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Granite-fluid interaction at near-equilibrium conditions: Experimental and theoretical constraints from Sr contents and isotopic ratios
CHEMICAL GEOLOGY INCLUDING
ISOTOPE GEOSCIENCE
Chemical Geology 121 (1995) 145-154
ELSEVIER
Granite-fluid interaction at near-equilibrium conditions: Experimental and theoretical constraints from Sr contents and isotopic ratios 1 Pierpaolo Zuddas a, Franqois Seimbille b, Gil Michard a aLaboratoire de GEochimie des Eaux, UniversitE Paris 7-Denis Diderot and lnstitut de Physique du Globe de Paris 2, place Jussieu, F- 75251 Paris Cedex 05, France bLaboratoire de GJochimie et Cosmoehimie, Universit~ Paris 7-Denis Diderot and lnstitut de Physique du Globe de Paris 2, place Jussieu, F- 75251 Paris Cedex 05, France
Received 5 May 1994; revision accepted 6 December 1994
Abstract Fluid-granite interaction at near-equilibrium conditions was investigated by doping an experimental system with S4Sr. Reactions were performed at 453 K, saturated vapour pressure, and fluid/rock ratio of 20. The initial composition of the experimental fluid corresponded to the theoretical saturation with respect to kaolinite, quartz, laumontite, low-temperature albite, calcite and adularia. Sr was used to trace parent-mineral dissolution, neogenic phase precipitation, and to identify the reactions at Srz+ steady state. The interaction is described, initially, by the isotopic mixing of the solution (84Sr/S6Sr)solutionratio and the granite (87Sr/ 86Sr)rock ratio. However, in the rock end-member every main mineral has a different 87Sr/S6Sr ratio. The evolution of 87Sr/86Sr ratio in solution as a function of 84Sr/S6Sr showed that in the early stage of the interaction (where the dissolution of the rock is the dominant reaction) STSr-enriched phases, such as biotite, strongly influence the isotopic budget. After some time, the 87Sr/ 86Sr ratio of the fluid approaches the Sr isotopic ratio of the bulk rock, suggesting that fluid-granite isotopic equilibration is attained. When the dissolution of the granite is stopped and once the Sr2÷ concentration reached steady state, variations of the 84Sr/S6Sr ratio with time can be interpreted by an isochemical dissolution-precipitation of neogenic phases.
1. Introduction Continental w a t e r - r o c k interaction mainly involves the reaction between meteoric water and granitic rocks. Chemical studies of hot springs, direct measurements in geothermal drill-holes, as well as experimental and theoretical investigations indicate that dissolution of parent minerals in host rocks and subsequent precipitation o f neogenic minerals are the dominant geochem[CA] 1 IPGP contribution No. 1339. 0009-2541/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSD10009-254 1 (94)001 59-6
ical reactions o f hydrothermal processes (Arnorsson et al., 1983; Giggenbach, 1984). In the temperature range o f 283-523 K, and at low fluid/rock ratios ( w / r ) , fluids at equilibrium produce new isochemical mineral assemblages (Michard and Fouillac, 1980; Giggenbach, 1981; Frape and Fritz, 1987; Michard, 1987, 1991). Assuming thermodynamic equilibrium between a fluid and neogenic phases, we can distinguish fluids that evolved from composition far from equilibrium to near equilibrium. This thermodynamic approach allows the system and the chemical composition o f the fluid to be predicted,
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P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
assuming reversibility of the reactions, mass conservation during the overall reaction, and that temperature, mobile ion content and mineralogical associations are known (Helgeson, 1968; Michard, 1987). The classical (i.e. thermodynamic) approach, however, does not provide much insight about the mechanism of mineral transformation. For this reason, kinetic studies were undertaken to elucidate these processes. Starting with the earlier work ofLasaga (1984) which provided the theoretical foundation of global geokinetics, the majority of the mineral dissolution rate data was described as a function of pH, temperature and solution composition at far from equilibrium conditions (Guy and Schott, 1989; Stumm and Wieland, 1990). Under such conditions, dissolution is more easily identified, but reaction rates and mechanisms may vary as the system evolves toward equilibrium (Busenberg and Clemency, 1976, 1986). However, crustal fluids realistically approach equilibrium more easily as temperature increases up to 520 K (Arnorsson et al., 1983; Michard, 1991); for this reason investigations were also performed under near-equilibrium conditions. Several individual minerals have already been investigated: potassium feldspar (Pauwels et al., 1989), calcite (Beck et al., 1992; Shiraki and Brantley, 1992; Zuddas and Mucci, 1994), kaolinite (Devidal et al., 1992), quartz (Crerar and Dove, 1990), laumontite (Savage et al., 1993), and plagioclase (Zuddas and Michard, 1993). The purpose of this work was to experimentally investigate the dynamics of neogenic phases generated during the near-equilibrium interaction between granite and fluids using Sr isotopic spiked fluids. In addition, the contribution of individual primary mineral phases of the granite to the dissolution pathway is discussed.
2. Experimental methods Experimental determinations of fluid-mineral reactions may be facilitated through the addition of multiple tracers that have anomalous isotopic compositions. During the experimental reactions the isotopic tracers added to the fluid and material of natural isotopic composition liberated from the granite mix are incorporated into neogenic minerals. By monitoring the isotopic composition in the fluid, this method may be successfully applied to systems at near-equilibrium conditions
where the reaction rates are slow and the chemical affinity is virtually zero.
2.1. Rock sample description A sample of porphyric granodiorite (K202) taken from a depth of 1800 m from the European Community experimental drill-hole (GRK1) of Soultz-sous-F6ret, Alsace, France, was used. Macroscopic and microscopic analyses showed that K-feldspar, quartz, oligoclase, biotite and hornblende are primary minerals, whereas sphene, metallic oxides, apatite and zircon are accessories. Veins in the rock are filled by phyllosilicates, carbonates, iron oxides and quartz. Perthitic K-feldspar, the most abundant phenocryst, contains small crystals of plagioclase, biotite and oxides. Oligoclase is compositionally zoned from Anz7 at the center to Anl4 at the rim. Oligoclase contains calcite and phyilosilicates in the core, whereas phyllosilicates, including mica, are present in the periphery. Small crystals of albite, probably deuteric, were also detected in the rock samples. Hornblende associated with biotite is sometimes transformed to chlorite. According to a petrographic study of the area by Genter (1989), two hydrothermal events affected this rock, a first pervasive alteration developed on the original minerals and a second around the fractures in the rock. The chemical composition of the bulk sample and the Sr isotopic ratios of some of its main mineral constituents are reported in Table 1. The internal isochron calculation gives an age of 325 + 10 Myr.
2.1. Solutions To determine the composition of solutions near equilibrium we considered the transformation of granite to an assemblage of quartz + kaolinite + low-temperature albite + adularia + laumontite + calcite at 453 K (Michard, 1987). For a given temperature and mineral association the equilibrium composition of the solution depends solely on the chloride content and, using a C I concentration of 0.005 mol k g - ~ (common value for geothermal fluids), we can calculate the composition of the experimental solutions using the thermodynamic data reported in Table 2. To prevent alteration of fluid composition through evolution or precipitation, solutions were prepared 1 hr
P. Zuddas et al. /Chemical Geology 121 (1995) 145-1.54 Table 1 Chemical and isotopic composition of the granite and of the main separate minerals Oxides (total rock)
(wt%)
Na20 MgO AI203 SiO2 P205 K20 CaO MnO TiO2 Fe203 ~to,~~
4.17 1.14 14.72 69.55 0.025 4.01 2.17 0.06 0.66 2.265
Total
98.77
Fire loss
0.07
147
One gram of powder and 20 ml of solution were introducted in PTFE(polytetrafluoroethylene) cells into steel autoclaves. Experiments were conducted at 453 K and vapour saturation pressure. Durations of the experiments were between 1 hr and 190 days. Thermal equilibrium was reached in ~ 15 min. The cell was weighed before and after the reaction to measure loss by evaporation. Leaching of F - from the P'ITE cell during the experiment was negligible ( < 10 -5 mol kg-l). At the end o f the reactions the autoclaves were rapidly cooled to 298 K in ~ I0 min. Solutions were removed using a syringe and filtered through 0.2-/xm MINISART® filters. Liquid samples were divided into three aliquots: one diluted ten times for silica, chloride and fluoride analyses, a second acidified with " S u p r a pur e ' ' (HNO3 Merck ®) for K, Na, Ca, Sr and AI analyses, and a third for Sr isotopic analyses.
