smectite during diagenesis: An experimental study

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The stability of illite/smectite during diagenesis: An experimental study BRUCE M. SASS’**,PHILIPE. ROSENBERG~ and JAMES A. KITTRICK~ 'Earth Sciences Department, Battelle, Pacific NorthLaboratories, Richland, WA 99352, U.S.A. wpartrnent of Geology, Washington State University, Pullman, WA 99 164, U.S.A. )Department of Agronomy and Soils, Washington State University, Pullman, WA 99164, U.S.A. (Received septet

19, 1986; accepted ilt r~is~~~

April 30, 1987)

Abstract-Illites from Goose Lake (Illinois) and Beavers Rend (Oklahoma) were equilibrated in 0.2 M and 2.0 M KClsolutions with excess kaolinite and either gibbsite, boehmite, quartz, or amorphous silica. Tefloulined reaction vessels were used to heat the charges from 25” to 25O’C at the vapor pressures of solution. Partial reversibility was demonstrated by approaching equilibria from both high and low K+/H+ activity ratios. These experiments provide evidence for but not unequivocal proof of the attainment of equilibrium. After quenching, solutions were removed by immiscible displacement techniques and analyzed to demrmine the activities of H’, K+, and SiO$. Ionic activities were calculated under experimental conditions using partial molai enthalpy and heat capacity data. The experimental data were used to construct isothermal, isobarie activity diagrams [log &+/au+) vs. log asiq] for the system [email protected]&-HzO. Stability regions have been defined for kaolinite, boehmite, and three other phases that are believed to be components of both ill&es. Assuming an Rz’-f&e stoichiometry, these components of natural ill&es have 0.24,0.67, and 0.90 K atoms per O,o(OHh repreaeming smectite, illite, and mica layers mspectively. At low temperatures (&O”C) ill&e-smectiteequilibrium is metastable with respect to calculated kaolinitemicrochne equilibrium. Retween 90’ and 110°C a phase transition involving illite and smectite occurs which stabilizes the assemblages illite-smectitekaolinite and illite-smectite-micro&e. This transition occurs at the same temperature as the ordering transition in natural I/s. Above 110°C the assemblage ilhte-smeetite is again me&stable with respect to kaolinite-microcline, but at 200°C and above ikite and smectite may

coexist stably with either kaolinite or mieroeline. However, smectite decomposes where u.& is eontrolled by quartz solubility yielding the StaMeassemblage ill&e-kaolinimquatiz I~ODU~ON

ILL~~AN~RT~C~y~n~ ingeolo&iCalaild soil environments, yet its nature and stability have not been clearly defined. An experimental investigation of the stability relationships of illite in the system KzO_ A1203-Si02-Hz0 was undertaken in an effort to resolve this problem. The term illite is used here, as recommended by BAILEY et al. (1984), to denote a non-expandable, dioctahedral, aluminous, K-bearing micaceous phase or ~om~nent that has less charge than phengite, whereas smectite refers to a related phase or component with expandable layers that generally has lower layer charge than illite. However, when used to identify a natural clay mineral (e.g. Goose Lake illite), the term illite refers to the bulk material. In many natural samples, illite layers are interstratified with other sheet silicates, typically smectite and/ or chlorite. The behavior of mixed-layer illite/smectite (I/S) during d&genetic and hydrothermal alteration in sedimentary rocks has been studied in some detail. Examinatio~ of Gulf Coast pelitic sediments by PERRY and HOWER (1970) have led to the widely accepted belief that I/S expandability and ordering are related to temperature. PERRY and HOWER (1970)found that the fraction of smectite (expandable) layers decreases with increasing depth and temperature to about 2096,

* Present address: Department ofchemistry, University of Pennsylvania, Philadelphia, PA 191046323, U.S.A.

and illite-micaquartz.

where a transition from random to ordered interstratigcation takes place with no further change in expandability. In a study of tuffaceous sediments in northern Japan, INOIJEetal. (1978) found that the expandability of I/S decreases in the direction of a hydrothermal ore zone. Furthermore, they observed that the transition from random to ordered takes place in the 20-4036 expandability range and that IS&ordering occurs in samples having < 15% expandability. HOWER and MOWATT (1966) showed that compositional variations in bulk samples of illite, smectite, and I/S were ~ntinuo~ up to a total layer charge of about -0.80 per O&OH)r . !%ROIX&eial. (1986) examined a broader range of samples using a more precise technique for determining expandability. Their statistical analysis revealed two kinds of illite layers present in approximately equal proportions. The K-contents of these layers are 0.55 and 1.O per half cell when expandabilities are less than about 50%. S~oDofi et al. ( 1986) postulate that the high-charge illite layers form by transformations of the low-charge layers in an aItemating fwhion. A third possible type of layer might have variable K-content and may be formed by repeated wetting and drying of high-charge smectite layers prior to burial diagenesis. Based on comparisons of TEM measurements of crystal thicknesses and X-ray difh-aetion (XBD) data, NADEAU et al. ( 1984a,b,c) proposed that ordered I/S is an artifact of interparticle difFraction. According to their hypothesis, the perceived ordering transition in I/S is actually caused by the disappearance of elemen-

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B. M. Sass, P. E. Rosenberg and J. A. Kittrick

2104

tary smectite particles and the formation of elementary ill&e particles. The fraction of interfaces between illite particles that adsorb water or organic molecules decreases as the illite units become thicker. When analyzed by XRD techniques, the material appears to be an ordered interstratification of illite and smectite Iayers. Longer range ordering is due to diffraction by thicker ilhte units. When the thickness of these particles becomes great enough, mixed-layering is no longer observed and the clay is identified as an end-member illite. NADEAU et al. ( 1984b,c) conclude that random I/S is a two-phase assemblage, and ordered I/S is simply end-member illite, in which apparent smectite is observed by XRD due to interparticle effects.

One of the most important, fundamental questions concerning mixed-layer minerals is whether component layers should be regarded as a single phase, or as two (or mom) phases (ZEN, 1962). According to ZEN (1962), a long sequence of a certain layer type, especially in random mixtures, could cause a component layer to behave the~odynami~ly as a unit; hence, minerals that contain com~sitionally different sequences could be polyphase mixtures. Alternatively, ordered mixed-layer minerals might exhibit single phase behavior with characteristics that depend on the proportion of layer types. If I/S behaves as a single phase solid solution, at equilibrium, the system will have one more degree of freedom than if it behaves as two phases. GARRELS (1984) used ground water compositions in apparent equilibrium with clay mineral assemblages (data from AAGAARDand HELGESON, 1983) to test heretics activity diagrams showing illite-montmorillonite (two phases) equilibrium as opposed to illite-montmorillonite solid solutions (one phase). He chose as components of the system KsO, AlzOs, SiOz, and H,O, neglecting Fe and Mg because changes in the site occupancies of these cations are not required in the conversion of interlayers from smectite to illite, and concluded that the two-phase solubility model best explained the data. ROSENBERG et al. (1985) identified a phase with a smectite-like composition which controts the solubility of Goose Lake illite at 25°C and inferred the presence of a second phase with higher K-content to account for the bulk composition. The primary objectives of this study are to identify the ~u~ib~urn~n~ouing phases in Goose Lake and beavers Bend illites, to test the proposed multi-phase and single-phase solid solution models for I/S (GARRELS,1984), and to predict phase relationships at the temperatures and pressures encountered during diagenesis. For these purposes, projection of illite compositions into the system K@-A&03-SiO&$O should prove adequate. MA~RIALS

