGUEL, 1977: PACES 1978). Ana-

nonstoichiometry more or less fixed point, only temperature dependent. mineral buffers of which the pair muscovite/K-feldspar is the one most commonly employed in fixing the pH of a geothermal fluid system, or a sliding scale buffer system with A&+-values related to CO*-partial pressures. As a consequence of the phase rule only one system can at any time exert control relegating the other to the role of indicator. The action of both, buffers and indicators. is based on essentially the same chemical process, the conversion of one set of reactants to another, this reaction resisting or responding to external changes depending on the relative amounts of reactants involved. For geothermal systems. the most important fluid component besides water is usually carbon dioxide. Together with calcium minerals forming a large part of a typical geothermal rock assemblage, the system calcium-ahtminium-silicate + carboq dioxide + caicite can obviously be expected to occupy a high position in the hierarchy of buffers governing the chemistry of solid and fluid phases within a geothermal system. Mineral equilibria involving carbon dioxide

According to Fig. 1, the temperature dependence of Pco, for the coexistence of calcium-aluminium-silicates with calcite and quartz may be evaluated by use of the reactions 3 grossular + 5COz + w

lytical aluminium conantrations may be considered to represent largely the sum of concentrations of aluminiumhydroxy complexes present in solution according to in,, = a,,l.tK,~a;l-;.

= 2 ciinozoisite + 5 calcite + 3q

(5)

and (3)

where m,,, is the analytically observed moiality of aluminium in solution. K. the equilibrium constants governing

clinozoisite + 3CO: + 2.5~ = 1.5 kaolinite + 2 calcite

(6)

as shown in Fig. 2 together with stability areas for margarite and anorthite. The latter is only stable at temperatures above Hx)“C and, therefore, does not K. = u,,,,,“,~ -.a;(. /aA,, . (4) appear in Fig. 1. Figure 2 also contains data points for Wairakei, and 7” the corresponding activity coefficients. Application of eqn (3) is likely to be affected by considerable errors Kawerau and Broadlands, three explored thermal arising from the need to evaluate equilibration pH’s and systems in New Zealand. Carbon dioxide partial activity coefficients for up to triply charged ions at elevated pressures were derived by use of data reported by temperatures and ionic strengths. Together with the uncerGiggenbach (1980) assuming CO, to be completely tainty in the validity of presently available data for the equilibrium constants K, and the possible formation of dissoived in a single liquid phase. The data points other aluminium complexes involving e.g. fluoride and sul- occupy positions within the zoisite stability field but fate, the problems associated with the correlation of ananot along any of its boundaries. These positions agree lytical aluminium contents and theoretical Al,-values with clinozoisite (epidote) forming the most important would suggest that, for the time being, AI” is best considered a purely theoretical parameter accessible to only secondary calcium aluminium silicate phase at eleindirect evaluation. vated temperatures but do not support the assumption of PcOr being controlled by a reaction involving In discussions of geothermal rock-water interactions frequently the existence of mineral buffers two coexisting calcium aluminium silicate phases. thought to control pH or redox conditions at depth Extending the range of possible reaction partners to are invoked (ELLIS 1%9; BROWNEand ELLIS 1970). primary phases, the general trend of the data points Based on the fmdings represented by Fig. 1, however, appears to follow closely an extension of the aninto the clinozoisite stabit appears necessary to consider two quite indepen- orthiteikaolinite-boundary dent systems possibly controlling AI” and therefore. ility field pointing to the possibility of Pco, being aluminium silicate phase relationships: within the effectively controlled by a reaction involving these variability induced by substitutional disorder and two mineral phases.

the formation AI(OHg-”

of

aluminium-hydroxy

compkxes

391

Geothermal mineral equilibria

4 bar) 0.1 I.0

10 loo

Fig. 2. Stability diagrams for calcium-aluminiumsilicate (a) and feldspar-mica (b) reactions in terms of Pco, and temperature together with data points for well discharges at Wairakei. Kawerau and Broadlands. -‘&~~I,. refers to the mol fraction of pistacite Ca2Fe3Si,012(0H) in epidote. In Fig. 2b microcline/muscovite and low albite/paragonite coexistence boundaries in the presence of epidote with X ,,ll,,eicrof 0, 0.1, 0.2 and 0.3 are shown.

In the construction of Fig. 2 the activities of solid phases were assumed to be unity corresponding to the presence of pure minerals of the stated composition. In natural systems, however. anorthite and clinozoisite are likely to occur only as components in minerals of the albite/anorthite-plagioclase and the pistacite/clinozoisite-epidote solid solution series respectively with considerably reduced activities. Valid comparison of observed and theoretical COz-partial pressures, therefore, should only be carried out on the basis of mineral phase activities ci, as e.g. for the anorthite/kaolinite reaction by use of the relationship PCO~

2 .f CO> =

akaolinitc Gad%norttti~c

K7

(7)

where Kf represents the equilibrium constant calculated for activities of unity of the phases involved. According to BROWNEand ELLIS (1970) and STEINER (1977) the primary plagioclase most commonly encountered in New Zealand systems is andesine with a composition taken to be close to alb,, an50. Accept-

ing some deviation from Raoult’s Law (KERRICK and DARKEN, 1975) the activity of the anorthite component in andesine is estimated to be about 0.35. On the basis of eqn (7) such unilateral reduction in the activity of anorthite would correspond to an increase in the partial pressure of CO2 required to convert plagioclase to kaolinite and calcite by a factor of three, or according to Fig. 2 to an expansion of the plagioclase stability field into lower temperatures. In view of the compensating effect of an additional likely reduction in the activity of the kaolinite component, and possibly calcite also required in eqn (7), however, and the better fit to the observed data points, the uncorrected anorthite,/kaolinite-line may still be assumed to adequately describe the conversion of plagioclase to secondary layer silicates including montmorillonites and illites according to the general, less specific reaction plagioclase + CO2 = clay + calcite.