Minerals
Sr (ppm)
Rb (ppm)
87Sr/8OSr
2.4. Analytical techniques
Biotite Plagioclase K-feldspar
7 854 698
837 122 363
2.665 0.7074 0.7122
Total rock
468
193
0.7111
pH was measured with an Ingold ® combination glass electrode immediately after sample extraction at 298 K. Three buffer solutions ( p H = 4 . 0 1 , 6.87 and 9.00 at 298 K) were used for electrode calibration. Carbonate alkalinity o f the solutions was measured by titration with HC1 using the Gran (1952) method. Cations were analysed by flame or flameless atomic absorption spectrometry (Hitachi ~ 180-70 and G B C ® 902). SiO2 and PO43- were measured colorimetrically and anions by ion chromatography. For isotopic measurement, Sr was isolated by standard ion-exchange methods (Birck, 1986). 84Sr/86Sr, 87Sr/86Sr and 86Sr/88Sr ratios were determined on a single-collector Thompson ® TSN 206 C thermal ionization mass spectrometer with precision of a single
before the start of the experiments using A1C13.6H20, CaC12, NaOH, KC1 and NaHCO3 analytical reagents. Silica was obtained by exchange of Na ÷, by H ÷ of Na2SiO3 • nNaOH ( Baker dilut It ~) with Bio-Rad ® A G X8 ion-exchange resin. Strontium was spiked with 84Sr (75 %) and added as a nitrate salt. Initial concentrations are reported in Table 3. 2.3. Interaction The granite sample was crushed and sieved to obtain powder beween 0.08 and 0.2 mm. It is known [see Berner (1981) for a review] that newly exposed surfaces of grains, or dissolving very fine grains that adhere to the larger grain as dust coatings, both have an increase o f the real w a t e r / r o c k ratio that could make influence on the dissolution pathway. W e previously demonstrated (Zuddas and Michard, 1993), however, that during long-term near-equilibrium interaction, the w a t e r / r o c k ratio does not influence the major-cation concentrations at steady state.
Table 2 Thermodynamicdata for the minerals used (Michard, 1983) Minerals
log K
Calcite Quartz Low-temperaturealbite Laumontite Kaolinite Adularia
- 0.39 - 2.55 - 14.38 -- 23.74 -- 42.66 - 15.74
Dissolution reactions are written using the following species: H4SiO4, Na ÷, K ÷ , Ca2+ , H ÷ and AI(OH)4-.
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P. Zuddas et al. / Chemical Geology 121 (1995) 145-154
Table 3 Evolution of experimental solution during time Time (days)
pH 298 K
Ca-"+ (raM)
Na + (mM)
K+ (raM)
Alk. (raM)
PO~ (/xM)
Si (mM)
AI Sr2+ ( p , M ) (/zM)
84Sr/86Sr 87Sr/86Sr 86Sr/SaSr
0 0.04 20 48 77 105 132 160 163 190
9.08
0.336 0.215 0.0588 0.100 0.086 0.0319 0.056 0.0283 0.020 0.0152
7.94
0.267
23.9
0.356 0.204 0.186 0.187 0.255 0.250 0.270 0.200
>0.1 >0.1 >0.1 > 0.1 > 0.1 >0.1 >0.1 > 0.1 >0.1 >0.1
2.80
7.52 7.58 8.00 7.75 8.13 8.04 7.56 7.07
2.50 2.00 1.78 2.08 1.78 1.47 1.46
3.14 2.30 2.26 2.46 2.25 2.10 2.15 1.69
48.5 42.7 19.5 12.4 17.6
17.98 1.45 0.277 0.304 0.437 0.420 0.3914 0.2671 0.261 0.320
9.01 9.66 9.34 9.4 9.36 8.34 9.01
1.23
19.0 18.9
0.292 1.25 0.193 0.192 0.300 0.188 0.178 0.178 0.201 0.151
0.481 0.696 0.7072 0.7077 0.7053 0.7049 0.7058 0.7076 0.7075 0.7062
0.2238 0.124 0.1201 0.1202 0.1206 0.1206 0.1205 0.1200 0.1200 0.1205
Alk. = alkalinity. isotopic ratio analysis being _ 0.01% (at 20-95% confidence level). The mass discrimination factor was corrected using double normalization of 84Sr/86Sr and 86Sr/SSSr ratios (Faure, 1986). The procedural blank is ~ 0.1 ng Sr and is negligible compared to the 100 ng Sr of each sample. The solid recovered from the vessel was dried and examined by scanning electron microscopy ( S E M ) . New solids were removed from the primary ones by ultrasonic cleaning, separated in a separatory funnel and analysed by X-ray power diffraction.
( 848r / 86Sr ) sp., = Xt ( 848r / 865r) rock + ( 1 - X t ) (845r/86Sr) sp.o
where sp.t indicates the isotopic ratio in solution at time t (the mixing proportion); and sp.0 indicates the ratio in the initial solution. The isotopic ratio of the rock, (84Sr/86Sr) rock, is constant and equal to 0.0566 (Faure, 1986). The mixing proportion, X,, can also be expressed with respect to 86Sr by the equation: X t = (86Srrock) / (868rrock + 86Srsp.o)
3. S r a s t r a c e r o f i n t e r a c t i o n : T h e o r e t i c a l considerations
Sr is a minor element in granitic rocks and is mainly located in feldspar, although accessory minerals like apatite and calcite also contain strontium. In this section we will show how Sr isotopic ratios can be used as tracers of mineral dissolution and what they can reveal about the dynamics of neogenic phase precipitation. By artificially enriching the experimental fluid in 845r, the least abundant natural isotope, the release of Sr during the dissolution of granite will necessarily result in a decrease of both 84Sr/86Sr and 84Sr/88Sr ratios in solution. Thus, the isotopic composition of the solution depends, at any time, on the relative proportions of the spiked solution and the contribution by dissolution of the mineral phases of the granite. The proportion, X,, with respect to 84Sr/86Sr may be expressed by:
( 1)
(2)
The 878r abundance in the primary minerals of the granite depends on its history and Rb concentration so that every mineral will have a different 875r/86Sr ratio. For instance, plagioclase is generally characterized by low 878r/86Sr ratios because it is poor in Rb, whereas biotite has a high 878r/86Sr as it is rich in Rb. The mixing parameter, X,, between the rock and solution can be defined with respect to 87Sr as: (878r/86Sr) sod
=
St(SVSr/86Sr)rock -1- ( 1 --X,) (87Sr/86Sr)sp.O
(3)
Rewriting Eq. l with respect to the mixing proportion
x,: X t -= [ (848r/86Sr)sp., - (848r/86Sr)sp.o] / [ (84Sr/86Sr)rock
- (84Sr/86Sr) sp.o]
(4)
P. Zuddas et al. / Chemical Geology 121 (1995) 145-154
and substituting Eq. 4 into Eq. 3 it is possible to calculate, at any time during the experiment, the contribution of the ( 875r/86Sr)rock end-member:
(875r/868r) rock = [ (84Sr/86Sr) rock
+ (88Sr/86Sr)sp.O] / [ 1 + (84Sr/86Sr) sp.t
-- (84Sr/S6Sr) sp.o] [ S7Sr/S6Sr] sp.t
+ (875r/86Sr)sp./-I- (88Sr/86Sr).~p.,
-- [ (84Sr/S6Sr) rock
X ( 1 - X,) Sff +
- (84Sr/86Sr)sp.Ol / [ (84Sr/S6Sr)sp.t - (S4Sr/86Sr) sp.o] [ 87Sr/S6Sr I sp.o
(5)
Eq. 5 permits identification of the principal mineral assemblage responsible for the isotopic signature of the experimental hydrothermal fluid. The amount of spiked strontium in solution can be calculated at any time during the experiments assuming: ( 1 ) the mass conservation law: Sr,2 + = 84Sr + 86Sr + 87Sr + 88Sr
(6a)
or: Sr 2+ = 86Sr( 1 + 845r/86 Sr + 87Sr/86Sr q- 88Sr/86Sr)
(6b)
and
86Sr = Sr,2 + / [ 1 + (84Sr/86Sr) + (87Sr/86Sr) + (88Sr/S6Sr) ]
(6c)
where Sr,2 ÷ is the mass of dissolved strontium. (2) the mixing proportion, X,, related to 868r coming from the initial solution:
86Srsp.t = 86Sr( 1 - X , )
(7)
Substituting Eq. 6b into Eq. 7:
86Srsp.t = [Sr~ + / [ 1 + (845r/86Sr)sp.t + (87Sr/86Sr)sp., + (88Sr/86Sr)sp.t] ( 1 - X t )
Substituting Eq. 9 into Eq. 8 we can calculated the amount of the strontium spike in solution at any time by the following equation:
Srsp., = [ 1 + (848r/86Sr)sp.o + (878r/86Sr)sp.O
-- (84Sr/86Sr)sp.o] / [ (84Sr/S6Sr)sp.t
+ (87Sr/S6Sr) sp.o
149
(8)
(3) Eq. 6c at time 0:
(10)
Neoformed phases will contain Sr in isotopic equilibrium with their parent solutions. As rock dissolution likely continues after the initial formation of neogenic phases, the Sr isotopic ratios of the solution will change and, in particular, the 84Sr/86Sr ratio should decrease. Thus each new generation of neogenic phases formed will have a different isotopic composition compared to that of the previous one. The evolution of a spiked solution, therefore, provides significant information about the dynamics of neogenic phase evolution. 4. Results and discussion
The duration of these investigations was up to 190 days, although one experiment was performed for only 1 hr. We observed that pH, K ÷ and Na ÷ concentrations do not vary significantly (Table 3) during the reaction period. In the first hour of the interaction, however, Sr 2÷ increased from 0.29 to 1.2 p.mol kg - l and the (SaSr/86Sr)solution ratio decreased from 17.98 to 1.45, reflecting the initial dissolution of the granite while the Ca 2+ and carbonate alkalinity depletions suggest the precipitation of carbonate minerals. Between 1 hr and 20 days, Ca 2 ÷ and alkalinity continue to decrease while a depletion in H4SiO4 and A I ( O H ) 4 suggests the precipitation of aluminosilicates. Microscopic examination of the solids confirmed the formation of neogenic phases, but their precise identification was not possible because of their crystal size and amount. Smectites were only tentatively detected by X-ray diffraction. The precarious distinction from the parent minor phases of the granite (minerals of the pervasive alteration of the sample) made a more precise characterization difficult.