ANT) METHODS

Materials Goose Lake and Beavers Bend ill&s were selectedfor this study because they have very different characteristics and ate thus, representative of a wide variety of ill&es. Goose Lake

ilhte, a Pennsylvanian underclay originally characterized by GIUMand BRADLEY(1939) as a non~nto~tic iBite ~n~ning small amounts of kaohnite, quartz and organic material, was obtained from the Illinois Clay Products Company; its cation exchange capacity of approximately 37 meq/ 100g is evidence that smectite is present. Mixed-layering was confirmed by GAUDETTEet ~1. (19&i), who determined the expandabiht to be between 25% and 30% and identifi~ packets of 10 1 units, which suggest that some illite and smectite layers are segregated. In the present study an XRD maximum was detected in K-saturated samples at about 17.7” 28 (CuKcr), indicating the presence of non-expandable 10 A layers(REYNOJ_DSand HOWER,1970). After Mg- or Ca-saturation and glycolation, an X-ray reflection was observed between 17.0” and 17.4” 29, indicating 10% to 20% expandability. The expandable layers are probably smectite; the presence of two reflections suggests that ilhte and I/S are segregated. Sequences with IS ordering are indicated by a XRD maximum at about 6.9” 28. The bulk composition of Goose Lake illite was reported to be Ko.r,Nao.~(All.x~~i~.~b.~~,~(OHb by GAUDE~ et ui. (1966). Beavers Bend illite, originally characterized by MANK~N and DODD(1963) as an ex-

ceptionallypure,wellcrystallizedilhte with only a smallquartz impurity,wasobtained from the Oklahoma Geological Survey.

The cation exchange capacity was determined to be 12 meq/ 100 g, which suggests that Beavers Bend illite contains little or no smectite interlayering. GALJDETTE ef ul. (1966) observed the 2 M, polytype, no mixed-layering, and a bulk composition equivalent to %.~IN~~.~~C~~.O~(A~,.~~F~~IF~~~M~~.IBT~O.M~

W.&b.&10(OH)2. The illites were pretreated to remove carbonates, iron coatings, and organic matter using the methods of K~WCK and HOPE (1963). Particle sizes between 0.2 and 5 grn were separated from aqueous suspensionsby gravity settling and centrifugation. The lower limit was chosen to minimize the amount of amorphous and poorly crystalline material. Exchange sites were then saturated with potassium by repeated soaking in 1 M WI and washing in me~ano~-~ter solutions to remove excess salt, as described by ROUTSONand ICrnwcr (1971). Other materials used include Georgia 2 kaolinite (Georgia Kaolin Co.), synthetic gibbsite (ALCOA), boehmite (crystallized from gibbsite above I 10°C), Cab-O-Sil (dehydrated amorphous silica), and quartz (Minas Gerais, Brazil), Except for grinding the quartz and sieving to retain the 5 to 50 em diameter particles, none of these materials was pretreated.

In the system K*~Al~O~Si~-H~O three solid phases can be in equilibrium with solution at a given temperature and pressure. inasmuch as natural illites appear to be multi-phase assemblages (SASSand ROSENBERG,1984; KITTRICK, 1984; ROSENBERGef al., 1985; SRODOQet al.. 1986), the solid mixtures used in these experiments may contain too many phases for all to be in upturn with solution ~~~. Thus, when the sample is equilibrated some solid phases may control the equilibrium solution composition and others will not. Furthermore, different solid assemblages may control the equilibrium solution composition as temperature and initial solution compositions are varied. The experimental strategy was to bring at least some of the component phases of natural illite to equilibrium in such a way that solution analyses could be made and equilibrium demonstrated. Experiments were conducted in 20 ml, Teflon (PTFE~lined reaction vessels which contained KCI solutions, illite, and two other solids: kaoiinite, quartz, gibbsite (boehmite), or Cab5 Sil and were designed to locate invariant points (and univariant curves) at d&rent temperatures by analyzing solutions in equilibrium with solid assemblages, Starting solutions which had total molarities of 0.2 M and 2.0 M and KCl/HCI and KCl/KOH ratios between 10’ and 106,were equilibrated with approximately equal weights of the mineral mixtures in order

Illite/smectite stability during diagenesis to bring solutions to equilibrium rapidly (solution equilibration experiments). Experiments were conducted in pairs, with identical solid mixtures and different solutions of the same ionic strength, such that equilibrium was approached from both low and high a&a,,+ ratios, in an effort to demonstrate reversibility. However, equilibrium with respect to silica was always approached from undersaturation. While these experiments provide considerable evidence for the attainment of equilibrium, they do not constitute a complete proof of equilibrium. Additional experiments (solid equilibration) with high so1ution:solid ratios (- 16011 weight ratio) were also conducted using equilibrium solution compositions (determined from solution-equilibration experiments) at the outset so that solids would undergo reaction to a greater extent. The products were examined by XRD to determine if any structural changes had taken place. The Teflon liners were tilled to about 8096capacity in all experiments and, therefore, internal pressures were controlled by the vapor pressure of solutions above 1OO‘C. The vessels were heated in constant temperature baths and ovens at ten selected temperatures between 25” and 250°C for periods of one week to four months; the longer durations were allowed at lower temperatures and for the solid equilibration experiments. After quenching the pressure vessels in water for IO to 20 minutes, the slurries from the solution equilibration experiments were transferred to polyethylene tubes containing CCL and centrifuaed at 3.0 X 10’ a for 30 minutes to sepat& the aqueous ph& by means of im&cible displacement (MUBARAKand OLSEN,1976; Krrrtuc~, 1983). Usually 4 to 7 ml of liquid were recovered and tran&red to glass (Pyrex) vials for immediate measurements of pH and dissolved silica. The solution from solid equilibration experiments was filtered through 0.22 pm Millipon membrane filters before analysis. The pH was measured at 23” + 2°C using a combination microelectrode (Microelectrodes, No. MI-410) and a Beckman PH-71 pH meter standardixed against two appropriate buffer solutions. Potassium ion activity at room temperature was measured using glass and Ag/AgCl reference microelectrodes (Microelectrodes, Inc., Nos. Ml-440 and MI-409, respectively). Standard curves were calculated using reagent KCl-HCl solutions and data from HAMER and WV (1972). Activity ratios (aK+/aH+), measured at 23” + 2°C were corrected to experimental temperatures following the approach of HEMLEY (1959) using more recently published heat capacity data, and by Debye-Hilckel calculations. A detailed account of these calculations and their results is given in SASS (I 984). Dissokd silica was measured by the molybdosilicate method (SKOUG STAD et al., 1979). The solid products of the experiments were centrifi.tgewashed twice with acetone to remove CCl, and three times with de-ionized water to remove KCl. Portions of some solids were evaporated from a slurry directly onto glass slides. Others were Mg-saturated by soaking in 1 M MgClzthree times for a total duration of one week, then washed to remove excess salt before depositing onto glass slides. When dry, all samples were stored overnight at 6O’C in a desiccator containing ethylene glycol. Solid products were characterized by X-ray dilfractometry using Ni-filtered CuKa radiation. Peek positions were deter-

mined by averaging at least two measurements at a scan rate of lo or 2’ 2Blminute. Quartz or bcehmite in the samples served as internal standards. The computer-generated XRD data of REYNOLDS and HOWER (1970) and techniques described by SRODO~~( 1980) were used to interpret peak posi-

tions and thus to evaluate expandability and ordering. RESULTS

AND DIBCUSSION

Critical experimental data are tabulated in Appendix A. Corrected solution activity measurements are shown in a series of isothermal, isobaric diagrams (Fig. 1). At