(8)

W. F.

398

The semi-empirical relationship dependence according to

GIWENBACH

with a temperature

log Pco, = 15.26 - 7850/(c + 273.15)

@a)

thus obtained suggests P,, in geothermal systems of the New Zealand type to be a unique function of temperature with equilibrium alteration mineralogy and relative activities of reactive fluid phase components likely to also depend on only one parameter: temperature. This extremely simple picture is likely to apply only to relatively high temperature systems occupied by a single, slow moving liquid phase in contact with calcite and plagioclase. As soon as boiling sets in, the acceleration in fluid movement caused by the increase

in fluid volume combine with the likely decrease in reaction rates with decreasing temperature to impede attainment of rock-water equilibrium and the chemistry of the system will increasingly deviate from that dictated by the plagioclase-clay reaction. The plagioclase/clay-calcite conversion reaction probably represents the most important primary mineral buffer within a geothermal system, at lower temperatures however, other competing mineral buffer systems, e.g. those based on the three pairs of Na-, K- and Mg-aluminium silicates shown in Fig. 1 may increasingly gain in importance. By accepting epidote to represent the major secondary Ca-aluminium sili-

cate in equilibrium with calcite over the likely cornpositional range of geothermal fluids, the coexistence boundaries of these possible three secondary buffers may be expressed in terms of Pco, by use of the reactions 3 micro&e

+ 2 ciinozoisite + 4CJOz + 2w = 3 muacovitc + 4 calcite + 6q

(9)

3 albite + 2 ciinozoisite + 4C02 + 2w = 3 paragonite + 4 calcite + 6q

(10)

15 phlogopite + 16 clinozoisite + 32CO2 + 28w = 9 clinochlore + 15 muscovite + 32 calcite + 2lq (11) As pointed out above, clinozoisite is likely to occur in natural systems largely as a component in minerals of the epidote solid solution series. An estimate of its activity in iron-substituted cpidote may be obtained by assuming iron lo predominantly occupy the M3 sites (KERRI~~, 1977) according to %li”.l~iSi,.= XALM3= l-3 XPI*UEil.

(12)

where XAISa represent the fraction of M3 sites occupied by the fraction of pistacite component aluminium and XP,SUOi,, present in the naturally occurring epidote. The e!Tect of such iron substitution on Pco, calculated for the clinozoisite/kaolinite-reaction (6) corresponds to PCO,

=

Wo&LSilc

(13)

as shown in Fig. 2a for pistacite fractions of 0.1, 0.2 and 0.3. Effects of the same order would result for reactions (9). (10) and (11). Because of the probably much higher and at preset unestimable uncertainties arising from the reduction in phlogopite and ciinochlore activities due to variations in iron contents of natural biotites and chlorites.

however, the chlorite/biotite-buffer [eqn (ll)] is omitted from further discussion. In the case of possible, but largely accidental compensation of activity corrections for this reaction, a behaviour closely corresponding to that of the paragonite/albite system can be expected. For the remaining two butier systems, involving alkalifeldapars and micas, the dependence of PC,, on mineral activities is given by where &to, again represents the equilibrium constants calculated for reactions (9) and (10) for activities of unity of the solid phases involved. In the above expression the activity of clinozoisite may be estimated by use of eqn (12) while it is probably quite safe to assume the activities of calcite and quartz to be close to unity. For the remaining reaction participants, the micas and feldspars, estimation of activity corrections is further complicated e.g. by the occurrenoe of non-stoichiometric and substitutionally disordered phases respectively. As shown in Fig. 2, the reduction in clinozoisite activity for natural iron-containing epidote gives rise to a shift in the coexistence boundaries representing reactions (6), (9) and (IO) to higher CO,-partial pressures or lower temperatures, the opposite e&t is observed if the equilibrium disordered feldspar (microcline, low albite) is replaced by the more highly disordered polymorph (sanidine, high albife), therefore, to a large degree compensating the etkcts of reduad clinoxoisite activity in natural epidote. A further such compenaatiag effec?can be expected to arise from the reduction in mica activities due to compositional variations. Concluding from the above, the major eliect of deviations in composition and degree of order from those of thermodynamically defined pure phases is likely to lead to a general broadening of the mineral coexistence boundaries shown in Figs 2a and b into less well d&ed coexistena ranges. Assuming these effects to be more or leas random for reactants and produots, the uncorrected boundaries baaed on reactions among pure phases, may still be considered to repmsent some kind of statistical mean corresponding to high coexistence probabilities with the evaluation of actual boundaries requiring detailed information on the compoitiou and the structure of the phases involved. Even if extenrive and auzurate intknation on the composition and degree of ordering of mineral phases in geothermal systems were available, their variability from thin section to thin section and within a thin section would again require the adoption of a more probabilistic attitude. In the present investigation the pure phaacs used to evaluate thermodynamic relationships have to be seen largely as representatives or symbols for a certain thermodynamic environmen t. Paragonite has never been reported for any of the New Zealand geothermal systems. Its use in the above calculations is due to the lack of thermodynamic information on any other Na-aluminium-silicate required in complementing albite in reactions delineating the stability of the feldspar at high Al”-values. The albite alteration product actually encountered is the montmoriilonite. Ona reliable thermodynamic information for this mineral becomes available, it may be substituted for paragonite in the above reaction without probably affecting the overall

conclusions reached here in any major way. Accepting the conclusions reached in the above discussions, the distribution of data points for Wairakei, Kawerau and Broadlands apparently follow trends

delineated by the feldspar/mica coexistence lines, with data points for high temperature systems (Kawerau, Broadlands) close to the low albite/paragonite coexistence line. Those for the lower temperature system of Wairakei, however. show better agreement with the line representing microcline/muacovite coexistence.

399

Geothermal mineral equilibria

On the basis of the limited data available, the distribution of points in Fig. 2 may tentatively be interpreted in terms of Pc,,-control in geothermal systems predominantly by the plagioclase/clay conversion reaction. Depending on the temperature or enthalpy of the fluid involved in this reaction, a system may then ‘lock on’ to a secondary buffer system invofving at lower temperatures the potassium phases microcline/muscovite, at higher temperatures low albite and its alteration products or possibly biotite/chlorite.

with caicite, its reaction with H$ in terms of commonly encountered geothermal mineral phases corresponds to 2 clinozoisite + 3 pyrite + 2 calcite + 3q + 15~ + 1.5Hz = 3 epidote + 2COs + 6H2S

(15)

which in addition to COr and H$ incfudes H2 among the fluid phase components. In combination with a recently derived empirical relationship (GIG GENBACH, 1980) describing the interaction of HIS with two-valent iron present in an as yet unidentified iron~uminium-siIi~t~ (FeO), according to