+ (87Sr/86Sr)sp.o + (88Sr/86Sr)sp.o] 4.1. Mineral source o f Sr
86Srsp.O = Sr 2+ / [ 1 + (845r/86Sr)sp.o + (87Sr/86Sr)sp.o q- (888r/86Sr)sp.O]
(9)
In order to identify the mineral assemblage controlling the input of Sr to the solution, we followed the
P. Zuddas et al. /Chemical Geology 12l (1995) 145-154
150
evolution of Sr isotopic ratios. A linear correlation was found between the 87Sr/86Sr and the 84Sr/86Sr ratios (Fig. 1) according to the equation: ( 8 7 S r / 8 6 S r ) solution
= -- 0.0146(84Sr/S6Sr) solution +0.71 15
( 1 1)
This correlation is defined for the range of (84Sr/ 86Sr) soJutionratios between 17.98 ( initial solution ratio) and 0.0566 (rock ratio), and we can apply Eq. 5 to calculate the (878r/86Sr)rock ratio corresponding to the average Sr isotopic composition of the dissolving mineral assemblage. In the first hour of interaction we find: (87Sr/86Sr)rock = 0.715 _ 0.001
(12a)
whereas for the whole experimental period (from 20 to 190 days) we obtain: (87Sr/86Sr) rock = 0.7 107 + 0.0003
( 12b)
The existence of two different rock end-members suggests that two main stages can be identified during this experimental interaction: one at the hour scale and another at the scale of days. In the early stage, dissolution is dominated by an end-member mineral assemblage having a 87Sr/86Sr ratio higher than the average 878r/86Sr ratio of the granite. In this assemblage, biotite, the main mineral of the granite with high 878r/86Sr ratio, could contribute strongly. In the second stage, the (87Sr/86Sr)rock ratio of the dissolving minerals is 0.720
, 13io
0.8
0.715 -
l
0'7~"~h
CD
"C o.7oa-
~ ,~.
e-.~.,
lo
20
O9
0.700 o
r
0.690-0.0
I 0.3
~
I 0.6
J
I 0.9
~
I 1.2
~
I 1.5
1.8
(84Sr/86Sr) Fig. 1.87Sr/8¢'Sr ratio as a function of the 84Sr/86Sr ratio in solution (Bio=biotite; K F = potassic feldspar; TR=total rock; Pl= plagioclase ).
Table 4 Amount of initial 84Sr spike in solution calculated by Eq. 10 Time (days)
Spiked Sr in solution (%)
0 0.24 20 48 77 105 132 160 163 190
100.00 76.85 2.43 2.66 6.37 3.64 3.22 2.22 2.41 2.21
close to the average ratio of the rock, reflecting no preferential contribution by a specific mineral. The amount of spiked strontium in solution was calculated by Eq. 10 (Table 4). At the beginning of the interaction the 0.292 ~ of Sr 2+ were 100% spiked in 84Sr. After 1 hr of interaction 25% of the 84Sr from the spike was removed from the solution by incorporation within neogenic phases while between 20 and 190 days virtually all the spike is removed from the solution. The amount of Sr in solution resulting from dissolution of the rock cannot, unfortunately, be calculated because of the simultaneous precipitation of neogenic phase from the first hour of the interaction. The dissolved Sr content and the isotopic ratios depend on the dissolution of plagioclase, biotite, Kfeldspar, apatite, sphene and hornblende. To evaluate the relative mineral weathering rate we assume that reactions with K-feldspar, apatite, sphene and hornblende are minor whereas variations in Sr solution composition are dominated by plagioclase and biotite dissolution. This is an acceptable approximation given that K-feldspar weathers at a much slower rate than plagioclase (Lasaga, 1984), both apatite and sphene are volumetrically very minor (Genter, 1989), and PO 3 was not detected. Since the Sr content of plagioclase is generally 20 times higher than that of hornblende, the latter mineral would only contribute a maximum of 0.5% of the Sr coming from plagioclase (Wedepohl, 1977). Assuming that the variation in the isotopic composition of Sr in the fluids results from the dissolution of plagioclase and biotite, we estimate that < 1% of the Sr in the fluids is derived from biotite.
151
P. Zuddas et al. / Chemical Geology 121 (1995) 145-154
To estimate the relative alteration rate of biotite with respect to plagioclase we assume that the proportion of dissolved Sr 2+ coming from both biotite and plagioclase dissolution is related to the relative Sr abundance of Sr in the mineral sources by: 2+ 2+ Srp. /Srb. =R(Srtot.p./Srtot.b.)
(13)
where Sre+p. and SrZ+b. are the dissolved strontium level coming from plagioclase and biotite alteration, respectively; Srtot.p. and Srtot.b. are the amount of Sr (grams) in plagioclase and biotite, respectively; and R is the relative rate of alteration. According to Eq. 13, if the rate of plagioclase dissolution is equal to the rate of biotite dissolution, R is equal to 1; if the rate of plagioclase dissolution is higher then that of biotite, R is higher than 1. We found experimentally that in the first hour of the reaction, R < 1. Srtot.p./Srtot.b. is equal to the ratio between the amount in every participating phase of the rock in our experimental system: Srtot.p./Srtot.b.
= (content ppm mass abund. %)p. (content ppm mass abund. %)b.
(14)
Assuming a plagioclase/biotite mass abundance ratio of ~ 2 (Genter, 1989) we evaluate, by Eq. 14, the Srtot.p./Srtot.b. ratio to be equal to 250. R can then be estimated by the following mixing equation:
(87Sr/86Sr) rock = [ 1 / ( 1 + 250R) ] (875r/86Sr)b. + [ 1/ { 1 + ( 1/250R) } ] (87Sr/86Sr)p.
(15) then, R = 1/250[ (87Sr/86Sr)b. -- (875r/86Sr)rock] / [ (875r/86Sr)rock -- (875r/86Sr)p.]
(15a) We evaluated that in the first hour of our experiment, under near-equilibrium conditions, plagioclase dissolved at the same rate as biotite, while in the later stage (from 1 hr to 20 days) it dissolved 2 times faster. The result of our calculation is consistent with the value of R = 2 derived from dissolution rate measurements conducted far from equilibrium on plagioclase (Mast and Drever, 1987) and biotite (Acker and Bricker, 1992).
1.0
0.8-
E) 0.6-oO GO0. 4 _ O0 0.2--
0.0
I
I
I
I
40
ao
12o
16o
2oo
Time (days) Fig. 2.84Sr/86Sr ratio in the solution as a function of time (days).