2105

each of the experimental temperatures data points can be visually differentiated into groups that are linearly distributed and intersect at common points. In general, each of the groups includes data points that were obtained using both illites and initial solutions having higher and lower ux+/uu+ values at two different ionic strengths. Thus, these groups appear to represent univariant equilibria and their intersections appear to represent invariant points. The four phases coexisting at the inferred invariant points are illite, smectite, kaolinite, aqueous solution at all experimental temperatures and illite, mica, kaolinite, aqueous solution at 90”, 125”, 175” and 25O’C (see Fig. 1). Smectite, illite and mica are defined in this paper in terms of their Kcontents as determined from the slopes of univariant phase boundaries. Isothermal, isobaric invariance suggests that illite and smectite are discrete phases rather than a single solid solution (GARRBLS, 1984; ROSENBERG et al., 1985). Mica is apparently also a discrete, component phase of both illites. Invariant points representing equilibrium between mica, kaolinite, boehmite, and solution must also exist in this system. Because this equilibrium is of less interest, fewer experiments were conducted, but approximate invariant points are shown at 90°C and above (Fig. 1). Most of the results represent metastable equilibrium due, no doubt, to the relatively low temperatures and short durations of the experiments. Although excess quartz was added in many experiments, silica concentrations in solution often exceed quartz solubility, especially at temperatures below 200°C. Apparently the aqueous solutions remain oversaturated because quartz is slow to precipitate at these temperatures. At 25°C solutions coexisting with gibbsite or boehmite are actually supersaturated with respect to quartz. Silica should react with gibbsite and boehmite, but this reaction did not control equilibrium in many of these experiments. At and below 200°C illite-kaolinite mixtures reach equilibrium with respect to &+/au+ ratios but not with respect to quartz in experiments of relatively short duration (-2 weeks at 200°C). The activity of silica gradually decreases toward quartz saturation in experiments of much longer duration (22 months at 200°C). The men&ability of silica in solution has actually been advantageous in this investigation. Because the K+/H+ activity ratio reaches equilibrium relatively rapidly as compared to the activity of SiOr, it has been possible to define lines on the stability diagrams (Fig. 1) that represent metastable univariant equilibria and to investigate equilibrium involving solutions that are supersaturated with respect to silica. A few aberrant data points do not lie close to phase boundaries (see Fig 1, 60°C); these may be due to disequilibrium with respect to silica. Compositions of solubilitpcontrolling phases Run products at every temperature were analyzed by XRD to determine whether structural or phase

B. M. Sass. P. E. Rosenberg and J. A. Kittrick

2106

,-

I-

I-

FIG.1. Activity diagrams at 25”, 60”, 90”, 125”, 175”, 2OO”,and 250°C for Goose Lake illite (without center circle) and Beavers Bend illite (with center circle) and kaolinite plus excess solid phases as indicated (quartz, Q; Cab-O&l, C; gibbsite, G; boehmite, B). Triangles and circles indicate results of I:1 and 160~1 solutioa:solid experiments, respectively. Filled and open symbols represent 0.2 M and 2.0 M KCI-HCI or KCl-KOH solutions, rqectively. The orientation of triangles indicates the direction of approach to equilibrium along vertical axes. Short dashed lines are metastable extensions of stable (solid) phase boundaries and dot-dashed lines are inferred phase boundaries.

changes had occurred. No differences in the illite diffraction maxima ( Io A, - 17 A) were observed between initial material and products of experiments with either 1:1 or 160: 1 solution:solid ratios. The K-saturated illites are completely non-expandable, even after a 48-hour ethylene glycol treatment at 6O’C. The Mg- and Casaturated products containing Goose Lake illite am 15% expandable and showed IS ordering as indicated by diEaction maxima at 6.9” 26 (REYNOLDS and HOWER, 1970). Saturation with Mg and Ca did not induce swelling of Beavers Bend illite. Because no structural changes were observed in either of the illites by XRD, it was necessary to rely exclusively upon solution.data for evaluating the re-

suits. The compositions of the solubility-controlling phases in illite were calculated from the activities of ions in solution using the method of MAITIGODand KITTRICK(I 979); it was assumed that smectite, illite, and mica can be represented by the general formula K,(Al~)(Al~i,+_,)O,o(OH)z, where x is the number of tetrahedral Al and interlayer K atoms. According to the convention of BAILEYet al. ( 1984), values of x less than 0.40 refer to smectite and x values above 0.60 refer to illite. Values of x between 0.8 and 1.O refer to mica (ROSENBERG,1984; VELDEand WEIR, 1979). For notational simplicity in the following reactions, illite (x) and (y) will refer to equilibrium-controlling phases of unknown composition. Then equilibrium between

Illite/smectite stability during diagenesis

2107

5

i m

‘2

4

0

H

3

2 -4

-3

-2

BOEHMITE

%iO,

FIG. I. (Continued)

illite(x) and kaolinite and between illite(x) and boehmite can be expressed by the equations: KXA12Al&_X0,0(OH)2 + xH+ + (1 + x/2)H20

KAA12XAl~i4-,)O,o(oH)2+ ((2 +x)/(2 + Y) 0 Y - xlK+ illite(x) = [(2 +x)/(2 + y)lKy[A121(Al~i,-,~~OH)2

illite(x)

illite( y)

= ((2 + x)/2]A12Siz05(OH), + xK+ + (2 - 2x)SiO$,, kaolinite

+[(2+x)/(2+y).y-x]H+ +I(4

-

x) - (2 +x)/(2 + Y)- (4 - YW02fwj.

(3)

(1) KAA12)(AISi,,_,)010(OH)2 + xH+ illite(x) = (2 + x)AlOOH + xK+ + (4 - x)SiO&+

(2)

boehmite. Equilibrium between two different illite(x, y) phases, where y > x, is given by the reaction:

Both natural illites contain small amounts of Fe2+ and Mg in octahedral coordination. While these cations may be conserved in reaction (3), this assumption cannot be made for reactions (1) and (2). However, the proportion of R2+ cations in octahedral coordination is small (
B. M. Sass, P. E. Rosenberg and J. A. Kittrick

2108

ratios of solutions do not change with increased duration of experiments suggesting that the role of Mg and Fe2+ is probably insignificant for purposes of this study. Equilibrium constant expressions for these reactions can be written as follows: REACTION

1

log aK+/aH+ = [(2x - 2)/x] log as, + (l/x) log 4

(4)

REACTION 2 log aK+/aH+ = [(x - 4)/x] log aREACTION

+ ( 1lx) log KZ

(5)

3

log aK+/aH+ = 3 log asa - log KS/

10 +x)/(2 +Y)*Y-xl.

(6)

These functions can be represented in activity diagrams such as Fig. 1. In Eqns. (4) and (5) the slopes of the illite(x)-kaolinite and iBite(boehmite phase boundaries are given by the terms (2.x - 2)/x and (x - 4)/x, respectively. The values of x can be determined by measuring the slopes of these lines. Extrapolation of the phase boundaries to log asiq = 0 permits the evaluation of the equilibrium constants. Equilibrium between two illite(x, y) phases is given by a phase boundary with slope of 3, according to Eqn. (6), regardless of the composition of either phase. Neither compositional variable can be determined from such a phase boundary, and the equilibrium constant for Eqn. (6) cannot be determined without prior knowledge of x and y. Furthermore, application of this equation is conditioned by the fact that it contains two independent Mass Action Law constraints. According to Eqn. (6), equilibrium between two illites(x, y) is indicated by a line with a slope of 3. Data points appear to define such a line at 25“C (Fig. 1). A line with slope of -6.23 was drawn through the set of data points that are believed to correspond to equilibrium between smectite [illite(x)] and kaolinite. The value of x was calculated using this curve and Eqn. (4), and found to equal 0.24. The value of y was not calculated at 25°C because the essential data are not available (Fig. 1). The phase with composition, Ko.2,(A12XAb.2,SiJ.,6t. O,o(OHh, probably represents the smectite component of the mixed-layer mineral. Because the bulk K-content of Goose Lake illite is 0.57 per halfunitcell [O,dOH),], it is evident that at least one additional phase of high K-content must be present in the samples. The composition of the other equilibrium-controlling phase cannot be determined from the data at 25’C. However, a phase with intermediate potassium content which has been found at all of the higher temperatures, may also be present at 25’C. The presence of this phase has been inferred in Fig. 1 (25’C, dot-dash line). The results obtained in these experiments with Goose Lake illite and kaolinite at 25°C are compared in Fig. 2 with those of KrrrrucK ( 1984) for experiments