In the above evaluation of reactions involving CO2 simultaneous control of Pm2 by reactions involving a primary (~o~hite) and a secondary (clino~i~te} Ca-aluminiumsilicate was postulated. This violation of the phase rule points to the inherent dilemma in any discussion of geo(16) (FeO) + 2H2S = Fe!& + Hz + (H*O) thermal rock-water interactions on the basis of ideal&d thermodynamic model. Geochemically a geothermal syswhere (HrO) represents the iron-free, protonated form tem provides one of the means for the conversion of a of the aluminium silicate, eqn (15) is converted to thermodynamically unstable to a stable mineral asscmblage. In the caseof a closed, finite system the major role clinozoisite + 0.75 pyrite + calcite of the ftuid phase would be that of a catalyst and to a +0.75(FeO) i- 1.59 + 2.5~ minor degree that of a reaction participant involved in (17) hydration and redox reactions. On attainment of full equi= 1.5 epidote + O.B(HsO) + COz + 1.5H2S librium within this system the composition of the fluid phases would be uniquely determined by the composition By use of the empirically derived temperature of the resulting alteration assemblage. dependence for reaction (16) In the case of an active Ethel system, however, continuing addition to and withdrawal of fluid mixtures log(Pn,/P;,,) = 0.16 + 504g/(t + 273.15) (18) from an e&ctively infinite rock system leads to the establishment of a chemical steady state retlecting the attainthat of eqn (17) is found to be ment to varying degrees of often only partial or local equilibrium with respect to the two main sub-reactions making up the above mineral requilibrium process: dissolution of unstable primary minerals and lion of stable, ~23.67 - 13 t75/(t + 273.15) (19f seco&uy phases. Digerent components of the geothermal fluid phases will Meet dlgerent degrees of attainment of The above relationships are valid for the interaction partial equilibrium with respect to different reactions. of pure phases of the stated compositions except for Major components. such as CO2 are likely to dominate the (FeO) and (H,O) whir& because of their empirical system through interaction with major phases of the derivation, already represent the activities of an primary rock matrix with minor components mereiy iron(H) aluminium silicate (prob~iy the daphuite or adjusting to the conditions estabfisbed by the initial, dominant reaction. annite components in natural chlorites or biotites)

In the above discussion two important components of a geothermal mineral-fluid system, iron and hydrogen sulfide, were not included. Under geothermal conditions they are likely to be linked through reactions involving the common geothermal minerals pyrite, epidote and probably chlorite and biotite to the aluminium-silicate reactions already discussed. Geothermal minerals

eqtiifibrio

inwlving

iron

containing

In two recent studies (BARTON et al., 1977; GIGGENBACH,1980) iron-containing chlorite was suggested to be one of the most likely phases complementing pyrite in reactions possibiy controlling oxidation potential and hydro~~~~e partial pressures in geothermal systems. Another important iron containing mineral phase commonly encountered in geothermal systems is epidote. In the absence of further information an idealised composition, Ca2FeAlzSiJ0r20H, and thermodynamic data given by HELGESONet al. (1978) are used. Again assuming epidote to represent the major secondary calcium aiuminium silicate in equilibrium

and its iron-free, protonated counterpart. The activities of the usually quite pure phases pyrite, calcite and quartz can safely be assumed to be close to unity. The activities of clinoxoisite and epidote in naturally occurring solid solutions may again be taken into account by use of eqn (12) for clinoxoisite, and a corresponding reiationship for epidote %ldol.

=

3&&r

(20)

Observed Pco,P$$val~s may, therefore, be assumed to be related to the equilibrium constant Kc derived above through p co2pt.5 Hz5 E K’G’

hwodkJd&ae

=Ko(l-3

Xpin*ot,c)/(3&Wte)1~5

(21)

In Fig. 3 variations in PHg with Pwt and pistacite content of epidote are given for reaction (17) together with the range of COr- and H,Spartial pressures for New Zealand well discharges producing from systems close to 300°C. Theoretical and observed gas partial pressures agree best for pistacite mole fractions of between 0.20 and 0.25. Unfortunately no sufficiently detailed analyses of naturally occurring epidotes from

W. F. GIGGENBACSI

400

buffers. The trends delineated by the annitejepidotesystem are close to ind~tin~iahab~e from those evaluated by use of the less specific (FeO). Assuming a pistacite mole fraction of 0.22 to be typical for geothermal epidotes generaify, theoretical and observed partial pressure products Pco,P& may be compared by use of a procedure previously developed (GIGGENWCK 1980). GW pW&I p~M!MRS Pi are related to total discharge gas contents x&i through Pi = X&,BfPH*/Df’

-2

-1

0

1

2

3 IogP

a2

Fig. 3. Variations in Pap vs Pco, as a function of mineral phase equilibria involving &dote, and (FeO) (see text) in the presence of pyrite, calcite and quartz. For the epidote reaction the dependence on iron content (XPiulo,,J is given. For comparison coexistence boundaries for reactions involving annite instead of (FeO) and anorthite instead of clinozoisite are shown.

1221

where B( represents the liquid/vapor-distribution coefficient of the gas i, P,, the saturated water vapor pressure, and Dr a parameter 4

=(I

-y+yBJ

(23)

allowing the composition of the fluid sampled to be expressed in terms of B1 and y the fraction of equilibrium vapor present in (I)+‘) or lost (B-t) from the discharge. In Fig. 4 values of log (x~~~~,x~&S)as a function of temperature