4.2. Sr 2 ÷ steady state and neogenic phase evolution
The hydrothermal alteration began with dissolution of mineral phases that were unstable under the experimental conditions of temperature and pressure. Under our experimental near-equilibrium conditions, the dissolution step is very rapid and is followed by both dissolution and concomitant precipitation of neogenic phases as the solution reaches steady state. The Sr 2+ concentration reached steady state after 20 days of interaction. The (848r/86Sr)solution ratio decreases strongly in the first 20 days while in the following days only small variations were observed (Fig. 2). The low (Snsr/86Sr)so]ution ratio found in the experimental period between 20 and 190 days may correspond to the end of the rapid dissolution of granite as the main source of 86Sr. The increase of the (845r/86Sr)so]ution ratio from 0.277 to 0.437 observed between 20 and 77 days is clearly related to the dissolution of phases with a high 84Sr content. In our system, only neogenic phases precipitated between 0 and 20 days are enriched in 84Sr. Because they are formed in the initial spiked solution, their subsequent dissolution must be responsible for the increase of the (845r/86Sr)solution ratio. It is then followed by a significant decrease from 0.437 to 0.261, observed between days 77 and 160, and is interpreted as the precipitation of a second generation of neogenic phases. A schematic representation of the evolution of the Sr composition of the solids and solution during the experiment is presented in Fig. 3. In A, the initial con-
152
P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
dition, we define two isotopic reservoirs: the strontium in the granite (open circles) and in the solution (solid circles). The dissolution of the granite is shown in B by the release of Sr (open circles) to the solution. Neogenic minerals contain strontium originating from both the original solution and the granite (C). In the following stage (D), the isotopic composition of the solution reflects an exchange of Sr between the third reservoir (the neogenic minerals) and the solution. Most of the Sr is released from the rock to the solution during the phase represented by B, whereas during phase D the isotopic composition of the solution depends only on the isotopic re-equilibration between neogenic minerals and solution, and consequently the isochemical dissolution-precipitation reaction of neogenic minerals. In this later stage of interaction we should introduce a new reservoir: "the exchangeable one". It can be represented as the amount of precipitating Sr able to be released back to the solution by "exchange reactions" after the precipitation of neogenic phases buffering the solution. The evaluation of the amount of strontium in the "exchangeable reservoir" is important in the definition of the Sr budget. Investigations presently under way will hopefully help to resolve this problem.
A
.
B
o . :. 0
oo
o. o.oO
o O
o 0
lc
• 0
0 0
0
.
o
[o o
Strontium
•
Initial
.
.
.
.
.
o in t h e
Strontium
mineral
r:
Parent minerals
~!i!!!~ NeoformeO minerels
Fig. 3. Schematic representation of the mechanism of neogenic mineral differentiation from the isotopic point of view (see text for discussion).
5. Implications to natural fluids
Changes in fluid chemical composition and the isotopic ratios result mostly from reactions occurring at an early stage during dissolution. Lasaga (1984) estimated that the time required to completely dissolve a 1-mm crystal of muscovite (at 298 K and p H = 5 ) is 2.7 Myr, 0.52 Myr for K-feldspar, 8 kyr for albite, and 112 yr for anorthite. Natural plagioclases being solid solutions can dissolve even more rapidly (Brantley, 1992). McNutt et al. (1990) suggested that high Sr concentrations in fluids associated with Canadian basement rocks (in an anorthositic environment) originate from the dissolution of feldspars, primarily the plagioclases. Edmunds et al. (1984, 1987) and Kay and Darbyshire (1986) concluded that feldspars, and especially the plagioclases, are the main source of Sr in brines associated with the Carnmenellis granite from Cornwall, southwestern England, although Sr from fluorite and carbonates may also be significant. Alternatively, Fritz et al. (1987) and Clauer et al. (1989) have identified chlorite, an alteration product of biotite, as a main Sr donor. Our experimental data indicate that the contribution of biotite in determining the fluid isotopic composition is important in the initial stage of the interaction whereas as a result of "long-term" interaction in near-equilibrium systems fluids reach the same isotopic signature as the rock. We believe that when natural fluids have a different isotopic signature compared to the host rock, the assumption of a closed system is not correct for isotopic considerations even if it is valid for the chemistry of major and minor elements. The precipitation of neogenic phases is intimately connected to the evolution of the fluid compositions during the alteration of the granitic rocks. Grimaud et al. (1990) showed that the evolution of major cations in Fennoscandian Stripa waters cannot always be correctly represented by a classical equilibrium (saturation) model between water and neogenic minerals. In fact, the same Ca 2 ÷ concentration can be buffered, if the Pco2 partial pressure of the system is fixed by a carbonatic equilibrium, by laumontite or prehnite or Ca-smectite (Zuddas and Michard, 1993). Due to the successive precipitation and dissolution of neogenic phases, being quantitatively minor with respect to the bulk rock, steady-state fluid compositions do not
P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
always c o r r e s p o n d to a true t h e r m o d y n a m i c equilibrium state.
6. Conclusions In this study, we w e r e able to demonstrate that the isotopic d o p i n g t e c h n i q u e is suitable for the experimental investigation o f f l u i d - g r a n i t e interaction under near-equilibrium conditions. O w i n g to the different Sr isotopic c o m p o s i t i o n o f constituent minerals in granite, doping initial fluid with 84Sr a l l o w e d us to trace the initial dissolution o f individual minerals and detect the precipitation o f n e o g e n i c phases. Our double spiked e x p e r i m e n t (fluid and rock) shows that the initial stage o f the granite dissolution is characterized by the release o f Sr having a higher 87Sr/ 865r ratio than the a v e r a g e o f the rock. W e evaluated that biotite dissolves at the s a m e rate than plagioclase in the first stage, w h i l e in the later one plagioclase dissolves 2 times faster than biotite. Initial dissolution o f biotite and plagioclase and conc o m i t a n t precipitation o f n e o g e n i c phases capturing strontium ( c a l c i t e + c l a y - m i n e r a l s ) are f o l l o w e d by steady-state [Sr 2 ÷ ] concentration. Our experimental data a l l o w e d us to s h o w that this steady state is determ i n e d by s u c c e s s i v e dissolution and precipitation o f n e o g e n i c phases and does not correspond to the end of the reaction. W e suggest that the concentrations o f natural g e o t h e r m a l fluids are g o v e r n e d by a re-equilibration o f n e o g e n i c m i n e r a l s and only a strong opening of the system can c h a n g e the c h e m i c a l c o m p o s i t i o n of the fluids.
Acknowledgements The authors wish to thank D. G r i m a u d , B. Dupr6 and J.-L. B i r c k for helpful discussions during the various stages o f this work. W e wish to thank S. W h i t t a k e r and A. M u c c i for reading an earlier version of the manuscript. Financial support was p r o v i d e d by the E u r o p e a n E c o n o m i c C o m m u n i t y ( E E C ) through research grant to P.Z. as well as funds m a d e available through the " D y n a m i q u e et Bilan de la T e r r e ( D B T ) " p r o g r a m o f the C N R S . J.I. D r e v e r and R.H. M c N u t t p r o v i d e d helpful and e n c o u r a g i n g r e v i e w s .
