The effect of fluid pressure on oxygen isotope exchange between feldspar and water

Geochimlca et Cosmochimxa Acta,1978, Vol.42, pp. 235 to 239. Pergamon Press. Printed m GreatBrltam

The effect of fluid pressure on oxygen isotope exchange between feldspar and water RICHARD A. YUND and THOMAS F. ANDERSON

Department of Geological Sciences, Brown University, Providence, RI 02912, U.S.A. and Department of Geology, University of Illinois, Urbana, IL 61801, U.S.A. (Received

2 November

1916; accepted

in revised form 22 November

1977)

Abstract-The rate of oxygen isotope exchange between adularia and 2 M KC1 solution has been measured at 650°C at pressures from 125 to 4000 bar. Isotropic diffusion coefficients calculated from these data show a positive dependence on the fluid pressure. This dependence is opposite to the predicted effect of hydrostatic pressure and is attributed to the activity of ‘water’ (HzO, H+ or OH-) in the feldspar.

INTRODUCTION THE RATES of oxygen isotope exchange between hydrothermal fluids and feldspars have received considerable attention in recent years because a knowledge of these rates is fundamental for understanding rock-fluid interactions. M~RIGOLJX (1968) reported oxygen isotope exchange data for the systems adularia-distilled water and albite-distilled water. Assuming that exchange was controlled by isotropic diffusion in the solid, he calculated diffusion coefficients for oxygen in these minerals (Fig. 1). However, MCrigoux also reported that when the feldspar and fluid are far from chemical equilibrium with respect to their Na/K ratios, oxygen isotope exchange proceeds much faster (M~RIGOUX, 1968). O’NEIL and TAYLOR (1967) have demonstrated that under the latter conditions the rapid exchange rate is due to a fine-scale dissolution and reprecipitation involving a reaction front which passes through the feldspar crystal. Very rapid isotope exchange between feldspar and a hydrothermal fluid at 8-15 kbar pressure has also been reported (MATSUHISA et al., 1976). Even though oxygen isotope exchange between a feldspar and fluid is slower when the system is essentially in chemical equilibrium, the rate below 700°C is surprisingly high compared to self-diffusion of K in feldspar under the same conditions (LIN and YUND, 1972; FOLAND, 1974). YUND and ANDERSON (1974) conclude that oxygen isotope exchange between microcline or adularia and a KC1 solution in equilibrium with the feldspar is governed by the migration of an unspecified oxygenbearing species in the solid rather than by a dissolution-reprecipitation mechanism. Their conclusion is based on the following evidence: (1) if dissolutionreprecipitation were important, then the newly precipitated phase would be a disordered feldspar (sanidine), at least at 600°C or above. However, no sanidine was 235

observed optically or by X-ray diffraction below 700°C. Although a small amount of sanidine was observed in the 700°C experiments, the proportion was considerably less than the fractional oxygen isotope exchange. (2) There was no significant change in the unit cell dimensions, and hence Al/Si distribution in the tetrahedral sites, of either microcline or adularia which had undergone up to 31-35% approach to isotopic equilibrium with the fluid. This argues that the oxygen migration cannot involve extensive disruption of the bonds between oxygen and the cations in tetrahedral sites. (3) As mentioned above, the calculated diffusion coefficients for oxygen isotope exchange in microcline are significantly greater than the corresponding values for potassium isotope exchange in the same microcline under identical experimental conditions below 700°C. Such a relation is impossible in a dissolution-reprecipitation process since potassium isotope exchange would have the same rate as oxygen isotope exchange via this mechanism. (4) Experiments for markedly different reaction times were conducted on adularia and on microcline at 700°C. Diffusion coefficients calculated for each experiment were identical (for each mineral) within experimental error, which indicates that oxygen isotope exchange was governed by a, volume diffusion mechanism. These results together with more recent studies (ANDERSON and KASPER, 1975; GILETTI et al., 1976, 1978) have helped to define the temperature dependence of the oxygen rate between feldspars and hydrothermal solutions in the absence of major chemical changes. On the other hand, the rate of oxygen exchange between feldspars and essentially dry gases at 1 atm and temperature up to 1100°C appears to be much slower (see Fig. 1). YUND and ANDERSON (1974) have suggested that the difference between the results for hydrothermal fluids and dry gases is due to the penetration of ‘water’ (either as Hz0 or dissociated OH-