and

.v are shown

together

with

data

points for Wairakei, Kawerau, BroadIands, Lardenllo and a number of other geothermal areas, The distribution of analytical data points closely agrees with that predicted by theory, except for the two points representing steam discharged from the low temperature Indian geothermal system at Manikaran. This deviation suggests that the overall equilibrium model is likeiy to break down at lower temperatures due to nonattainment of equilibrium or to Pco, and PHs being connected through a difkient set of mineral phases. THE LHBTBUWTION OF I~%INBBAL PHASES WFlwtN GEO-MAL SYSTEMS In the above discussion comparison of theoretical and observed COz- and H&contents and partial Fig. 4. Theoretical and observed values of log (~,.~o,.x:;&) pressures of geothermal fluids was restricted to disas a function of temperature and the fraction ,v of steam charges from sufficiently deep wells considered to propresent in or lost from the equilibrium liquid phase discharged from a weft. Data points for Larderelto were duce liquid-vapor mixtures closely representative of obtained from data reported by TRUIBDELLand bb3UUNG the deep rock/water-equilibration fluid Within this (1978). those for Wairakei, Kawerau. Broadlands and framework the chemistry of the fluids sampled and Manikaran from GIGGENBACH (1980). the rest from ELLIS the distribution of mineral phases at the production and MAHON (t977). (Xpi,IaE,,.= 0.22). Open c&ts replevel were found to be best described in terms of the resent vapor dominated closed circles liquid dominated systems. attainment of some intermediate steady state retlecting to varying degrees the effects of initial fluid interaction with the primary rock matrix and those of close approach to equ~ibrium with the thermodynaNew Zealand geothermal systems are available to mically stable secondary mineral assemblage. The allow the above value to be confirmed. For one epidote crystal from Broadlands an iron content of 13O$ simple picture obtained, however, is likely to apply only to relatively high temperature systems occupied is reported (BROWNISand Eurs, 1970), relatively ironrich, as required by the above evaluation. Figure 3 by a slow moving liquid phase. As soon as boiling sets also contains stability boundaries for two other as- in the acceleration in fluid movement caused by the increase in fluid volume is likely to combine with the semblages possibly controlIing refative COz- and decrease in reaction rates with decreasing temperature H&partial pressures. Both reactions reproduce obto progressively impede further attainment of minserved partial pressures very closely at all temperatures, and therefore, represent, after application of eral/tluid equilibrium. Another major factor likety to affect the distribution of mineral phases within a geosuitable activity corrections. possible alternative loo”

MO”

Geothermal mineral equilibria

401

Table 1. Carbon

dioxide partial pressures (b) of an initially one-phase. 1 molq; solution of CO2 during the adiabatic expansion from temperature tl to t2

HI (Jig) t, (‘C) 12

w

75 loo 125 150 175 200 225 250 275 300 325 350 l1-12

25 20 15 10 5 0

419 100

632 150

852 200

1085 250

1344 300

1671 350

0.09 53.76

0.03 0.12 0.52 69.75

0.02 0.06 0.17 0.51 1.83 68.29

0.01 0.04 0.10 0.25 0.60 1.50 4.51 54.71

0.01 0.03 O.Of 0.16 0.34 0.70 1.47 3.16 7.81 37.80

0.01 0.02 0.05 0.11 0.22 0.43 0.81 1.48 2.68 4.89 9.48 23.33

0.09 0.14 0.23 0.41 0.95 53.76

0.52 0.14 1.13 1.93 4.26 69.15

1.83 2.54 3.71 5.98 11.99 68.29

4.51 5.97 8.25 12.27 21.07 54.71

7.81 9.75 12.49 16.64 23.63 37.80

9.48 11.01 12.92 15.63 18.65 23.33

tI = temperature

of first boiling tt = fluid temperature

&o (b) 0.39 1.01 2.32 4.76 8.92 15.54 25.48 39.13 59.42 85.81 120.40 165.13

after adiabatic cxpan-

sion, H, = enthalpy of fluid.

thermal system is the change in fluid compositions due to exsolution of gases during adiabatic expansion of an initiaIly single liquid phase fluid rising from depth. Variations

in jiuid composition with depth At sufficiently high pressures and depths all the gases

present in a rising thermal fluid can be assumed to JX completely dissolved in a single liquid phase and the activity of the gas i at total pressures < 100 bar corresponds closely to its partial pressure Pi as given by

Pi =

&I.i.tXl.i

where &.I., is the Henry’s law constant and x1.i the mole fraction of the gas i in the liquid phase at temperature 2. With decreasing total pressure a vapor phase is likely to form and the gas i will distribute itself according to XI%&,+ &Al - &I = x0.i

(251

X,.Jxl.i = Bi,f,

(26)

and where h,i and Xl,i are the mol-fractions of gas i in the vapor and liquid phase respectively. xosi that in the original deep liquid phase, ~12~ the fraction of steam formed adiabaticaIly, and .E)Ithe vapor-liquid gas distribution co&icients as definad by eqn (26). Values for Bi suitable at low ionic strengths (~0.06 molal) are reported by GIMENBACH (1980). By combining eqns (25) and (26) the mol-fraction xX.$of gas in the liquid phase is given by X1.i = xO.J(l - Jiv,,+ B.,,Y,,)

.

(271

By msertmg these values for x1 .i into eqn (24) a relationship is obtained allowing P, to be evaluated as a function of y,,, the fraction of steam formed after adiabatic expansion of a fluid with a given enthalpy H, to a temperature t2 according to

where HI,,, and &.,, an the enthaMts ofliquid water and steam at the temperature tz respectively. ln the case of carbon dioxide. its partial pressure then is given by pco,,, = ~n.co,.t~xO,co2l~~ - a, + ko,,*,Yz,)

(29)

In Table 1 values of Pcola, are given for tota1 fluid C02-contents of one mol%. C02-partial pressures for other values of X&,~ may be obtained by use of By neglecting the effects of salinity (Haas, 1971) and those of the other gases present and assuming total pressure in rising geothermal fluids to correspond to Pf f %o

+ pco,

(31)

the values Pm, and PHp given in Tabk 1 may be combined to obtain boiling-point versus depth curves for geothermal fluids as a function of X&content and enthaipy as represented by the temperature of &rst boiling, tl. A mathematically more demanding evahzation of the pressure-temperature curves for the two phase CO,-water systern has been carried out by SW (1976) and Surro~ and MCNABB(1977). Their curves, however, closely agree with those derived by use of the above more straightforward procedure.

In Fig. 5 the variations in Pco, for liquid-vapor mixtures forming from a liquid phase with enthalpies corresponding to liquid water temperatures for 150 to 3OO’C and a C02-content of l.Omol% are shown. Pco, vs t-curves for fluids with higher or lower initial liquid phase CO+ontents are obtained by shifting the starting point for a given temperature verticaliy to the curve representing the desired liquid phase COt-content. The C02-isocompositional curves shown in Fig. 5 were drawn by ose of drita presented by TAKENOUCHIand KENNEDY(1%5). They suggest that CO,-partial pressures caiculated for liquid phase

402

Fig 5. Carbon dioxide partial pressure vs temp. curves for fluids containing 1 mol”/; CO2 with temperatures of first boiling of 150. 200. 250, 300 and 350°C together with stability boundaries for geothermal mincrai pairs as calcuIatcd by use of eqns (S),(4),(9) and (10).(H, represents enthaipy of fluid mixture in J/8).