153
References Acker, J.P. and Bricker, O.P., 1992. The influence of pH on biotite dissolution and alteration kinetics at low temperature. Geochim. Cosmochim. Acta, 56: 3073-3092. Arnorsson, S., Gunnlaugsson, E. and Svarvasson, H., 1983. The chemistry of geothermal waters in Iceland, II. Mineral equilibria and independent variables controlling water compositions. Geochim. Cosmochim. Acta, 47: 547-566. Beck, W.J., Berndt, M.E. and Seyfried, Jr., W.E., 1992. Application of isotopic doping techniques to evolution of reaction kinetics and fluid/mineral distribtion coefficients: An experimental study of calcite at elevated temperatures and pressures. Chem. Geol., 97: 125-144. Bemer, R.A., 1981. Kinetics of weathering and diagenesis. In: A.C. Lasaga and R.J. Kirkpatrick (Editors), Kinetics of Geochemical Processes. Mineral. Soc. Am., Rev. Mineral., 8:111-134. Birck, J.L., 1986. Precision K-Rb-Sr isotopic analysis: application to Rb-Sr chronology. Chem. Geol., 56: 73-83. Brantley, S.L., 1992. Kinetics of dissolution precipitation - - Experimental and field results. In: Y.K. Kharaka and A.S. Maest (Editors), Proceedings of the 7th International Symposium on WaterRock Interaction (WRI7), 1: 3~5. Busenberg, E. and Clemency, C.V., 1976. The dissolution kinetics of feldspars at 25°C and 1 atm. of COz partial pressure. Geochim. Cosmochim. Acta, 40: 4149. Busenberg, E. and Clemency, C.V., 1986. A comparative study of the dissolution and crystal growth kinetics of calcite and aragonite. In: F. Mumpton (Editor), Studies in Diagenesis. U.S. Geol. Surv., Bull., 1578: 139-168. Clauer, N., Frape, S.K. and Fritz, B., 1989. Calcite veins of the Stripa granite (Sweden) as records of the origin of the groundwaters and their interactions with granite body. Geochim. Cosmochim. Acta, 53: 1777-1781. Crerar, D.A. and Dove, P.M., 1990. Kinetics of quartz dissolution in electrolyte solutions using a hydrothermal mixed flow reactor. In: Y. Noack and D. Nahon (Editors), Geochemistry of the Earth's Surface and of Mineral Formation. Chem. Geol., 84: 301304 (special issue; extended abstract). Devidal, J.L., Dandurand, J.L. and Schott, J., 1992. Dissolution and precipitation kinetics of kaolinite as a function of chemical affinity (T= 150°C, pH=2 and 7.8). In: Y.K. Kharaka and A.S. Maest (Editors), Proceedings of the 7th International Symposium on Water-Rock Interaction (WRI7), 1: 93-96. Edmunds, W.M., Andrews, J.N., Burgess, W.G., Kay, R.L.F. and Lee, D., 1984. The evolution of saline and thermal groundwaters in the Carnmenellis granite. Mineral. Mag., 48: 407-424. Edmunds, W.M., Kay, R.L.F., Miles, D.L. and Cook, J.M., 1987. The origin of saline groundwaters in the Cammenellis granite, Cornwall (U.K.): Further evidence from minor and trace elements. In: P. Fritz and S.K. Frape (Editors), Saline Waters and Gases in Crystalline Rocks. Geol. Assoc. Can., Spec. Pap., 33: 127-143. Faure, G., 1986. Principles of Isotopic Geology. Wiley, New York, N.Y., 589 pp. Frape, S.K. and Fritz, P., 1987. Geochemical trends from groundwaters from the Canadian shield. In: P. Fritz and S.K. Frape
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(Editors), Saline Waters and Gases in Crystalline Rocks. Geol. Assoc. Can., Spec. Pap., 33: 19-38. Fritz, B., Clauer, N. and Kam, M., 1987. Strontium isotopic data and geochemical calculation as indicators for the origin of saline waters in crystalline rocks. In: P. Fritz and S.K. Frape (Editors), Saline Waters and Gases in Crystalline Rocks. Geol. Assoc. Can., Spec. Pap., 33: 121-126. Genter, A., 1989. GEothermie roches chaudes sEches - - Le granite de Soultz-sous-For~t (Bas-Rhin, France). Ph.D. Dissertation, Universit6 OrlEans, OrlEans, 201 pp. Giggenbach, W.F., 1981. Geothermal chemical equilibria. Geochim. Cosmochim. Acta, 45: 393-410. Giggenbach, W.F., 1984. Mass transfer in hydrothermal alteration systems - - A conceptual approach. Geochim. Cosmochim. Acta, 48:2693-2711. Gran, G., 1952. Determination of the equivalence point in potentiometric titrations, II. Analyst (London), 77:661~571. Grimaud, D., Beaucaire, C. and Michard, G., 1990. Modelling of the evolution of ground water in a granite system at low temperature: the Stripa ground waters, Sweden. Appl. Geochem., 5:515-525. Guy, C. and Schott, J., 1989. Multisite surface reaction versus transport control during the hydrolysis of a complex oxide. In: J. Schott and A.C. Lasaga (Editors), Kinetic Geochemistry. Chem. Geol., 78:181-204 (special issue). Helgeson, H.C., 1968. Evaluation of irreversible reactions in geochemical processes involving minerals and aqueous solutions, I. Thermodynamic relations. Geochim. Cosmochim. Acta, 32, 853-877. Kay, R.L.F. and Darbyshire, D.P.F., 1986. A strontium isotope study of groundwater-rock interaction in the Carnmenellis granite. 5th Int. Symp. on Water-Rock Interaction, pp. 329-332. Lasaga, A.C., 1984. Chemical kinetics of water-rock interactions. J. Geophys. Res., B89: 4009-4025. Mast, M.A. and Drever, J.l., 1987. The effect of oxalate on the dissolution rates of oligoclase and tremolite. Geochim. Cosmochim. Acta, 51: 2559-2568. McNutt, R.H., Frape, S.K., Fritz, P., Jones, M.G. and MacDonald, I.M., 1990. The (87Sr/86Sr) values of Canadian Shield brines and fracture minerals with applications to groundwater mixing, fracture history, and geochronology. Geochim. Cosmochim. Acta, 54: 205-215.
Michard, G., 1983. Recueil des donnEes thermodynamiques concernant les Equilibres eaux-minEraux dans les reservoirs hydrothermaux. Rapp. CEE (Comm. Eur. Econ.) t~nergie Bruxelles, Rapp. No. EUR 8590 FR. Michard, G., 1987. Controls of the chemical composition of geothermal waters. In: H.C. Helgeson (Editor), Chemical Transport in Metasomatic Processes. Reidel, Dordrecht, pp. 323-353. Michard, G., 1991. The physical chemistry of geothermal systems. In: F. D'Amore (Editor), Application of Geochemistry in Geothermal Reservoir Development. UNITAR/UNDP, New York, N.Y., pp. 197-214. Michard, G. and Fouillac, C., 1980. Contr61e de la composition chimique des eaux thermales sulfurEes sodiques du Sud de la France. In: Y. Tardy (Editor), GEochimie des interactions entre les eaux, les minEraux et les roches, l~lEments, Tarbes, pp. 147166. Pauwels, H., Zuddas, P. and Michard, G., 1989. Behaviour of trace elements during feldspar dissolution in near-equilibrium conditions: preliminary investigation. In: J. Schott and A.C. Lasaga (Editors), Kinetic Geochemistry. Chem. Geol., 7 8 : 2 5 5 - 2 6 7 ( special issue). Savage, D., Cave, M.R., Haigh, D., Milodowski, A.E. and Young, M.E., 1993. The reaction kinetics of laumontite under hydrothermal conditions. Ear. J. Mineral., 5: 523-535. Shiraki, R. and Brantley, S., 1992. Precipitation kinetics of calcite at elevated temperatures. In: Y.K. Kharaka and A.S. Maest (Editors), Proceedings of the 7th International Symposium on WaterRock Interaction (WRI7), 1:111-114. Stumm, W. and Wieland, E., 1990. Dissolution of oxide and silicate minerals: rates depend on surface speciation. In: W. Stumm (Editor), Aquatic Chemical Kinetics. Wiley, New York, N.Y., pp. 367400. Wedepohl, K.W., 1977. Handbook of Geochemistry. Springer, Berlin. Zuddas, P. and Michard, G., 1993. Experimental mineral-fluid interaction in the N ~ C a - ( S r ) - A 1 - S i system. Eur. J. Mineral., 5: 807-818. Zuddas, P. and Mucci, A., 1994. Kinetics of calcite precipitation from seawater, I. A classical kinetics description for strong electrolyte solutions. Geochim. Cosmochim. Acta, 58: 43534362.
ISOTOPE GEOSCIENCE
Chemical Geology 121 (1995) 145-154
ELSEVIER
Granite-fluid interaction at near-equilibrium conditions: Experimental and theoretical constraints from Sr contents and isotopic ratios 1 Pierpaolo Zuddas a, Franqois Seimbille b, Gil Michard a aLaboratoire de GEochimie des Eaux, UniversitE Paris 7-Denis Diderot and lnstitut de Physique du Globe de Paris 2, place Jussieu, F- 75251 Paris Cedex 05, France bLaboratoire de GJochimie et Cosmoehimie, Universit~ Paris 7-Denis Diderot and lnstitut de Physique du Globe de Paris 2, place Jussieu, F- 75251 Paris Cedex 05, France
Received 5 May 1994; revision accepted 6 December 1994
Abstract Fluid-granite interaction at near-equilibrium conditions was investigated by doping an experimental system with S4Sr. Reactions were performed at 453 K, saturated vapour pressure, and fluid/rock ratio of 20. The initial composition of the experimental fluid corresponded to the theoretical saturation with respect to kaolinite, quartz, laumontite, low-temperature albite, calcite and adularia. Sr was used to trace parent-mineral dissolution, neogenic phase precipitation, and to identify the reactions at Srz+ steady state. The interaction is described, initially, by the isotopic mixing of the solution (84Sr/S6Sr)solutionratio and the granite (87Sr/ 86Sr)rock ratio. However, in the rock end-member every main mineral has a different 87Sr/S6Sr ratio. The evolution of 87Sr/86Sr ratio in solution as a function of 84Sr/S6Sr showed that in the early stage of the interaction (where the dissolution of the rock is the dominant reaction) STSr-enriched phases, such as biotite, strongly influence the isotopic budget. After some time, the 87Sr/ 86Sr ratio of the fluid approaches the Sr isotopic ratio of the bulk rock, suggesting that fluid-granite isotopic equilibration is attained. When the dissolution of the granite is stopped and once the Sr2÷ concentration reached steady state, variations of the 84Sr/S6Sr ratio with time can be interpreted by an isochemical dissolution-precipitation of neogenic phases.