0

FIG. 2. Activity diagram at 25°C. Triangles (Goose Lake illite), this study; circles (Goose Lake illite) and squares (BeaversBend illite), for z-300day equilibrations, KIITRKK ( 1984). The kaolinite-microcline and smectite-microcline phase boundaries were calculated as described in Appendix B. of over 300 days duration. The line representing smectite-kaolinite equilibrium satisfies all three data sets. Based on linear regression analysis of averaged data points the structural formula of the smectite in Goose Lake illite has been estimated to be I<0.2J(A12)(Ab.2$i3.,S)0,0(OH)2 (ROSENBERGet al., 1985) which is nearly identical to the formula based on the data of the present study alone. Calculated phase boundaries (see Appendix B) for kaolinite-microcline and smectitemicrocline equilibria pass through some of KIlTRICK’s ( 1984) data. Microcline had not been observed in any of the starting materials. Nevertheless, if the calculated phase boundaries are correct, then all of the equilibria in the present experiments at 25°C are metastable with respect to microcline. Thus, under these conditions, in the presence of microcline (in fact, any K-feldspar), stable equilibrium between smectite and the illite should not be possible. Phase relations at 60”, 90”, 1 lo’, 125”, 150”, 175’, 2OO”,and 250°C are similar to those at 25°C (selected results are shown in Fig. 1). Equilibrium assemblages consist of kaolinite, smectite of the same composition as that observed at 25°C and an illitic phase with a calculated composition (Eqn. 4) of K&Al+ (Al,,.&,&0,0(OH)2, which may represent the illite component of Goose Lake and Beavers Bend ill&s. At 60°C (Fig. 1) for example, some of the data points lie on a line with a slope of -6.23 representing equilibrium between &.2J0,dOH)2(smectite) and kaolinite while other data points lie on a line with a slope of representing equilibrium between I
2109

Illite/smectite stability during diagenesis probably because of slow reaction kinetics. If stable equilibrium had been attained in these experiments, all of the data would have been located at the invariant points, instead of on the univariant phase boundaries. The experimental data at 90”, 125’, 175”, and 25O’C suggest that another phase is present and controls equilibrium (see Fig 1). A line with slope of -0.22 was fit to these data points; the composition of this equilibrium-controlling phase (Eqns. 1 and 4) is Ko.soA12(Ab.gOSi3.,,)010(OH)*.Thus, this phase is Kdeficient relative to muscovite but lies within the compositional range of mica, as defined in the present study. Phase boundaries with a slope of 3 suggest equilibrium between mica and illite (Eqn. 6). Mica-boehmite equilibrium was estimated from the experimental data (Eqn. 5) and calculated values for the kaolinite-boehmite phase boundaries (Appendix B). Because equilibrium involving mica is observed at temperatures as low as 9O”C, mica is probably a natural component of both illites, rather than a product of the experiments. Kinetic effects may prevent the appearance of mica as a solubility-controlling phase below 90°C and at some temperatures above 90°C. Solid solutions between muscovite and pyrophyllite are well-known in synthetic systems (ROSENBERG,1984; VELDEand WEIR, 1979). At 200°C there is evidence of metastable equilibrium between illite and boehmite. A best-fit straight line through five data points for charges containing Goose Lake illite, kaolinite, and boehmite (Fig. 1, 200°C) has a slope of -4.97 (Eqn. 5), corresponding to K,,67/0,0(OH)Z which is exactly the composition calculated for the illite component from kaolinite-illite equilibria. In two solid-equilibration experiments at 25O’C (Fig. 1, circles) kaolinite is absent in the products, although it was included with Goose Lake illite and either quartz or boehmite in the starting materials. The experiment with quartz lies on the illite-smectite phase boundary, while the experiment with boehmite lies on the boehmite-kaolinite phase boundary close. to an extension (metastable) of the mica-illite boundary. Kaolinite dissolves during the experiments because the initial solution compositions lie outside of the kaolinite stability field (initial log uK+/aH+= 2.7, asiq controlled by quartz and boehmite solubilities) and because solid assemblages are forced to equilibrate with the aqueous solutions due to high solution:solid ratios. In the absence of kaolinite, stable univariant illite-smectite equilib-

TPC

:o" 90 110 125 150 175 200 250

1.6

2.0

2.2

2.4

2.6

2.8

3.0

3.2

3.4

,COO/Y”l”,“~

FIG. 3. Log K values for Goose Lake illite and equilibria as a function of 1000/T (OK).Curves were fit separately to the data (see Table I) between 25” and 90°C and between 1IO’ and 250°C. Coefficients for the equations are given in Table 2. rium is attained in one case and metastable mica-illite equilibrium is attained in the other. The activity diagrams (Fig. 1) suggest that Goose Lake and Beavers Bend illites contains three phases that may be classified in terms of composition as smectite, illite, and mica. The assemblages illite-smectite-kaolinite, mica-kaolinite-illite and mica-kaoliniteboehmite may be in invariant equilibrium but are usually in metastable univariant equilibrium with solutions under isothermal, isobaric conditions in these experiments. The results suggest that in addition to the expandable smectite layers, two kinds of non-expandable layers are present in natural illites, illitic layers [Kea,/ 0,0(OH)2] and mica layers [KeL90/0,dOH)2]. Similar results have been reported by SRODO~~er al. (1986) who found that illite is composed of at least two kinds of layers which have approximately 0.55 and 1.0 K atoms per 0,0/(OH)2. Equilibrium constants and enthalpies of reactions

Equilibrium constants (Table 1) for the smectitekaolinite, illite-kaolinite and smectite-illite reactions have been calculated from the experimental data using Eqns. (4) and (6). Mica-kaolinite and mica-boehmite equilibrium constants have not been calculated because of insufficient data. Generally, log K values are observed to decrease with increasing temperature for illitekaolinite equilibrium and to increase with increasing Table 1. EquilibriumconstanBofmactions. temperature for smectite-kaolinite and smectite-illite logk laqn.(311 equilibria (Fig. 3). loa K [Em (111 x-O.24 x-O.57 x.0.24. ~10.67 Above 90°C however, equilibrium constants for smectite-kaolinite and illite-kaolinite reactions decrease -3.77 2.04 -5.60 abruptly, suggesting that a phase change occurs. The -5.12 -3.39 1.92 -2.71 -4.45 1.90 data between 25” and 90°C and between 110” and -3.07 1.34 -4.29 -2.92 1.17 -3.99 250°C were fit to curves having the form, log K = u -2.62 1.31 -3.90 + b/T + c/T2, using least-squares regression; the coef-2.50 1.25 -3.91 -2.40 1.01 -3.32 ficients are given in Table 2. These curves (Fig. 3, solid -2.49 0.37 -2.07 lines) indicate that the Gibbs free energies of illite and

B. M. Sass, P. E. Rosenberg and J. A. Kittrick

2110

equilibriumconstantfunction, IogK-a+tYT+wl?

Table 2. Coefficients for the

Reedion

a. b. c. d.

Smectire - Kaolinite Smectite - Kaolinite Mite - Kaolinite llliie - Kaolinite

TemparaturePC

25-90 11 O-250 25-90 11 O-250

smectite decrease markedly, relative to kaolinite between 90” and 110°C. At lOO’C, the calculated Gibbs free energy differences are -6.0 kJ/mole and -6.2 kJ/ mole for the smectite-kaolinite and illite-kaolin& reactions, respectively. Equilibrium constants for the smectite-ilIite reaction (Fig. 3, dashed lines) were calculated from the data in Table 2 by summation in order to assess the compatibility of the calculated and experimental values. The agreement of predicted and observed log K values (Fig. 3) is evidence that the transition involves illite and smectite and not kaolinite. No abrupt change in log Kis observed for the smectiteillite reaction near 1OO’C because the magnitude of change is about the same for illite and smectite. The approximately equal differences in AG, for the illitekaolinite and smectite-lcaolinite reactions between 90” and 1lO“C suggest that the stabilities of illite and smectite are interdependent.