CO,-contents up to 1 mole/, correspond closely to those expected for Henry’s law behaviour, at higher liquid phase CO1-contents the solubility of CO2 dccreases considerably, especially at lower temperatures, and equiliirium CO,-partial pressures required to maintain a given iiquid phase C&content are substantially higher compared to those resulting from ap plication of Henry’s law. Liquid phase CO+ontents within the accessible parts of a geothermal system however are unlikely to ever exceed 2 molx and the Pco, vs temp. relationship described by eqn (29) can be expected to be valid for most systems investigated. Figure 5 also contains boundaries for some of the most important mineral-bu&r-indicator pairs. With decreasing temperature the mineral assemblages likely to be stable in contact with the rising liquid changes from one consisting essentially of fcldspars to assembiages increasingly made up of micas. ill&es and montmorillonites until at some stage the kaolinite boundary is reached. In Fig. 6a P,o, vs temp. curves representing fiuids corresponding to those encountered at Wairakei and

Broadlands are shown. For some 40°C after boiling sets in the fluid system follows closely the plagioclase/ clay conversion line, with further decreasing temperature the fluid systems leave the feldspar stability areas. In the case of Wairakei low albite can be expected to be stable down to around 140°C. at Broadlands it should theoretically be limited to temperatures above 220°C. This marked difference in theoretical mineral stability ranges between the two areas corresponds closely to that observed in drill-cores (BROWNEand ELLIS,1970). At Wairakei secondary albite is encountered at depths as shallow as 9Om at BroadIands hydrothermal albite is only observed at depths greater than 450m corresponding to cold hydrostatic pressures of 10 and 45 bar and liquid/vapor-coexistence temperatures of 180 and 250°C respectively. Analyses of geothermal steam samples are frequently limited to the determination of CO2 and HzS with the COljH2S mol ratio representing one of the main parameters available to characterise a geothermal gas discharge. In Fig. 6b theoretical values of RG = x,,Co~;.q,H3 are shown as a function of tem-

403

Geothermal mineral equilibria

GROSSULAR

KAOLINITE

I

GRGSSULAR

KAWNITE

2. .

0 Woirokei 0 Broadlands

3. d

ZcF

Fig. 6. Stability diagrams of Ca-aluminium-silicates. For the stability field of epidote microcline/muscovite and albite/paragonite coexistence lines (a) and COJH&mol-ratio contours (b) are shown together with P,,, vs depth curves for fluids with temperatures (260 and 300°C) and C02-contents (1.0 and 0.06 mol%) typical for Wairakei and Broadlands well discharges.

perature and Pm, as calculated by use of the relationship Rc = P:&B,s/K’E’3&o,

(32)

and values of K& represented by eqn (21). Equation (32) is valid for liquid dominated systems neglecting the effects of the presence or loss of equilibrium vapor in the discharge sampled. For vapor-dominated systems gas-partial pressures may be assumed to correspond to Pi = PHpxi and the relationship between RG and Pm, is given by RG = PgJK;;13

(33)

For the two examples chosen in Fig. 6b the agreement between theoretical COs/H$ ratios and those actually observed of around 5-10 for Wairakei and of 60-100 at Broadlands is excellent and again supports the assumption that relative COr- and H&contents in geothermal systems of the New Zealand type are controlled by a mineral assemblage closely resembling that represented by reaction (17). Assuming CO,-partial pressures during the rise of the thermal buid for the two systems to follow Pco, vs

temp. curves as given by eqn (2!J), CO1/HIS-ratios can be expected to decrease for the frrst 50°C after boiling sets in, from then onwards relative H+contents are likely to decrease dramatically and in the case of Broadlands HsS should have been effectively removed when the fluid emerges at the surface. The CO~/H&ratio in steam collected from e.g. Ohaki Pool, the major natural discharge of the Broadlands geothermal area, of around 60, however, suggests that the rate of rock/water interaction as compared to that of fluid movement within the two phase region of a geothermal system is much too slow to allow the fluids to re-equilibrate in accordance with the theoretically expected trends. By presupposing complete lack of re-equilibration of a geothermal fluid within the two phase region of a liquiddominated geothermal system the discharge composition can be assumed to correspond to that of a fluid rising from depth along the single liquid phase, plagioclase/clay-conversion line with CO1/H+ratios reflecting conditions at the point of first boiling. By expressing Pco, in eqn (32) in terms of eqn (8) the temperature dependence of Rc = ~t.~~Jxt,~~ in a

W. F. GEGENBACH

404

single liquid phase system at the onset of boiling is

given by log RG = 14.10 - 663O/(r + 273.15)

(34)

with the temperature c in C. In this evaluation PcO, is assumed to be controlled by the plagioclase/clayconversion reaction while relative C02/H&contents correspond to those for reaction (17). By use of analytical data reported by GIGGENBACH(1980). COz/ H&equilibration temperatures of 220 to 250°C are obtained for Wairakei, 250 to 260°C for Kawerau, and 265 (Ohaki Pool) to 280°C for Broadlands, close to those measured or derived by use of other geothermometers. Because of the generally observed absence of epidote from geothermal assemblages at lower temperatures (STEINER,1977; BROWNEand ELLIS, 1970), evaluation of C02/HzS-equilibration temperatures by use of eqn (34) should be limited to well discharges from systems for which deep temperatures of at least 200°C have been determined by other means. Vapor-liquid separation process also have a dramatic effect on the distribution of COz and HIS between liquid and vapor phase and consequently on the ratio COt/H2S (GLOVER. 1970: GIGGENBACH, 1980), care should therefore be taken to select data as representative of the undisturbed equilibration liquid phase as possible. The satisfactory agreement between theoretical and observed gas partial pressures suggest close attainment of equilibration among vapor and liquid phase constituents and a secondary mineral assemblage including epidote. pyrite, calcite, quartz and probably chlorite or biotite. This reaction, likely to control relative CO2 and H&contents. represents another obviously very powerful buffer consisting largely in the interaction of CC+ and H2S with the ahtminiumsilicate phases to form calcite and pyrite and, therefore. may link deposition of sulfide ore minerals directly to variations in CO,-partial pressures. The form&on