1. Introduction Continental w a t e r - r o c k interaction mainly involves the reaction between meteoric water and granitic rocks. Chemical studies of hot springs, direct measurements in geothermal drill-holes, as well as experimental and theoretical investigations indicate that dissolution of parent minerals in host rocks and subsequent precipitation o f neogenic minerals are the dominant geochem[CA] 1 IPGP contribution No. 1339. 0009-2541/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSD10009-254 1 (94)001 59-6
ical reactions o f hydrothermal processes (Arnorsson et al., 1983; Giggenbach, 1984). In the temperature range o f 283-523 K, and at low fluid/rock ratios ( w / r ) , fluids at equilibrium produce new isochemical mineral assemblages (Michard and Fouillac, 1980; Giggenbach, 1981; Frape and Fritz, 1987; Michard, 1987, 1991). Assuming thermodynamic equilibrium between a fluid and neogenic phases, we can distinguish fluids that evolved from composition far from equilibrium to near equilibrium. This thermodynamic approach allows the system and the chemical composition o f the fluid to be predicted,
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P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
assuming reversibility of the reactions, mass conservation during the overall reaction, and that temperature, mobile ion content and mineralogical associations are known (Helgeson, 1968; Michard, 1987). The classical (i.e. thermodynamic) approach, however, does not provide much insight about the mechanism of mineral transformation. For this reason, kinetic studies were undertaken to elucidate these processes. Starting with the earlier work ofLasaga (1984) which provided the theoretical foundation of global geokinetics, the majority of the mineral dissolution rate data was described as a function of pH, temperature and solution composition at far from equilibrium conditions (Guy and Schott, 1989; Stumm and Wieland, 1990). Under such conditions, dissolution is more easily identified, but reaction rates and mechanisms may vary as the system evolves toward equilibrium (Busenberg and Clemency, 1976, 1986). However, crustal fluids realistically approach equilibrium more easily as temperature increases up to 520 K (Arnorsson et al., 1983; Michard, 1991); for this reason investigations were also performed under near-equilibrium conditions. Several individual minerals have already been investigated: potassium feldspar (Pauwels et al., 1989), calcite (Beck et al., 1992; Shiraki and Brantley, 1992; Zuddas and Mucci, 1994), kaolinite (Devidal et al., 1992), quartz (Crerar and Dove, 1990), laumontite (Savage et al., 1993), and plagioclase (Zuddas and Michard, 1993). The purpose of this work was to experimentally investigate the dynamics of neogenic phases generated during the near-equilibrium interaction between granite and fluids using Sr isotopic spiked fluids. In addition, the contribution of individual primary mineral phases of the granite to the dissolution pathway is discussed.
2. Experimental methods Experimental determinations of fluid-mineral reactions may be facilitated through the addition of multiple tracers that have anomalous isotopic compositions. During the experimental reactions the isotopic tracers added to the fluid and material of natural isotopic composition liberated from the granite mix are incorporated into neogenic minerals. By monitoring the isotopic composition in the fluid, this method may be successfully applied to systems at near-equilibrium conditions
where the reaction rates are slow and the chemical affinity is virtually zero.
2.1. Rock sample description A sample of porphyric granodiorite (K202) taken from a depth of 1800 m from the European Community experimental drill-hole (GRK1) of Soultz-sous-F6ret, Alsace, France, was used. Macroscopic and microscopic analyses showed that K-feldspar, quartz, oligoclase, biotite and hornblende are primary minerals, whereas sphene, metallic oxides, apatite and zircon are accessories. Veins in the rock are filled by phyllosilicates, carbonates, iron oxides and quartz. Perthitic K-feldspar, the most abundant phenocryst, contains small crystals of plagioclase, biotite and oxides. Oligoclase is compositionally zoned from Anz7 at the center to Anl4 at the rim. Oligoclase contains calcite and phyilosilicates in the core, whereas phyllosilicates, including mica, are present in the periphery. Small crystals of albite, probably deuteric, were also detected in the rock samples. Hornblende associated with biotite is sometimes transformed to chlorite. According to a petrographic study of the area by Genter (1989), two hydrothermal events affected this rock, a first pervasive alteration developed on the original minerals and a second around the fractures in the rock. The chemical composition of the bulk sample and the Sr isotopic ratios of some of its main mineral constituents are reported in Table 1. The internal isochron calculation gives an age of 325 + 10 Myr.
2.1. Solutions To determine the composition of solutions near equilibrium we considered the transformation of granite to an assemblage of quartz + kaolinite + low-temperature albite + adularia + laumontite + calcite at 453 K (Michard, 1987). For a given temperature and mineral association the equilibrium composition of the solution depends solely on the chloride content and, using a C I concentration of 0.005 mol k g - ~ (common value for geothermal fluids), we can calculate the composition of the experimental solutions using the thermodynamic data reported in Table 2. To prevent alteration of fluid composition through evolution or precipitation, solutions were prepared 1 hr
P. Zuddas et al. /Chemical Geology 121 (1995) 145-1.54 Table 1 Chemical and isotopic composition of the granite and of the main separate minerals Oxides (total rock)
(wt%)
Na20 MgO AI203 SiO2 P205 K20 CaO MnO TiO2 Fe203 ~to,~~
4.17 1.14 14.72 69.55 0.025 4.01 2.17 0.06 0.66 2.265
Total
98.77
Fire loss
0.07
147
One gram of powder and 20 ml of solution were introducted in PTFE(polytetrafluoroethylene) cells into steel autoclaves. Experiments were conducted at 453 K and vapour saturation pressure. Durations of the experiments were between 1 hr and 190 days. Thermal equilibrium was reached in ~ 15 min. The cell was weighed before and after the reaction to measure loss by evaporation. Leaching of F - from the P'ITE cell during the experiment was negligible ( < 10 -5 mol kg-l). At the end o f the reactions the autoclaves were rapidly cooled to 298 K in ~ I0 min. Solutions were removed using a syringe and filtered through 0.2-/xm MINISART® filters. Liquid samples were divided into three aliquots: one diluted ten times for silica, chloride and fluoride analyses, a second acidified with " S u p r a pur e ' ' (HNO3 Merck ®) for K, Na, Ca, Sr and AI analyses, and a third for Sr isotopic analyses.
Minerals
Sr (ppm)
Rb (ppm)
87Sr/8OSr
2.4. Analytical techniques
Biotite Plagioclase K-feldspar
7 854 698
837 122 363
2.665 0.7074 0.7122
Total rock
468
193
0.7111
pH was measured with an Ingold ® combination glass electrode immediately after sample extraction at 298 K. Three buffer solutions ( p H = 4 . 0 1 , 6.87 and 9.00 at 298 K) were used for electrode calibration. Carbonate alkalinity o f the solutions was measured by titration with HC1 using the Gran (1952) method. Cations were analysed by flame or flameless atomic absorption spectrometry (Hitachi ~ 180-70 and G B C ® 902). SiO2 and PO43- were measured colorimetrically and anions by ion chromatography. For isotopic measurement, Sr was isolated by standard ion-exchange methods (Birck, 1986). 84Sr/86Sr, 87Sr/86Sr and 86Sr/88Sr ratios were determined on a single-collector Thompson ® TSN 206 C thermal ionization mass spectrometer with precision of a single
before the start of the experiments using A1C13.6H20, CaC12, NaOH, KC1 and NaHCO3 analytical reagents. Silica was obtained by exchange of Na ÷, by H ÷ of Na2SiO3 • nNaOH ( Baker dilut It ~) with Bio-Rad ® A G X8 ion-exchange resin. Strontium was spiked with 84Sr (75 %) and added as a nitrate salt. Initial concentrations are reported in Table 3. 2.3. Interaction The granite sample was crushed and sieved to obtain powder beween 0.08 and 0.2 mm. It is known [see Berner (1981) for a review] that newly exposed surfaces of grains, or dissolving very fine grains that adhere to the larger grain as dust coatings, both have an increase o f the real w a t e r / r o c k ratio that could make influence on the dissolution pathway. W e previously demonstrated (Zuddas and Michard, 1993), however, that during long-term near-equilibrium interaction, the w a t e r / r o c k ratio does not influence the major-cation concentrations at steady state.