Discussion of phase relationships Experiments with Goose Lake and Beavers Bend illites suggest that equilibrium in solutions between 25’ and 25O’C is controlled by phases that are compositionally equivalent to smectite, illite and mica (Fig. 1). The solid assemblages illite-smectite-kaolinite, micaillite-kaolinite and mica-kaolinite-boehmite are, by implication, at invariant equilibrium with solutions at their vapor pressures. The kaolinite-microcline and kaolinite-smectite univariant lines (Fig. 2) intersect at an invariant point at which kaolinite, microcline and smectite are in equilibrium with solution. Kaolinite-microcline is the stable assemblage above log a&a,.,+ = 4.2, while kaolinite-smectite is stable below this value. Figure3showsthevariationoflogKwith l/T(“K) suggesting that the illite and smectite components of the natural illites undergo phase changes between 90” and 110°C that lower their Gibbs free energies by ap proximately equal amounts. The exact nature of this phase transition is unknown, but it may correspond to the transition between I/S random and I/S ordered that has been observed at approximately 100°C during diagenesis (PERRY and HOWER, 1970;HOWER et al., 1976). Although XRD maxima do not differ for illites equilibrated above and below the transition temperature, the similarity of the transition temperature to that of ordering suggests that these phenomena are related. A change of approximately 6 kJ/mole is reasonable for the proposed transition. The difference in the free

a

27.75 -11.64 9.60 -10.60

bXl0‘3

CXlo-5

-18.69 8.975 -5.128 16.42

27.71 -21.85 8.571 -33.67

energies of formation of 7 A and 14 A clinochlore, for example, is 4.60 kJ/mole (HELGESONet al., 1978),

while between kaolinite and dickite the difference is 3.06 kJ/mole (ROBIEet al., 1978) at 25°C. Inasmuch as the observed difference in AGg is somewhat larger than in these isochemical examples, it is possible that some dehydration of smectite accompanies the phase transition. The difference in AGg for the phase change from halloysite to kaolinite which involves dehydration, is 18.65 kJ/mole (ROBIEet al., 1978). Figure 4 is a composite diagram showing the illitesmectite-kaolinite invariant point at different temperatures and vapor pressures. The data points were calculated from the regression data in Table 2 in order to obtain smooth univariant curves. With increasing temperature, solutions in equilibrium with illite-smectite-kaolinite are shifted to lower K+/H+ activity ratios and generally to higher aqueous silica activities. A discontinuity in the univariant curve above 9O”C, due to the illite-smectite phase transition, is accompanied by an abrupt decrease in log aK+/aH+. At 25”C, illite-smectite equilibrium has been shown to be metastable with respect to microclinekaolinite equilibrium (Fig. 2). Similar calculations have been

6

i -4

-3

-2

-1

‘w %i02 Flc. 4. Composite diagramshowingil&-smecti~kaolinite invariant points at different temperatures and at the vapor pressures of solution. The invariant points were calculated from the data in Table 2, and are u-mnected by univariant curves; the break between 90’ and 110°C may be caused by a phase change in illite and smectite.

lllite/smectite stability during diagenesis

2111

mica may approach the muscovite-pyrophyllite solid solution series described by ROSENBERG( 1984). In contrast to the experimentally determined stability of mica relative to end-member muscovite, when high temperature and/or calorimetric data for muscovite are extrapolated to 25°C the log (uk+/au+) for kaolinite-muscovite equilibrium is approximately 3.5 (e.g. BOWERSet al., 1984) which indicates that muscovite is much more stable than mica. This must certainly cast doubt on the current extrapolations of muscovite data to 25’C. The high temperature and/or calorimetric data may also exaggerate the stability of microcline at low temperature, but in the absence of actual low temperature experimental evidence to the contrary, the extrapolated microcline data (Appendix B) are shown in Fig. 2. Application to diagenesis I

-4

I

I

-3

-2

log%iO,

FIG.5. Phase relationships at selected temperatures derived from the experimental data and calculated phase boundaries

for microeline-kaolinite (Appendix B). Symbols: I = illite, K = kaolinite, M = rnicrocline, and S = smectite. Vertical lines show quartz saturation.

made at other temperatures using the data in Appendix B; the predicted phase relationships are shown at selected temperatures in Fig. 5. At 25” and 90°C the assemblage illite-smectite-kaolin&e is metastable (Fig. 5, dashed curves), because the microcline-kaolinite boundaries lie at lower aqueous silica activities than the illite-smectite-kaolinite invariant points. At 110°C the invariant point is slightly to the right of the microcline-kaolin&e boundary and, therefore, the assemblage illite-smectite-kaolinite is again metastable. At 200°C and above, illite and smectite are stable and may coexist with either kaolinite or microcline. An extrapolation of the univariant curve representing illite-smectite-kaolinite equilibrium at high temperatures (> 1lO”C, Fig. 4) intersects the microcline-kaolinite boundary slightly below 110°C. Thus, illite-smectite equilibrium may be stable in the narrow temperature range above the transition temperature and below 110°C. With increasing temperature illite-smectite equilibrium is again metastable with respect to microcline-kaolinite equilibrium ( 110” to about 19O’C) becoming stable only above 190°C. The activity diagrams (Fig. 1) must be interpreted in relation to the stability of muscovite and K-feldspar at the temperature of interest. If the experimentally determined kaolinite-gibbsite line (KIl’TRICK, 1980) is superimposed upon the line for natural muscovites in equilibrium with gibbsite (MATTIGOD and KITTRICK, 1979), the resulting log @x+/au+) for the gibbsite-kaolinite-muscovite triple point is approximately 8.2 for 25’C at a log Us of -4.5, indicating that mica [K& O,dOH)z] is stable relative to muscovite at 25OC. The

During lithification of fine-grained sediments where solid/solution ratios are high, pore fluid compositions will be determined by the solids that they contact. In sediments that contain quartz, microcline, kaolinite and illite (including ill&e, smectite and mica compo nents) pore fluids can equilibrate with up to three solid phases. At temperatures below the phase transition (N lOO’C), stable invariant equilibrium at constant temperature and pressure is possible for the solid assemblage microcline-kaolinitequartz and, at solution silica activities higher than those representing quartz saturation, for the solid assemblage microcline-kaolinite-smectite (see Figs. 2 and 5). At temperatures immediately above the phase transition and again above 190°C in the presence of solutions that are supersaturated with respect to quartz, the solid assemblage illite smectite-kaolinite and (at even higher silica activities in solution) illite-smectite-microcline are stable. Micabearing assemblages coexist with solutions that are largely undersaturated with respect to quartz solubility (see Fig. 1). Smectite reacts with K-feldspar above 50°C (PERRY and HOWER, 1970) to form I/S, chlorite and quartz, the proportion of illite in I/S gradually increasing to 55% before ordering begins at approximately 100°C (SRODOIG and EBERL, 1984). Although experimental data for this reaction are not available, the assemblage illite-smectite appears to be stable immediately above the phase transition (< 110°C) and at higher temperatures (> 190°C). If the observed phase change corresponds to the I/S ordering transition (PERRY and HOWER, 1970; HOWER et al., 1976) the driving force for the transition may be the mutual stability of illite and smectite. According to HOWER et al. (1976) with increasing depth of burial, K-feldspar disappears, kaolinite-content stabilizes after declining and I/S undergoes ordering with a concomitant decrease in expandability to about 20% (or less in other areas, SRODO~~and EBERL, 1984); the assemblage I/S-ordered, kaolinite persists from 95’ to 175’C. Thus, illite-smectite-kaolinite ap-