R. A. YUND and T. F.

236 -10

1\’

-II

-12-

\

ANDERSON

(101

Ii?), I91

f G-13B

z-14-

I

s

(9)

-15 -

-l6-

14009 0.5

0.6

1200* 0.7

1000. 0.6

600. 0.9

1.0 I/Txl03

I.1

400-c 1.2

1.3

1.4

1.5

OK

Fig. 1. The diffusivity of oxygen or an oxygen-bearing species in feldspars as a function of temperature. Curves (l)-(7) are for exchange with hydrothermal solutions at the pressures indicated below and pressure. Curve (1) for the remaining data are for exchange with air, O2 or CO* at atmospheric adularia and (2) for albite at 250-600 bar, from MBRIGOUX (1968). Curve (3) for adularia and (4) for microcline at 2 kbar, from YUND and ANDERSON (1974). Curve (5) for albite at 2 kbar from ANDERSON and KASPER (1975). Curves (6) for anorthite (An,,) and (7) for albite at 1 kbar, from GILETTI et al. (1976). (8) are maliimum values for K-feldspar from YUND and ANDERSON (1974) and (9) and (10) are for synthetic anorthite and a natural plagioclase from MUEHLENBACHS and KUSHIRO (1974). See references for details.

and HC) in the lattice and its participation in the oxygen exchange mechanism. The purpose of this study was to determine the effect of the water pressure on the rate of oxygen exchange in feldspar. This is important for interpreting oxygen exchange rates under crustal conditions and for a better understanding of the mechanism of oxygen exchange. EXPERIMENTAL

METHODS

The adularia for this study was from Kristallina, Switzerland (American Museum of Natural History No. 26545). It was exchanged twice in molten KC1 at 900°C to produce an essentially pure potassium feldspar. The 149%179pm size fraction was then used for the oxygen exchange experiments. The experiments were done in pressure vessels at 650 k 5°C. The feldspar was placed in sealed gold tubes (3 cm long for all run; except at 125 bar) with ai approximatelv eaual weight of 2 M KC1 solution whose 1sO/‘6O

ratio ‘was’approxymately + 56 per mil relative to the ‘adularia. For the 125 bar experiment, the adularia was enclosed in a crimped but unsealed platinum tube inside a sealed gold tube 12cm long containing the fluid. This was necessary to prevent rupture of the gold tube due to a higher internal than external pressure. Crushed gold strips were placed in the gold tube to minimize convection of the fluid, and the temperature of the feldspar is believed to have been within 5” of 650°C. In all other respects the experimental techniques, analytical procedures, and methods of calculating the diffusion coefficients are identical to those given by YUND and ANDERSON (1974). Oxygen was liberated from the adularia using bromine pentafluoride (CLAYTON and MAYEDA, 1963). The approach to isotopic equilibrium was calculated from the difference in the 1sO/160 ratio of the adularia before and after the experiment, corrected for the feldspar-

water equilibrium fractionation

from the data of O’NEIL and TAYLOR (1967). Diffusion coefficients were calculated by assuming isotropic diffusion in spherical grains from a well-stirred reservoir. Details of this method and a discussion of the errors involved are given by LIN and YUND (1972). DISCUSSION

OF RESULTS

The experimental data and results as well as the calculated diffusion coefficients are listed in Table 1. The uncertainty in the D values includes only the analytical uncertainty in determining the oxygen isotope composition of the adularia from each experiment ($0.2 per mil). Since we are primarily concerned here with relative exchange rates, we did not include other factors which contribute to the total uncertainty, such as: the error in the isotopic ratios of the adularia and the solution before the experiments, the variation in grain size, the departure from spherical shape, and the anisotropy of the diffusion. GILETTI et al. (1976, 1978) used an ion microprobe to determine the 18O/‘6O ratio along a traverse normal to a grain’s surface after partial oxygen isotope exchange with a hydrothermal solution. Their diffusion data for adularia agree very well with YUND and ANDERSON’S (1974) data, which were obtained using bulk isotopic exchange measured by a conventional mass spectrometer. This indicates that an isotropic diffusion model is approximately correct and that the uncertainty in the size and shape of the grains is not significant in the bulk exchange experiments. The diffusion coefficient is plotted as a function of fluid pressure in Fig. 2, which demonstrates that the