Table 2 Thermodynamicdata for the minerals used (Michard, 1983) Minerals
log K
Calcite Quartz Low-temperaturealbite Laumontite Kaolinite Adularia
- 0.39 - 2.55 - 14.38 -- 23.74 -- 42.66 - 15.74
Dissolution reactions are written using the following species: H4SiO4, Na ÷, K ÷ , Ca2+ , H ÷ and AI(OH)4-.
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Table 3 Evolution of experimental solution during time Time (days)
pH 298 K
Ca-"+ (raM)
Na + (mM)
K+ (raM)
Alk. (raM)
PO~ (/xM)
Si (mM)
AI Sr2+ ( p , M ) (/zM)
84Sr/86Sr 87Sr/86Sr 86Sr/SaSr
0 0.04 20 48 77 105 132 160 163 190
9.08
0.336 0.215 0.0588 0.100 0.086 0.0319 0.056 0.0283 0.020 0.0152
7.94
0.267
23.9
0.356 0.204 0.186 0.187 0.255 0.250 0.270 0.200
>0.1 >0.1 >0.1 > 0.1 > 0.1 >0.1 >0.1 > 0.1 >0.1 >0.1
2.80
7.52 7.58 8.00 7.75 8.13 8.04 7.56 7.07
2.50 2.00 1.78 2.08 1.78 1.47 1.46
3.14 2.30 2.26 2.46 2.25 2.10 2.15 1.69
48.5 42.7 19.5 12.4 17.6
17.98 1.45 0.277 0.304 0.437 0.420 0.3914 0.2671 0.261 0.320
9.01 9.66 9.34 9.4 9.36 8.34 9.01
1.23
19.0 18.9
0.292 1.25 0.193 0.192 0.300 0.188 0.178 0.178 0.201 0.151
0.481 0.696 0.7072 0.7077 0.7053 0.7049 0.7058 0.7076 0.7075 0.7062
0.2238 0.124 0.1201 0.1202 0.1206 0.1206 0.1205 0.1200 0.1200 0.1205
Alk. = alkalinity. isotopic ratio analysis being _ 0.01% (at 20-95% confidence level). The mass discrimination factor was corrected using double normalization of 84Sr/86Sr and 86Sr/SSSr ratios (Faure, 1986). The procedural blank is ~ 0.1 ng Sr and is negligible compared to the 100 ng Sr of each sample. The solid recovered from the vessel was dried and examined by scanning electron microscopy ( S E M ) . New solids were removed from the primary ones by ultrasonic cleaning, separated in a separatory funnel and analysed by X-ray power diffraction.
( 848r / 86Sr ) sp., = Xt ( 848r / 865r) rock + ( 1 - X t ) (845r/86Sr) sp.o
where sp.t indicates the isotopic ratio in solution at time t (the mixing proportion); and sp.0 indicates the ratio in the initial solution. The isotopic ratio of the rock, (84Sr/86Sr) rock, is constant and equal to 0.0566 (Faure, 1986). The mixing proportion, X,, can also be expressed with respect to 86Sr by the equation: X t = (86Srrock) / (868rrock + 86Srsp.o)
3. S r a s t r a c e r o f i n t e r a c t i o n : T h e o r e t i c a l considerations
Sr is a minor element in granitic rocks and is mainly located in feldspar, although accessory minerals like apatite and calcite also contain strontium. In this section we will show how Sr isotopic ratios can be used as tracers of mineral dissolution and what they can reveal about the dynamics of neogenic phase precipitation. By artificially enriching the experimental fluid in 845r, the least abundant natural isotope, the release of Sr during the dissolution of granite will necessarily result in a decrease of both 84Sr/86Sr and 84Sr/88Sr ratios in solution. Thus, the isotopic composition of the solution depends, at any time, on the relative proportions of the spiked solution and the contribution by dissolution of the mineral phases of the granite. The proportion, X,, with respect to 84Sr/86Sr may be expressed by:
( 1)
(2)
The 878r abundance in the primary minerals of the granite depends on its history and Rb concentration so that every mineral will have a different 875r/86Sr ratio. For instance, plagioclase is generally characterized by low 878r/86Sr ratios because it is poor in Rb, whereas biotite has a high 878r/86Sr as it is rich in Rb. The mixing parameter, X,, between the rock and solution can be defined with respect to 87Sr as: (878r/86Sr) sod
=
St(SVSr/86Sr)rock -1- ( 1 --X,) (87Sr/86Sr)sp.O
(3)
Rewriting Eq. l with respect to the mixing proportion
x,: X t -= [ (848r/86Sr)sp., - (848r/86Sr)sp.o] / [ (84Sr/86Sr)rock
- (84Sr/86Sr) sp.o]
(4)
P. Zuddas et al. / Chemical Geology 121 (1995) 145-154
and substituting Eq. 4 into Eq. 3 it is possible to calculate, at any time during the experiment, the contribution of the ( 875r/86Sr)rock end-member:
(875r/868r) rock = [ (84Sr/86Sr) rock
+ (88Sr/86Sr)sp.O] / [ 1 + (84Sr/86Sr) sp.t
-- (84Sr/S6Sr) sp.o] [ S7Sr/S6Sr] sp.t
+ (875r/86Sr)sp./-I- (88Sr/86Sr).~p.,
-- [ (84Sr/S6Sr) rock
X ( 1 - X,) Sff +
- (84Sr/86Sr)sp.Ol / [ (84Sr/S6Sr)sp.t - (S4Sr/86Sr) sp.o] [ 87Sr/S6Sr I sp.o
(5)
Eq. 5 permits identification of the principal mineral assemblage responsible for the isotopic signature of the experimental hydrothermal fluid. The amount of spiked strontium in solution can be calculated at any time during the experiments assuming: ( 1 ) the mass conservation law: Sr,2 + = 84Sr + 86Sr + 87Sr + 88Sr
(6a)
or: Sr 2+ = 86Sr( 1 + 845r/86 Sr + 87Sr/86Sr q- 88Sr/86Sr)
(6b)
and
86Sr = Sr,2 + / [ 1 + (84Sr/86Sr) + (87Sr/86Sr) + (88Sr/S6Sr) ]
(6c)
where Sr,2 ÷ is the mass of dissolved strontium. (2) the mixing proportion, X,, related to 868r coming from the initial solution:
86Srsp.t = 86Sr( 1 - X , )
(7)
Substituting Eq. 6b into Eq. 7:
86Srsp.t = [Sr~ + / [ 1 + (845r/86Sr)sp.t + (87Sr/86Sr)sp., + (88Sr/86Sr)sp.t] ( 1 - X t )
Substituting Eq. 9 into Eq. 8 we can calculated the amount of the strontium spike in solution at any time by the following equation:
Srsp., = [ 1 + (848r/86Sr)sp.o + (878r/86Sr)sp.O
-- (84Sr/86Sr)sp.o] / [ (84Sr/S6Sr)sp.t
+ (87Sr/S6Sr) sp.o
149
(8)
(3) Eq. 6c at time 0:
(10)
Neoformed phases will contain Sr in isotopic equilibrium with their parent solutions. As rock dissolution likely continues after the initial formation of neogenic phases, the Sr isotopic ratios of the solution will change and, in particular, the 84Sr/86Sr ratio should decrease. Thus each new generation of neogenic phases formed will have a different isotopic composition compared to that of the previous one. The evolution of a spiked solution, therefore, provides significant information about the dynamics of neogenic phase evolution. 4. Results and discussion
The duration of these investigations was up to 190 days, although one experiment was performed for only 1 hr. We observed that pH, K ÷ and Na ÷ concentrations do not vary significantly (Table 3) during the reaction period. In the first hour of the interaction, however, Sr 2÷ increased from 0.29 to 1.2 p.mol kg - l and the (SaSr/86Sr)solution ratio decreased from 17.98 to 1.45, reflecting the initial dissolution of the granite while the Ca 2+ and carbonate alkalinity depletions suggest the precipitation of carbonate minerals. Between 1 hr and 20 days, Ca 2 ÷ and alkalinity continue to decrease while a depletion in H4SiO4 and A I ( O H ) 4 suggests the precipitation of aluminosilicates. Microscopic examination of the solids confirmed the formation of neogenic phases, but their precise identification was not possible because of their crystal size and amount. Smectites were only tentatively detected by X-ray diffraction. The precarious distinction from the parent minor phases of the granite (minerals of the pervasive alteration of the sample) made a more precise characterization difficult.