B. M. %ss, I’. E. Rosenberg and J. A. Kittrick

2112

pearsto be an ~u~ib~urn assemblage at ~rn~mt~ considerably above the phase transition, whereas the experimental data suggest that the stabihty of illitesmectite-kaoIinite assembIages is interrupted between 100” and 190°C. Since natural illites contain small but significant amounts of Mg (and Fe), their stabilities will depend on the Mg-content as well as the K-content of coexisting solutions; an inverse relationship exists between log (Use+/&) and log (uk+/ffu+) at fixed values of log asi%, The isothermal, isobaric stability of illite is limited by kaolin&e-muscovite equilibrium at low values of log &gl+/&+) and by i~ite~hlo~te at high values of Iog (a~&&+); it can be shown from the diagrams Of BOwBRSet al. (1984) that corresponding values of log (ux+/un+) differ by about 2 units. Thus, the presence of Mp in solution enhances the stability of illite and may account for the apparent stability of illite-smectite relative to microcline-kaolinite between 110” and 190°C in nature. Experiments designed to test this inference are now underway. Furthermore, once K-feldspar has dissolved and illite has crysmlhzed, kinetic constraints may prevent the crystallization of K-feldspar even if it becomes thermodynamically stable. Experiments at temperatures below 90°C show that ate-smote ~~lib~urn is metastable with respect to k~~ni~rnic~~ne; hence one or both of these phases should dissolve. The fact that illite and smectite coexist at low temperatures in nature suggests that metastable equilibrium can persist for extended periods of time at this temperature. Although mica has been shown to be an equilibriumcontrolling phase in natural illite, thermodynamic considerations dictate the incompatibility of mica and smectite layers. The disappearance of I/S X-my diffraction reflections accompanied by the reduction of ex~~i~ty to 5% or less, occurs at about 200°C (PRODS& and EBERL, 1984). VELDE( 1977) proposed that mixed-layer I/S decomposes above 100°C yielding illite and chlorite. Decomposition of smectite would be expected if silica activities in solution are controlled by quartz solubilities at elevated temperatures. Stable assemblages in the quartemary system K@-A1203SiOr-Hz0 would be illite-kaolinitequartz or illite-micaquartz (see Fig. 1,250’C).

(Ab,wSip.,o)OIO(OH)2]. The illite and mica phases may comprise the non-expandable components of natural I/S as su8gested by SRODOrj ef al. ( 1986). Thus, in agreement with the conclusions of ROSENBERG et al. ( 1985), Goose Lake illite appears to behave as a multiphase mixture. Below a critical temperature (- lOO*C) illite and smectite are incompatible in the presence of K-feldspar. At low temperatures illite-smectite equilibrium is metastable with respect to kaolinite-microcline equilibrium but between 90” and 110°C the Gibbs free energies of illite and smectite decline by about 6 kJ/ mole, thus stabilizing the assemblages illite-smectitekaolinite and illite-smectite-micro&e. This transition occurs at approximately the same temperature as the ordering transition in natural I/S (PERRYand HOWER, 1970; HOWERet al., 1976). Thus, at temperatures just above the phase transition I/S is a stable ordered mixedlayer (mixed~pha~) mineral. With increasing temperature illite-smectite equilibrium is again metastable with respect to microcline-kaolinite equilibrium (1 10°C to 19O’C) but kinetic constraints may prevent the reappearance of K-feldspar during diagenesis. Furthermore, the presence of Mg (and/or Fe) in solution may enhance the stability of illite-smectite relative to microcline-kaolinite in nature. At 200°C and above illite and smectite may coexist stably with either kaolinite or microcline. However, smectite decomposition occurs where silica activities am controlled by quartz solubility yielding the stable assemblages, illitekaolinitequa~ and illit~mi~qua~. The present investigation in which illite and smectite compositions are projected into the system K@-Al~03SiOr-HzO, provides an insight into the stability relations of l/S. Further studies in this system and in the system with Mg and Fe are now in progress in an effort to confirm the above inferences and to fully define the stability relationships of illite.

CONCLUSiONS

Acknowledgements--This study was supported by the donors of the Petroleum Research Fund, administered by the Amer-

The stabihty relation~i~ of Goose Lake and Beavers Bend ilhtes have been investigated in the system K20_A1rOr-Si@-Hz0 between 25” and 250°C; the presence of small amounts of Mg and Fe*’ in these illites has been neglected. Me&stable equilibrium was apparently attained in both 0.2 M and 2.0 M solutions, by starting with solutions having hi8h and low K+/H+ activity ratios, indicating that K+/H+, and SiOz are solubility-controlled species. Three phases in these illites have been found to equilibrate with kaolinite and boehmite over a broad temperature range (25” to 250%). These experiments provide evidence for but

not unequivocal proof of the at~nment of equilibrium. Based on their chemical compositions, which were calculated from solubility measurements assuming an R2+-free stoichiometry, these phases have been referred to as smectite [K,,.2s(A12)(Ab.2&.7&3i~OH)~], illite

IKo.6,(Ala)(Ab.6,Si3.~~)O,~(OH)21, and mica Ko.wOW

ican Chemical Society under Grants PRF#12049-AC2 and 15X%AC!.

Editorial handling:J. I. t&ever REFERENCES AAGAARDP. and HELGE~~NH. C. (1983) Activity/compo-

sition relations among silicates and aqueous solutions: II. Chemical and thermodynamic consequencesof ideal mixing of atoms on homological sites in montmorillonites, illites

and mixed-layer clays. Clays Clay Minerals 31, 207-2 17.

BAILEYS. W., BRINDLEY G. W., FANNINGD. S., KODAMA H. and MARTINR. T. (1984) Report of the Clay Mineral

2113

Illite/smectite stability during diagenesis Society, Nomenclature Committee for 1982 and 1983. Clays Clay Minerals 32,239-240. BOWERST. S., JACKSON K. J. and HELGESONH. C. (1984) EquilibriumActivityDiagramsfor CoexistingMineralsand Aqueous Solutions at Pressures and Temperatures to 5Kb and 600°C. Springer-Verla& New York. 397p. GARRELSR. M. (1984) Montmorillonite/illite stability diagrams. Clays Clay Minerals 32, 16 I- 166. GAUDE~E H. E., EADES J. L. and GRIM R. E. (1966) The nature of illite. Clays Clay Minerals, 13th Natl. Conf., 3348. GRIM R. E. and BRADLEYW. F. (1939) A unique clay from the Goose Lake, Illinois, area. J. Amer. Ceram. Sot. 22, 157-164. HAMERW. J. and WV Y. C. (1972) Osmotic coefficients and mean activity coefficients of uni-univalent electrolytes in water at 25°C. J. Phys. Chem. Ref: Data 1, 1047-1099. HELGESONH. C., DELANEYJ. M., NF_SBITH. W. and BIRD D. K. ( 1978) Summary and critique of the thermodynamic properties of rock-forming minerals. Amer. J. Sci. 278-A, l-229. HEMINGWAYB. S., Hess J. L. JR. and ROBINSONG. R. JR. ( 1982) Thermodynamic properties of selected minerals in the system AlrOs-CaO-SiOr-Hz0 at 298.15 k and 1 bar ( 10’ Pas&s) pressure and at higher temperatures. U.S. Geol. Suw. Bull. 1544, 70~. HEMLEYJ. J. (1959) Some mineralogical equilibria in the system K@-AlrOs-SiOrH20. Amer. J. Sci. 257.241-270. HOWER J.and MOWATTT. C. (1966) The mineralo8y of ilhtes and mixed-layer illite/montmorillonites. Amer. Mineral. 51, 825-854. HOWERJ., ESLINGERE. V., HOWER,M. E. and PERRYE. A. (1976) The mechanism of burial metamorphism of a& laceous sediments: 1.Mineralogical and chemical evidence. Bull. Geol. Sot. Amer. 87,725-737. INOUEA., MINATOH. and UTADA M. (1978) Mineralogical properties and occurrence of illite/montmorillonite mixedlayer minerals formed from Miocene volcanic &ss in WagaOmono district. Clay Science 5, 123- 136. KEENAN J. H., KEYES F. G., HILL P. G. and MOORE J. G. (1978) Steam Tables, Thermodynamic Propertiesof Water Including Vapor, Liquid, and Solid Phases (International System of Units-SI). John Wiley and Sons, 156~. KITIXICKJ. A. (1980) Gibbsite and kaolinite solubilities by immiscible displacement of equilibrium solutions. Soil Sci. Sot. Amer. J. 44, 139-142. KITTRICKJ. A. (1983) Accuracy of several immiscible displacement liquids. Soil Sci. Sot. Amer. Proc. 47, 10451047. KITTRICKJ. A. (1984) Solubility measurements of phases in three ilhtes. Clays Clay Minerals 32, 115-I 24. KITTRICKJ. A. and HOPE E. W. (1963) A procedure for particle-size separation of soils for X-ray diffraction analysis. Soil Science %,3 19-325. LEWISD. (197 I) Studies of redox equilibria at elevated temperatures. I. The estimation of equilibrium constants and standard potentials for aqueous systems up to 374°C. Arkiv for Kemi 32,385-404. MANK~N C. J. and DODD C. G. (1963) Proposed reference illite from the Ouachita Mountains of southeastern Okrahoma. Clays Clay Minerals, 10th Natl. Conf., 372-379. MA~GOD S. V. and KIURICK J. A. (1979) Aqueous solubility studies of muscovite: Apparent nonstoichiometric solute activities at equilibrium. Soil Sci. Sot. Amer. J. 43. 180-187. MLJBARAK A. and OVEN R. A. ( 1976) An improved technique for measuring soil uH. Soil Sci. Sot, Amer. J. 40.880-882. NADEAUP. H.;TAI~ J. M., MCHARDY W. J. and WILSON M. J.(1984a) Interstratified XRD characteristics of physical mixtures of elementary clay particles. Clay Minerals 19, 67-76.