of metasrable geothermal mineral phases

In the above discussion, mineral-fluid interactions were evaluated in terms of primary mineral phases and secondary phases stable according to thermodynamic data available. Within this framework, frequently observed geothermal minerals such as the calcium-aluminium-silicates wairakite. prehnite or laumontite had to be omitted as they were found to be theoretically unstable. Discussion in terms of thermodynamically stable phases also preclude the evaluation of the effects of the interaction of thermal fluids with an important component of a primary rock matrix. the glassy ground-mass commonly encountered in geothermal systems with close magmatic-volcanic associations. The thermodynamic approach as described above provides a means of distinguishing thermodynamically stable from unstable phases without giving an indication as to the degree of instability of a given phase. Obviously some phases will be less unstable

than others, these differences in instability largely determining the likelihood of a phase to persist or to form metastably during the transition of a geothermal system from an initially thermodynamically unstable state to the final. thermodynamically stable state (HELGESON,1968). The formation of intermediate metastable phases is a simple consequence of the Ostwald step rule stating that the transition of a chemical system from an unstable to a stable state may occur via a series of intermediate metastable states, the thermodynamic instability of these states decreasing with increasing reaction progress. The parameter reflecting the relative stability of a chemical system is its free energy whose values can rigorously be determined only for a closed system. Another approach to obtain a measure of the relative stability of geothermal mineral phases consists in the determination of the relative solubilities of a series of mineral phases, the least soluble phase, of course. being the most stable. The solubility of an aluminium silicate may be expressed in terms of eqn (1). By rewriting it in the form Yn = (Li.q.v - mAlu)/n

(35)

a relationship is obtained allowing the solubihty of a mineral to be expressed in terms of Yn = log (ai/‘&’ ) and theoretical values of Li.4.\y representing the logarithm of the hydrolysis constant of the mineral Y.Al,-silicate. The remaining parameter Aln required to evaluate Yu is obtained by accepting Al,-values in a liquid dominated geothermal system to be essentially determined by the interaction of carbon dioxide with plagioclase according to the reaction anorthite + CO2 -I- 6H’ = calcite + 2A13’ + 2q + 3w (36) and PcO, to correspond to that evaluated for the adiabatic expansion of a fluid with a given COz-content and enthalpy during its rise to the surface [eqn (29)]. The resulting values of Yu, representing relative solubilities at a common pH, are shown for calcium phases in Fig. 7. The least soluble phase at temperatures above 240°C is epidote. to be replaced at lower temperatures by Ca-montmorillonite. This finding closely agrees with observations in geothermal systems such as Broadlands (BROWNEand ELLIS. 1970) where formation of epidote is restricted to temperatures above 25O’C. the prevalent Ca-aluminium-silicate at lower temperatures being the montmorillonite. The dashed lines in Fig. 7 represent the shift in epidote solubility with variations in XPtltsCLtC.With increasing iron content the stability of epidote can be expected to extend to lower temperatures. At temperatures above 250°C the solubilities of the major geothermal Ca-aluminium silicates as shown in Fig. 7 are very similar and only slightly above that of epidote. Anorthite is more soluble than another common geothermal Ca-phase. wairakite. but less soluble than prehnite or laumontite. Assuming mineral con-

Geothermal mineral equilibria

405

Fig. 7. Solubility of Ca-aluminium-silicates in terms of CaH = log (atiz+ /ai- I for a fluid system formed from a liquid phase containing 1mol9. CO2 at 300°C. For epidote and plagioclase solid solutions curves representing mole fractions of pistacite of 0, 0.1. 0.2 and 0.3 and of albite of 0 and 0.5 fandesine) are @en. Reactions involving volcanic glass are discussed in the text.

version reactions to obey the Ostwald step rule, at 280°C the dissolution of primary anorthite can be expected to lead initially to the formation of wairakite and then epidote as the least solubk phase. At lower temperatures dissolution of plagioclase is likely to result also in the formation of laumontite as an intermediate, metastable phase followed possibly by margarite and finally Ca-montmorillonite. Similar reaction sequences are likely to accompany the alteration of the volcanic glass fraction of the original rock. Because of its low thermodynamic stability and correspondingly high solubility, the alteration products of volcanic glass may initially include a considerable proportion of Ca-montmorillonite frequently observed (BROWNEand ELLK, 1972) at temperatures above that of the thermodynamic montmorillonite/ epidote coexistence range, with prehnite, laumontite and wairakite representing other possible me&table alteration phases. The actual sequence of phases encountered, however. can be expected to depend strongly on kinetic factors such as relative rates of nucleation, crystallisation and recrystallisation and. therefore. also on the residence time of an assemblage within a system. This time dependence suggests that the type of assemblage encountered may provide an indication of the ‘maturity’ of a system, with assemblages containing high proportions of metastable phases likely to represent a system ‘younger’ than one made up largely of a thermodynamically stable, well equilibrated assemblage. In general. the solubilities represented in Fig. 7 depend to a significant degree on the choice of mineral system assumed to effectively control Al,-values.

The assumption of anorthite-COr-calcite control is likely to be only valid for a temperature range of between 200 and 350°C. By selecting appropriate A&,-buffer systems. however. it should be possible to extend application of the above technique to lower temperature diagenetic and weathering. or higher temperature metamorphic systems. General discussion ofgeothermal

rock-fluid

interactions

As stated above, the main process affecting the distribution of minerals and solutes in a geothermal system consists in the gradual conversion of mineral phases thermodynamically unstable under geothermal conditions to a stable assemblage according to unstable, primary minerals + stable. secondary minerals

(37)

representing an overall irreversible reaction. By assuming the transition from an initially unstable to the final, thermodynamically stable state to represent a succession of close to reversible reactions reaching partial or local equilibrium with respect to the previous or next reaction. HELGEWN(1968. 1971. 1979) was able to evaluate the kinetics of mass transfer reactions during hydrothermal alteration processes as a function of reaction progress towards full attainment of equilibrium. In the case of a closed system the distribution of components is given by the mass balance fluid,,itid + primary phases = fluid,,,, + secondary phases

(38)

406

W.

F.