231

Fluid pressure effect on oxygen isotope exchange Table 1. Experimental results for oxygen isotope exchange between adularia and 2 M KC1 solution at 650°C

Pressure

Time

(bar)

(W

125 250 400 600 1000

336 333 342 342 336 332 336 336 334

2500 3000 4000

rate of oxygen isotope exchange between

Isotopic Equilibration (%)

Fluid/Solid (weight) 0.85 1.08 0.90 0.97 1.08 1.00 1.11 1.07 1.05

adularia

(6.4 + 1.4) (1.8 + 0.2) (2.2 _+0.3) (2.8 + 0.3) (3.6 + 0.4) (5.1 + 0.4) (5.6 + 0.4) (6.4 + 0.4) (7.9 + 0.5)

6.1 9.1 10.8 11.8 12.8 15.3 15.5 16.7 18.5

and

fluid shows a small positive dependence on fluid pressure between 125 and 4000 bar. These results and the observed pressure dependence are in good agreement with other data on oxygen isotope exchange rates in feldspar-fluid systems. Although YUND and ANDERSON(1974) did not do an experiment on adularia at 650°C interpolating their Arrhenius relation for this temperature predicts a D value of 5.2 x lo-l4 cm*/sec. The D value obtained at 2 kbar in this study (5.1 k 0.4 x lo-l4 cm*/sec, Table 1) agrees with this predicted value. The observed pressure dependence on D would account for most of the difference between M~~RIGOUX’S (1968) data for adularia at 3255600 bar water pressure and YUND and ANDERSON’S (1974) data for 2 kbar. [Compare curves (1) and (3) in Fig. 1.1 However, a similar pressure dependence for albite is not large enough to account for the difference between the results of ANDERSONand KASPER (1975) and GILETTI et al. (1976) which were at 2 and 1 kbar, respectively. [Compare curves (5) and (7) in Fig. 1.1 The observed pressure dependence is opposite in sign to the predicted dependence of diffusion in solids on pressure, which is commonly expressed as

D

(cm’/sec) x x x x x x x x x

lo-r5 lo-l4 lo-l4 10-l“ lo-l4 10-r‘+ lo-l4 lo-r4 lo-r4

D(P) = Do exp (-PAV*/RT), where AV* is the activation volume for diffusion, P is pressure, T is absolute temperature, and R is the gas constant. AV* is positive (generally some fraction of the partial molar volume) and essentially constant over a pressure interval of 4 kbar (LAZARUS and NACHTRIEB,1963). Therefore an increase in hydrostatic pressure should result in a systematic decrease in the diffusion coefficient and hence in the rate of oxygen isotope exchange. The fact that the observed diffusion coefficient increases with fluid pressure suggests that ‘water’ plays a special role in the mechanism of oxygen isotope exchange between adularia and water. Except for the point at 125 bar, the data show a linear dependence of the diffusion coefficient on the fluid pressure between 250 and 4000 bar. (The fit is equally good if water fugacity is plotted instead of fluid pressure.) Within the uncertainty of the data, an equally good fit is obtained if D is plotted as function of the square root of the fluid pressure. The penetration of ‘water’ into crystalline and amorphous silicates is often described by one of these two empirical pressure-dependence relations (MOULSON and ROBERTS, 1960; SHAFFER et al., 1974). Because we cannot choose one relation or the other, it is not

1

t

Fluid

Pressure

bars

Fig. 2. Calculated diffusion coefficient (D)as a function of fluid pressure for adularia at 650°C. The line is a least-squares fit to the data excluding the 125 bar point.