NADEAUP. H., WILSONM. J., MCHARDY W. J. and TAIT J. M. (1984b) Interparticle diffraction: A new concept for interstratified clays. Clay Minerals 19, 757-769.

NADEAUP. H., WILSONM. J., MCHARDY W. J. and TAIT J. M. (1984c) Interstratified clays as fundamental particles. Science 225,923-925. PERRYE. A. and HOWERJ. (1970) Burial diagenesis of Gulf Coast pelitic sediments. Clays Clay Minerals l&165-177. REYNOLDSR. C. and HOWER J. (1970) The nature of interlayering in mixed-layer illite-montmorillonites. Clays Clay Minerals 18, 25-36. ROBIE R. A., HEMINGWAYB. S. and FISHER J. R. (1978) Thermodynamic properties of minerals and related substancesat298.15Kand 1bar(lOJPascals)pressureandat higher tempemtures. U.S. Geol. Suw. Bull. 1452, 456~. ROSENBERGP. E. (1984) Synthetic muscovite solid solutions in the system KrO-A&O3SiOr-H20. Geol. Sot. Amer. Ann. Mtg. Abstr.. 639. ROSENBERGP. E., KITTR~CKJ. A. and SASS B. M. (1985) Implications of illite/smectite stability diagrams: A discussion. Clays Clay Minerals 33, 56 l-562. ROUT~ONR. C. and KI-I-~RICKJ. A. (1971) Illite solubility. Soil Sci. Sot. Amer. Proc. 35, 7 14-7 18. SASSB. M. (1984) Stability relationships of iBite in solutions between 25” and 250°C. M.S. thesis, Washington State Univ. SASSB. M. and ROSENBERGP. E. ( 1984) Multi-phase solubihty in natural ii&s between 25” and 250°C. Program and Ab stracts, Ann. Mtg., Clay Minerals Society, Baton Rouge, 101. SKOUGSTADM. W., FISHMANM. J., FRIEDMAN L. C., ERD MAND. E. and DUNCANS. S. (1979) Techniques of water resources investigations of the United States Geological Survey, Chapter Al: Methods for determination of inorganic substances in water and fluvial sediments. Book 5. U.S. Government Printing Office, Washington, D.C. SRODorj J. (1980) Precise identification of ilhte/smectite interstratifications by X-ray powder diffraction. Clays Clay Minerals 28,40 1-4 I 1. SRODO~;~ J. and EBERL D. D. (1984) Illite. In Micas (ed. S. W. BAILEY),Reviews in Mineralogy Vol. 13, pp. 495544. Mineralogical Society of America. SROE& J., MORGAND. J., ESLINGERE. V., EBERLD. D. and URLINGER M. R. ( 1986) Chemistry of Elite-smectite and end-member illite. Clays Clay Minerals. 34.368-378. VELDE B. (1977) Clays and Clay Minerals in Natural and SyntheticSystems. Else&r, 2 I8p. VELDEB. and WEIR A. H. ( 1979) Synthetic ilhte in the chemical system KrO-A120s-Si02-Hz0 at 300°C and 2kb. In Developmentsin Sedimentology(ads. M. M. MORTLANDand V. C. FARMER),Vol. 27, pp. 395-404. Elsevier. WALTHERJ. V. and HELGESON H. C. (1977) Calculation of the thermodynamic properties of aqueous silica and the solubility of quartz and its polymorphs at high pressures and temperatures. Amer. J. Sci. 277, 1315- I35 1. ZEN E-AN. (1962) Problem of the thermodynamic status of the mixed-layer minerals. Geochim. Cosmochim. Acta 26, 1055-1067.

APPENDIX T*d~llm

zz& 2321 2322 2327

ancm Quwa Ouwo

:::: 2340 2341 2323 2324 2332

%s OLKlcI OLllM Ovwa QWM QWA

z% 2376 2377 2378

zkz OuKlo Wva MM

71 71 34 34 13 13 132 71 71 79 2 81 01 01

0.2 0.2 0.2 0.2 0.2 0.2 0.2 2.0 2.0 0.2 0.2 0.2 0.2 2.0 2.0

3s 6.50 3.59 OS0 3.59 9.25 a.25 4.03 6.52 2.07 6.38 3.46 9.25 3.87 IO.09

A

3.41 3.40 3.40 3.44 3.54 3.58 3.38 3.a 3.50 3.01 i::; 3.03 3.75 3.76

7.42 7.50 7.49 7.56 7.66

0.08 0.93 0.85 0.01 0.17

::: 6.90 6.01 7.50 7.55 8.25 7.75 7.64 7.65

:::: -0.04 -0.1 0.79 0.09 0.79 01s -0.07 -O.oI

6.54 6.57 6.(14 :::i 6.80 6.62 e.94 6.97 6.80 ::ii :::t 7.73

B. M. Sass, P. E. Rosenberg and J. A. Kittrick

2114

PO%

2QeC 2250

G-

12

0.2

2251 2” 2240 2241 2242 2275 2367

G%z OuMl GbWQ GWAG GmiE‘G GlAUG

36 47 1s 12 33 64 47 19

iii: 0.2 2.0 2.0 2.0 2.0 2.0

EE 2263 2284 2266

x Moo Ouwo G-

32 “: 43 18

0.2 2.0 2.0 2.0

7.31 6.36 6.53 6.62

2262 2263 2278 2265 2266

BBAUG Ewlua 66/K/O BBilUG BBWG

12 34 30 49 49

0.2 0.2 0.2 2.0 2.0

Tz2s.s

GGlJKIO GGblUG GQ-

26 26 26 26 32 32

2290 2362 2268 2315 2363 ;:N;@)

GiiwQ ux: GMIQ GbUJG GMVC fLU&

29 ii 27 29 42

2361) 2372(d)