GIGGENBACH

At any intermediate stage during this conversion process the composition of the fluid is a function of reaction progress and likely to re!Iect both the effects of earlier formation conditions and increasingly those of approach to equilibrium. In the case of an active geothermal system open to the passage of mobile phases the approach to equilibrium is greatly compiicated by the existence of large temperature and pressure gradients and is generally interrupted when the fluid reaches the surface. Even for essentially undisturbed, well established systems, the fluids accasible to sampling then are likely to have reached some compiex steady state comgrosition r&acting the combined effects of initial fluid composition, the kinetics of primary mineral dissolution and secondary mineral deposition at decreasing temperatures and pressures, in addition to vapor loss, dilution and mixing with fluids of different origin.

In the evaluation of the kinetics of mass transfer among silicate phases and aqueous solutions, H-N (f971) assumed dissolution rates to be controlled by the formation of a surface layer of intermediite reaction products thus closely linking dissolution and redeposition processes. Reorntiy, however, LAoAcXi?(1976) showed that the dissolution of feidspars corresponds to a rektivdy simpk surface controlled process with the rate of dissoiution proportional to the degree of saturation reached by the solution and the precipitation of alteration products taking place in a series of essentially unrelated reactions. These findings were later on substantiated by Rrnovrc (1976) and HotDEN and BERNER (1979) who also showed that the surface presented by the dissolving feldspar to the solution remains chemically largely unaltered. The two hydrothermal alteration subreactions, dissolution and redeposition, can therefore be assumed to take place &se to indapen&ntiy with the geothermal solution initially striving towards dimolution equilibrium with the primary mineral assemblage. The mode of attainment of this initial dilution equilibrium can be expected to be a function of the relative amoums and ‘reactivities’ of fluid components and primary rock constituents involved. Tba most reactive components of a geothermal fluid are (20% and HrS. With CO2 by far predominating, it is safe to assume that HrS plays only a subordinate role in this battle for geochemieal control. In order for the rock dissolution reaction to be able to progress an etIicient means of removing reaction products accumulating in the reaction fluid has to be provided. In the case of C02-attack an excess of the reaction products Ca” and HCO; is removed through precipitation of calcite, excess silica in the form of quart& and the rest consisting largely of Na’, K’, M8”, Fe” and A13’ in the form of secondary aluminiumdlicates. Kinetic barriers, however, are likely to deiay nucleation and ~s~~~ation and allow the Suid system to persist to varying degrees in a state of supersaturation with respect to the formation of a stable secondary assemblage, while still reflecting to varying degrees the approach to equilibrium with respect to primary phases. Tire alteration assemblage can, therefore, be considered to form largely in response to conditions established by the primary rock/fluid interaction process and, therefore. represents ‘truly’ secondary phases. Accepting CO+zontents by plagiociase/ciay-oalte

to be essentially controlled conversion [eqn (St], cor-

responding Pco,-values can be used to derive temperature/depth-curves by assuming the geothermal fluid to be close to va~r~iquid~~~sten~. Total Ruid pressures may then be obtained by use of eqn (311. A

correlation with depth is established if the thermal guid pressure is taken to be balanced by the hydrostatic pressure of a column of cold water as shown in Fig. ga together with maximum ~tto~ole temperatures for Broadlands wells as reported by MAHONand FINLAYSON(1972). Most of the data points occupy positions over a region delineated by the anorthitet kaolinite and ep~dote/kaolinite coexistence lines. Points outside this range represent wells subject to mixing with cooler waters at the margin of the geothermal system as outlined by increasing resistivity contours. The position of the data points below the anorthite/kaolinite-line is likely to be due to several factors. One already mentioned is dilution or mixing with cooler waters lowering measured temperatures. Well discharges can also be expected to represent fluids having reached only partial equilibrium with respect to piagiociase interaction. For fluids with limited access to the primary rock matrix, COxpartial pressures are likely to remain well above those corresponding to full equilibrium The position of data points in Fig 8a therefore, would suggest most of the fluids to have reached only partial quilibrium with the thmmticd ~~~/~ay~~e representing the upper limit for a fully equilibrated system. This curve represents actually another ‘boiling-point vs depth-atrve based on the assumption of liquid/ vapor coexistence as given by eqn (311 and therefore again only an upper limit. The position of data points then could simply reflect equilibration at depths and pressures exceeding those required for the gases to remain completely dissolved in a singly liquid phase. Figure 8a also contains maximum temperature/ depth-curves for microcline/muscovite- and low albite/paragonits+coexistence. Their positions in relation to the ~o~~t~~~~~~~urv~ curma~ond to those shown in Fig 6a. In terms of thermodynamically stable phases, secondary albite should show a high probability to form only at temperatures below 280°C. with the probable formation of microcline theoretically restricted to a small temperature interval at around 250°C. At Wairakei maximum discharge temperatures of 240 to 260°C are generally observed at a depth of around 500 m with close to constant temperatures beneath this level. Such behaviour suggests equiiibration well within a single liquid phase until the boiling point vs depth-curve for CO#,O-mixture represen-

tative of Wairakei discharges is reached. Similar patterns are observed for wells drilled in geothermal areas of Iceland (ARNORSSON et al., 1978) and at El Tatio (CUSICANQLJ~ er ut.. 1975), with maximum down hole temperatures obtained at depths well below those expected for the vapor-pressure curve of pure water or saline solutions (HAAS, 1971). Positions above the pia~oci~~aolinite~nve~ion line might be observed in geothermal systems where one of the primary components of the piagioclase/ciay-caicitereaction, carbon dioxide or anorthite, is present in amounts too small to allow the reaction to act as an

Geothermal minerai equilibria

407

Fig, 8. Boiling point vs depth_(a) and fluid pressure P&b) curves for a s&s of mineral -b&s. Wairalcei, Kawcrsu, Broadlands, El Yatio and Iceland (ARHCJRSSON et al, 1978)wells correspond to the point at which measured temperature/depth and temperature/pressure curves change from close to constant temperature with depth behaviour to that indicating boiling point vs depth conditions.