238

R. A. YUND

possible to derive any information about the mechanism of oxygen isotope exchange from the pressure dependence. The point at 12.5 bar appears to depart slightly from the linear relation of the other data in Fig. 2. Although this departure is slight, we believe it may be real and is consistent with a transition to the slower oxygen exchange rate under anhydrous conditions. Oxygen isotope exchange between feldspars and dry gases at around 7OO’C yields diffusion coefficients which are at least three to four orders of magnitude below the values obtained in hydrothermal systems (Fig. 1). Accordingly, we suggest that as fluid pressure decreases below a few hundred bars the rate of isotope exchange decreases rapidly to values approaching those obtained in ‘dry’ experiments. In addition to demonstrating the effect of fluid pressure on oxygen isotope exchange, the present results provide some additional information on the importance and perhaps unique role of ‘water’ in accelerating oxygen exchange with feldspars. At least two rather distinct roles for water can be suggested. One possibility is that the increased exchange rate in hydrothermal compared to dry experiments is due to an increased rate of oxygen transfer across the fluid-feldspar interface. However, this type of interface control is not significant in the 1 and 2 kbar hydrothermal experiments. The ‘*O/i60 profiles from the surface into the interior of the grains (GILETTI et al., 1976, 1978) are exactly what we would predict using our calculated diffusion coefficients. Hence the transfer of oxygen across the interface must be rapid compared to the diffusion rate within the feldspar in these experiments. We do not know whether this is also true for feldspars exchanged in essentially dry gas. However, even if the kinetics of the interface transfer is important in dry gases and <250 bar water pressure, we would still need to explain the rather surprising observation that the oxygen exchange rate is faster than that for the alkalis in hydrothermal experiments. Therefore we suggest that water as molecular H,O, OH- or H+ in the feldspar structure increases the oxygen exchange rate. This may occur by a mechanism similar to that suggested by DONNAY et al. (1959). They envisaged the rapid diffusion of hydrogen ions and the hydrolyzation of tetrahedral ionoxygen bonds. The slower diffusing hydroxyl ions or water molecules could then exchange with these hydrolyzed lattice oxygens. The net transport of ‘water’ during the dehydroxylation of micas also must involve the migration of hydroxyl ions or water molecules as the rate-limiting step (ROUXHETet al., 1969). Further support for this mechanism is provided by the Frank-Griggs model of slip in hydrolytically-weakened quartz (GRIGGS, 1967). According to this model, glide occurs when a dislocation moves by exchanging hydrogen bonds with a neighbor Si-O-Si bridge which has become hydrolyzed by the diffusion of water or hydrogen through the lattice.

and T. F. ANDERSON

The results of our experiments would thus be explained by the difference in the concentration of ‘water’ in the feldspar in the different experiments. Up to a few hundred bars fluid pressure, the concentration of water in the feldspar increases rapidly and so does the oxygen exchange rate. Above this pressure the concentration of water in the feldspar is only slightly higher and the effect on the oxygen exchange rate is less pronounced. At still higher fluid pressures, the increased solubility of the feldspar may lead to a dissolution-reprecipitation mechanism for isotope exchange (MATSUHISAet al., 1976). Further studies should attempt to determine whether hydrothermal treatment at a very low fluid pressure or at a low activity of water affects the oxygen isotopic exchange rate. It would also be desirable to determine the concentration of hydrogen or water in the feldspar after hydrothermal exchange at various conditions. This, however, may involve very low concentrations and the measurements may be difficult by standard methods. Returning to the more immediate geological implications of the pressure effect, we would make the following comments. Judging from the difference between M~RIGOUX’S (1968) data and YUND and ANDERSON’S(1974) data (Fig. l), the effect of fluid pressure on the rate of isotope exchange in alkali feldspar is essentially constant from 800°C to at least 5OO’C. If this relation holds at lower temperature, then the effect of water pressure would be minimal in the upper crust. Only at very shallow depths or when the water pressure is much less than the fluid pressure would the exchange rates be significantly different. Finally, it is important to reiterate that the results of this and previous investigations demonstrate that appreciable oxygen isotope exchange between feldspars, and possibly other silicates, and hydrothermal fluids can occur in the absence of major chemical or structural changes. Whether diffusion or dissolution-reprecipitation will be the dominant mechanism in nature depends on the solubility and the dissolution kinetics of the solid relative to the diffusion rate of oxygen in the solid. If the fluid and feldspar are essentially in chemical equilibrium, oxygen exchange by diffusion is likely to be the dominant mechanism, at least below 4 kbar.