GG-

2409 2410 Z2Sl 22s2 2406 2401 2402

BBWO BBAUO BBKIO BBI100 BB#O 6Bnm B&lb3

z:; 2.0 2.0

:::; 10.09

3.1. 3.14 3.33 3.29 3.24 3.23 3.45 4.32 4.14 3.99 3.72 3.72 4.06 4.06

2.0 2.0 2.0 2.0

3.67 7.10 3.97 1O.M)

3.16 3.40 3.76 3.77

7.21 6.35 6.44 6.71

-0.06 -0.01

2.01

3.06 3.04 2.81 2.93 2.96 2.96 3.02

7 63 7.57 7.29 7.47 6.25 6.25 7.30 7.41 7.60 7.79 7.64

0.67 0.94 0.63 O.SO -0.10 -O.oS 9.05 0.83 0.91 0.62 0.69

6.75 6.63 6.46 6.57 6.43 6.42 7.63 8.56 6.69

2.71 2.61 2.31 2.43 2.62 3.40 3.34 3.32 3.61

6.56 6.70 5.32 3.09 4.41 7.25 7.22 6.16 6.34

0.63 0.W 0.63 0.91

5.75 5.60 4.49

2.72

0.65

3 E%i 2343 GIMm

II: 50

0.2 0.2

~~~ 2m4 2342 2267 2266

z&z Gwm Gwm &4UG GL4’X

:: 41 59 (II 61

zi 2344 2379 2360

z OLMlB GIAW GUlW

3b 3 61 61

0.2 2.0 2.0 2.0 0.2 0.2 0.2

2267 2266 2393 2394

BBIKIO B&%/O 6BWG B&W0

62 62 69 69

4.32 3.4s 9.23 9.23 3.00 5.33 10.29 6.63 ::Zi 2.94

:::: 9.16 3.67 6.W IO.23 1.36 :::: 9.16 2392 2391 :::

B&?UG BBmo s

46 ii:

i:: 0.2

2427

BME

32

!if

zz

m

::

iI::

g4

plc&

2

2.0

IIp% 2261 22o6

Gwm Glnm

4S 12

0.2 2.0

2: 2276 2262

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z

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:ii

2313 2314 2397 2396

BMCiQ B&X/G BMUG BBWG

2250 2239 2403 2404

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i:;: 3.911 3.67

7.00 6.30 7 6, 7.50 6.11

0.96 0.97 0.01 0.97 9.01

F:Z

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0.14 0.93

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:ii i41 6.29 %

2.0

2.94 3.66 3.59 2.63 3.44 4.76 2.29 6.65

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135 133 41 41

2.0 2.0 2.0 2.0

2.29 6.63 2.w IO.09

2d6 2.96 3.44 3.43

3.06 3.43 3.30 5.29

-0.07 Q.oS

5.17 5.57 5.43 5.42

GM/G GIJWG GUKC GUKC

46 46 46 46

0.2 0.2 0.2 0.2

2.56

ZEt GLWQ QGUlUi Gi6U3

6 40 46 52 52

2.0 2.0 2.0 2.0 2.0

2.66 2.67 3.71 3.75

3.46 5.53 3.47 3.32 4.6S 4.52 4.25 4.24 4.90 4.99

0.67 0.76 0.60 0.86

i;ii 2231 2232 2303 2304

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24oC 2466 2311 2312 2423

BBlwQ B5wG BBNM MtlW BGWG

33 33 62 32 26

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26 19 19

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0.64 0.91 .o.o6 .O.OS 0.05 0.93

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2.37 9.16 2.29 6.66 2.37 9.16 2.69 IO.09

LIIp% 2264 2265

m W

16 16

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O.S2 0.89

2199 2lBS 2245 2307 2302

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317 31 36 36

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2363 2310 2307 236S

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2.39 2.66 2.92 2.64

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4.73 4.70

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1.10

lms

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402

4.56 4.15 4.34 4.60 4.60 4.65 5.05 5.25

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3.60

2273 2274 2273 2272 2317 2316

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2.21

3.72

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::: 4.76

2.37 2.42 2.40 2.43 2.42

5.93 5.66 5.55 3.37 3.16

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2.24 2.26 2.16 2.1, 2.43 2.47

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3.46 3.69 2.76 2.53

0.2 0.2 2.0 2.0 2.0 2.0

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2.20 1.19 2.19 2.17

0.67 0.73 -0.20 -0.12 .O.l2

2.92 2.69 2.30

::

Il:i: 2.0

:::: 2.62

2.02 2.34 2.32 2.55

3.96 3.7S 2.46 2.18 2.56 3.10 3.95 4.04 3.07

36 36 16 16 43 3s 39

0.2 0.2 2.0 2.0 2.0 2.0 2.0

1.33 5.S6 2.29 4.76

2.22 2.23 2.23 2.13 2.1s 2.52 2.57

4.42 4.42 4.37 2.33 2.2o 2.24 2.36

0.91 0.74

111

i:::

:: fi: IO.09

1.60

0.63 0.93

-0.12

:::: 3.w 4.27 3.70 ::g 1.40 3.10 3.31 ::Z 1.76 ::II: 4.60 3.46 3.09

$2

1.93

2.32 2.61 2.95 2.94 2.76 3.45 3.62 4.49 2.50 2.12 2.16 2.22

w

APPENDIX B

Calculation of microciine-kaolin&e and kaolinite-boehmite stability boundaries at 2PC and higher temperatures In order to include a region of potassium fehdsparstability in the ex~~rnen~lly derived phase diagrams, ~uilib~um between the low temperature potassium feldspar, microcline. and kaolinite was calcuiated from available thermodynamic data. Equilibrium is described by the following mass action equation: 2 Microcline + 2H+ + Hz0 = Kaolinite + 2K+ + 4SiO$,,.

03.1)

The equilibrium constant for this reaction can be written in the following form for graphical representation: log aK+/un+ = -2 log a& + ( I/2) log Kr,

(B.2)

Vahtes of the equilibrium constant (&) were calculated following the method of ROBIEet al. (t 978); thermodynamic data used in these calculations are given below. The locations of all other pertinent phase. boundaries are easily determined for each isothermal, isobaricactivity diagram where the kaolinite-microcline phase boundary has been drawn. The smectite-microcline boundary must pass through the invariant point, smectite-microcline-kaolinite-solution, and has a slope of - 1.48. Similarly, the illite-microciine boundary has a slope of -2.34. Phase relationships are shown in Fig. 5.

lllite/smectite stability during diagenesis Thermodynamic data’ pertaining to microcline-kaolinite equilibrium (Eqn. B. 1) Aeg, = 70.79 kJ

AC;= 5.555 +O.O2491T-2.813X

10-5T2kJ/oKb.

’ Thermodynamic data were obtained from the following sources: Microcline, ROMEef al. (1978); Kaolinite, HEMING WAYet al. ( 1982); H20, KEENANet al. ( 1978); SiO! , WALTHER and HELGESON (1977); K+, LEWIS(197 1). b T is temperature in Kelvins.

c Thermodynamic data were obtained from the following sources: boehmite and kaolinite, HEMINGWAYet al. (1982); HrO, KEENANet al. (1978); Si@, WALTHERand HELCFSON

( 1977).

d T is temperature in Kelvins.

The region ot hoehmite stability can be calculated using similar methods. Equilibrium is described by the following mass-action equation: Kaolinite = 2 Boehmite + 2SiO&,,, + H20.

A.!&=O.l565kl

2115

(B.3)

The equilibrium constant for Eqn. (B.3) can be written as follows: log a& = ( l/2) log KT (B.4) Values of the equilibrium constant were calculated from the thermodynamic data given below. The boehmite-mica phase boundary must pass through the invariant point, boehmite-kaolinite-mica-solution, and has a slope of -3.44. Thermodynamic data’ pertaining to kaolinite-boehmite equilibrium (Eqn. 8.3) AHbs = 89.98 kJ AS’&, = 0.0857 kJ/K AC;= -2.627 +O.O1196T- 1.337 x 10-sT2kJ/“Kd.