The data points for

effective buffer. Such a situation may arise in the case of non-magmatic systems deriving their heat largely through deep circulation without an opportunity to acquire enough CO*, or in the case of systems associated with rocks high in alkali feldspars and low in the anorthite component. In this case Pm, may be expected to be controlled by albite to paragonite or microcline to muscovite conversion reactions. Restricting further considerations to systems for which Pco, is likely to be controlled by the piagiocl~/clay-buffer, the com~ition of fluids discharged at the surface can be expected to closely retlect conditions at the point of first boiling. With Pm, cmtrolled by the plagioclaselclay-buffer, this temperature is solely a function of the enthaipy of the rising fluid which in turn is iargely determined by the relative proportions of magmatic or metamorphic fluid and cooler meteoric water making up the discharge. In the case of highly permeable systems good hydraulic connections and resulting mixing with invading cooler groundwaters lower the enthalpy of the rising fluid column and the point of first boiling will be at shal-

low levels. With decreasing permeability and dilution the point of first boiling will shift deeper and deeper and iv extreme cases the deep liquid may be prevented from reaching the surface aftogether leading to the establishment of vapor-dominated zones over the accessible parts of a geothermal system (WHITE et al., 1971). Under these circumstances the hydraulic connection between the high temperature geothermal fluid column and that of the surrounding cold water is interrupted and pressures within the geothermal fluid system may well remain below cold hydr~~tic causing temperature/depth-data points in Fig. 8a to apparently deviate to values below those predicted by the theoretical curves. Because of this lack of pressure depth correspondence, especially for few permeability systems, comparison between theoretical and observed temperature/pressure relationships is better carried out on the basis of variations in downhole t~~ratur~ with actually measured downhole pressures rather than depths as shown in Fig. 8b. The position of data points fully agrees with the assumption of Pco,-control initially by the plagioclase/clay-buf

408

W. F. GIGGENBACH

Fig. 9. Schematic cross-section of a geothermal system of the Broadlands type. COs, H2S and the other constituents of the geothermal fluid are assumed to be largely derived through interaction of gases (HsO, COr, SO*, HCl, HF) released from a magma located at an arbitrary depth of around 8 km with deeply circulating groundwaters and rock. Stability regions for secondary minerals were drawn by use of temperature/pressure-relationships corresponding to those of Figs 6 and 8. A region of high formation probability for microcline, theoretically unstable under Broadlands conditions, is shown.

later on by the P, vs depth curve for CO,/watermixtures. In the case of liquid dominated systems, rock permeability can be assumed to be sufficiently high for pressures at a given depth to be close to the hydrostatic pressure of a corresponding column of cold water and the information provided by Fig. 8 may be translated into a schematic cross-section of a geothermal system as shown in Fig. 9. Equilibrium temperature/depth coexistence contours for three secondary mineral pairs, represented by the pure phases microcline/muscovite, low albite/paragonite and micakaolinite are delineated for a fluid system with an enthaipy of 1344 J/g (Broadlands). Isotherms were drawn by assuming a geothermal gradient of SO”C/km outside the thermal system and temperature/depth relationships within the system corresponding to those given in Figs 6 and 8. The system is assumed to derive its heat from a magmatic intrusion with the top of the magma column at an arbitrarily selected depth of some 8 km beneath the surface. From the geochemical point of view the system may be subdivided into two sub-systems: a lower magmatic-volcanic system and the upper geothermal system proper. The dividing line is considered to be represented by the plagioclase/cIay conversion curve (Fig. 2a) separation COz-consuming

diagenetic to geothermal rock alteration systems from CO*-producing metamorphic and magmatic systems. The latter is taken here to be the source system for most of the thermal tluid components. The generation of this fluid can be expected to be a close to equiiibrium process with the composition of the vapor phase released assumed to correspond closely to that of volcanic surface gas discharges with H20, CO;?, SOr, HCI and HF forming the major components (GIG GSNBACH and LEGUERN, 1976). Absorption of these gases by a deeply circulating originally meteoric water phase should lead to the formation of highly reactive solutions and intense rock/water interaction until the most aggressive components, SO*, HCI and HF are converted to less reactive products such as Fe&, HIS, NaCl and CaFz respectively (GIGGENBACH, 1977). Only then can the system be expected to approach behaviour governed by the plagioclase/clay buffer. Assuming full attainment of thermodynamic equilibrium between mineral and fluid phases and uniform and isotropic permeability conditions possible formation of secondary K-feldspars is limited to a narrow zone along the 250°C isotherm, at around 500 m depth. Secondary Na-feldspars should be stable over a somewhat larger portion of the system. Outside the albite stability volume, the alteration assemblage should consist exclusively of sheet silicates such

Geothermal mineral equilibria as micas. illites and montmorillonites. Local variations in permeability and rock/fluid contact however, are likely to lead to large deviations from this idealised distribution pattern. For instance. in zones of limited permeability CO1 may rapidly be used up by initial interaction with the anorthite component of primary plagioclase allowing secondary albite to form in a close to isochemical recrystallisation process well away from the thermodynamic stability region. Concluding from the thermodynamic distribution of phases within a geothermal system of the Broadlands type, the preferential formation of feldspars over a medium depth region in the center of the thermal upflow zone should be accompanied by a corresponding reduction in clay formation over a depth interval from around 300-1OOOm. For three wells (BR-2. 4 and 9) in the center of the field clays were found to be absent in drillcores from a depth of around 450 m down to around 1100 m, this depth range corresponding to a maximum in the presence of secondary feldspars. a&aria and albite (BRO~JNEand ELLIS, 1970). For a well at the margins of the thermal upflow. BR-6. a more uniform distribution of clays with depth is observed with secondary feldspars virtually absent in the drillcores. Instead of predicting the distribution of mineral phases within a geothermal system the thermodynamic techniques derived above may equally well be used to delineate temperature/pressure-conditions for existing or fossil thermal systems on the basis of observed mineral assemblages. On the other hand. with sufficient refinement it should be possible to derive techniques allowing the two most important parameters characterising a geothermal fluid at depth. temperature and gas-content, to be determined by use of a chemical composition of a limited number of surface discharge samples.

409

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HELGE~CIV H. C. (1971) Kinetics of mass transfer among silicates and aqueous solutions. Geochim. Cosmochim. wishes to thank W. A. work and A. J.ELLIS, R. W. HENLEY,T. M. SEWARDand J. L. HAU, JR who reviewed a draft of this paper for helpful comments and especially D. K. BIRD whose comments prompted the author to pay more attention to the effects of variations in mineral phase activities and. therefore, led to a probably somewhat more realistic presentation. Acknowledgements-The

author

SINGERS for her help with the computational

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