Acknowledgements-We

wish to thank comments an earlier draft of the manuscript. The was supported by National Science GA-21145 and EAR75-21791 to RAY.

J. TULLIS for their helpful

Drs. B. GILETTIand and for reviewing experimental work Foindation grants

REFERENCES ANDERSON T. F. and KASPERR. B. (1975) Oxygen self-diffusion in albite under hydrothermal conditions. Trans. Am. Geophys. Union 56, 459.

Fluid pressure effect on oxygen isotope exchange CLAYTONR. N. and MAYEDAT. K. (1963) The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta 27, 43-52.

DONNAYG., WYARTJ. and SABATIERG. (1959) Structural mechanism of thermal and compositional transformations in silicates. Z. Krist. 112, 161-168. FOLANDK. A. (1974) Alkali diffusion in orthoclase. In Geochemical Transport and Kinetics (editors A. W. Hofmann et a/.), pp. 77-98. Carnegie Institution of Washington. GILETTIB. J., SEMETM. P. and YUND R. A. (1976) Oxygen self-diffusion measured in silicates using an ion microprobe. Trans. Am. Geophys. Union 57, 350. GILETTIB. J., SEMETM. P. and YUND R. A. (1978) Studies in diffusion--III. Oxygen in feldspars, an ion microprobe determination. Geochim. Cosmochim. Acta 42, 45-57. GRIGGSD. T. (1967) Hydrolytic weakening of quartz and other silicates. Geophys. J. 14, 19-31. LAZARUSD. and NACHTRIEBN. H. (1963) Effect of high pressure on diffusion. In So/ids Under Pressure (editors W. Paul and D. M. Warschauer), pp. 43-69. McGrawHill. LIN T.-H. and YUND R. A. (1972) Potassium and sodium self-diffusion in alkali feldspar. Contrib. Mineral. Petrol 34, 177-184. M~RIGOUXH. (1968) Etude de la mobilit& de l’oxygine dans les feldspaths alcalins. Bull. Sot. Fr. Mineral. Crist. 91, 51-64.

c.c * 42,) 13

239

MATSUHISA Y., GOLDSMITH J. R. and CLAYTONR. N. (1976) Oxygen isotopic fractionation in the system quartzalbite-anorthite-water. 1976 Annual Meeting, Geol. Sot. Am., Abstracts,

999.

MUEHLENBACHS K. and KUSHIROI. (1974) Oxygen isotope exchange and equilibrium of silicates with CO2 or Oz. Carnegie

Inst. Wash.

Yearb. 73, 232-236.

MOULSONA. J. and ROBERTSJ. P. (1960) Water in silica glass. Trans. Brit. Ceram. Sot. 59, 388-399. O’NEIL J. R. and TAYLORH. P. (1967) The oxygen isotope and cation exchange chemistry of feldspars. *Am. MineraL ogist 52. 1414-1437.

ROUXHETP. G., TOUILLAXR., MESTDASHM. and FRIPIAT J. J. (1969) New considerations about the dehydroxylation processes of minerals. Proc. Int. Clay Co@, 1969, Vol. 1, pp. 109-119. SHAFFERE. W., SANGJ. SHI-LAN,COOPERA. R. and HEUER A. H. (1974) Diffusion of tritiated water in P-quartz. In Geochemical Transport and Kinetics (editors A. W. Hofmann et al.), pp. 131-138. Carnegie Institution of Washington. YUND R. A. and ANDERSONT. F. (1974) Oxygen isotope exchange between potassium feldspar and KC1 solution. In Geochemical Transport and Kinetics (editors A. W. Hofman et al.), pp. 99-105. Carnegie Institution of Washington.