+ (87Sr/86Sr)sp.o + (88Sr/86Sr)sp.o] 4.1. Mineral source o f Sr
86Srsp.O = Sr 2+ / [ 1 + (845r/86Sr)sp.o + (87Sr/86Sr)sp.o q- (888r/86Sr)sp.O]
(9)
In order to identify the mineral assemblage controlling the input of Sr to the solution, we followed the
P. Zuddas et al. /Chemical Geology 12l (1995) 145-154
150
evolution of Sr isotopic ratios. A linear correlation was found between the 87Sr/86Sr and the 84Sr/86Sr ratios (Fig. 1) according to the equation: ( 8 7 S r / 8 6 S r ) solution
= -- 0.0146(84Sr/S6Sr) solution +0.71 15
( 1 1)
This correlation is defined for the range of (84Sr/ 86Sr) soJutionratios between 17.98 ( initial solution ratio) and 0.0566 (rock ratio), and we can apply Eq. 5 to calculate the (878r/86Sr)rock ratio corresponding to the average Sr isotopic composition of the dissolving mineral assemblage. In the first hour of interaction we find: (87Sr/86Sr)rock = 0.715 _ 0.001
(12a)
whereas for the whole experimental period (from 20 to 190 days) we obtain: (87Sr/86Sr) rock = 0.7 107 + 0.0003
( 12b)
The existence of two different rock end-members suggests that two main stages can be identified during this experimental interaction: one at the hour scale and another at the scale of days. In the early stage, dissolution is dominated by an end-member mineral assemblage having a 87Sr/86Sr ratio higher than the average 878r/86Sr ratio of the granite. In this assemblage, biotite, the main mineral of the granite with high 878r/86Sr ratio, could contribute strongly. In the second stage, the (87Sr/86Sr)rock ratio of the dissolving minerals is 0.720
, 13io
0.8
0.715 -
l
0'7~"~h
CD
"C o.7oa-
~ ,~.
e-.~.,
lo
20
O9
0.700 o
r
0.690-0.0
I 0.3
~
I 0.6
J
I 0.9
~
I 1.2
~
I 1.5
1.8
(84Sr/86Sr) Fig. 1.87Sr/8¢'Sr ratio as a function of the 84Sr/86Sr ratio in solution (Bio=biotite; K F = potassic feldspar; TR=total rock; Pl= plagioclase ).
Table 4 Amount of initial 84Sr spike in solution calculated by Eq. 10 Time (days)
Spiked Sr in solution (%)
0 0.24 20 48 77 105 132 160 163 190
100.00 76.85 2.43 2.66 6.37 3.64 3.22 2.22 2.41 2.21
close to the average ratio of the rock, reflecting no preferential contribution by a specific mineral. The amount of spiked strontium in solution was calculated by Eq. 10 (Table 4). At the beginning of the interaction the 0.292 ~ of Sr 2+ were 100% spiked in 84Sr. After 1 hr of interaction 25% of the 84Sr from the spike was removed from the solution by incorporation within neogenic phases while between 20 and 190 days virtually all the spike is removed from the solution. The amount of Sr in solution resulting from dissolution of the rock cannot, unfortunately, be calculated because of the simultaneous precipitation of neogenic phase from the first hour of the interaction. The dissolved Sr content and the isotopic ratios depend on the dissolution of plagioclase, biotite, Kfeldspar, apatite, sphene and hornblende. To evaluate the relative mineral weathering rate we assume that reactions with K-feldspar, apatite, sphene and hornblende are minor whereas variations in Sr solution composition are dominated by plagioclase and biotite dissolution. This is an acceptable approximation given that K-feldspar weathers at a much slower rate than plagioclase (Lasaga, 1984), both apatite and sphene are volumetrically very minor (Genter, 1989), and PO 3 was not detected. Since the Sr content of plagioclase is generally 20 times higher than that of hornblende, the latter mineral would only contribute a maximum of 0.5% of the Sr coming from plagioclase (Wedepohl, 1977). Assuming that the variation in the isotopic composition of Sr in the fluids results from the dissolution of plagioclase and biotite, we estimate that < 1% of the Sr in the fluids is derived from biotite.
151
P. Zuddas et al. / Chemical Geology 121 (1995) 145-154
To estimate the relative alteration rate of biotite with respect to plagioclase we assume that the proportion of dissolved Sr 2+ coming from both biotite and plagioclase dissolution is related to the relative Sr abundance of Sr in the mineral sources by: 2+ 2+ Srp. /Srb. =R(Srtot.p./Srtot.b.)
(13)
where Sre+p. and SrZ+b. are the dissolved strontium level coming from plagioclase and biotite alteration, respectively; Srtot.p. and Srtot.b. are the amount of Sr (grams) in plagioclase and biotite, respectively; and R is the relative rate of alteration. According to Eq. 13, if the rate of plagioclase dissolution is equal to the rate of biotite dissolution, R is equal to 1; if the rate of plagioclase dissolution is higher then that of biotite, R is higher than 1. We found experimentally that in the first hour of the reaction, R < 1. Srtot.p./Srtot.b. is equal to the ratio between the amount in every participating phase of the rock in our experimental system: Srtot.p./Srtot.b.
= (content ppm mass abund. %)p. (content ppm mass abund. %)b.
(14)
Assuming a plagioclase/biotite mass abundance ratio of ~ 2 (Genter, 1989) we evaluate, by Eq. 14, the Srtot.p./Srtot.b. ratio to be equal to 250. R can then be estimated by the following mixing equation:
(87Sr/86Sr) rock = [ 1 / ( 1 + 250R) ] (875r/86Sr)b. + [ 1/ { 1 + ( 1/250R) } ] (87Sr/86Sr)p.
(15) then, R = 1/250[ (87Sr/86Sr)b. -- (875r/86Sr)rock] / [ (875r/86Sr)rock -- (875r/86Sr)p.]
(15a) We evaluated that in the first hour of our experiment, under near-equilibrium conditions, plagioclase dissolved at the same rate as biotite, while in the later stage (from 1 hr to 20 days) it dissolved 2 times faster. The result of our calculation is consistent with the value of R = 2 derived from dissolution rate measurements conducted far from equilibrium on plagioclase (Mast and Drever, 1987) and biotite (Acker and Bricker, 1992).
1.0
0.8-
E) 0.6-oO GO0. 4 _ O0 0.2--
0.0
I
I
I
I
40
ao
12o
16o
2oo
Time (days) Fig. 2.84Sr/86Sr ratio in the solution as a function of time (days).
4.2. Sr 2 ÷ steady state and neogenic phase evolution
The hydrothermal alteration began with dissolution of mineral phases that were unstable under the experimental conditions of temperature and pressure. Under our experimental near-equilibrium conditions, the dissolution step is very rapid and is followed by both dissolution and concomitant precipitation of neogenic phases as the solution reaches steady state. The Sr 2+ concentration reached steady state after 20 days of interaction. The (848r/86Sr)solution ratio decreases strongly in the first 20 days while in the following days only small variations were observed (Fig. 2). The low (Snsr/86Sr)so]ution ratio found in the experimental period between 20 and 190 days may correspond to the end of the rapid dissolution of granite as the main source of 86Sr. The increase of the (845r/86Sr)so]ution ratio from 0.277 to 0.437 observed between 20 and 77 days is clearly related to the dissolution of phases with a high 84Sr content. In our system, only neogenic phases precipitated between 0 and 20 days are enriched in 84Sr. Because they are formed in the initial spiked solution, their subsequent dissolution must be responsible for the increase of the (845r/86Sr)solution ratio. It is then followed by a significant decrease from 0.437 to 0.261, observed between days 77 and 160, and is interpreted as the precipitation of a second generation of neogenic phases. A schematic representation of the evolution of the Sr composition of the solids and solution during the experiment is presented in Fig. 3. In A, the initial con-
152
P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
dition, we define two isotopic reservoirs: the strontium in the granite (open circles) and in the solution (solid circles). The dissolution of the granite is shown in B by the release of Sr (open circles) to the solution. Neogenic minerals contain strontium originating from both the original solution and the granite (C). In the following stage (D), the isotopic composition of the solution reflects an exchange of Sr between the third reservoir (the neogenic minerals) and the solution. Most of the Sr is released from the rock to the solution during the phase represented by B, whereas during phase D the isotopic composition of the solution depends only on the isotopic re-equilibration between neogenic minerals and solution, and consequently the isochemical dissolution-precipitation reaction of neogenic minerals. In this later stage of interaction we should introduce a new reservoir: "the exchangeable one". It can be represented as the amount of precipitating Sr able to be released back to the solution by "exchange reactions" after the precipitation of neogenic phases buffering the solution. The evaluation of the amount of strontium in the "exchangeable reservoir" is important in the definition of the Sr budget. Investigations presently under way will hopefully help to resolve this problem.
A
.
B
o . :. 0
oo
o. o.oO
o O
o 0
lc
• 0
0 0
0
.
o
[o o
Strontium
•
Initial
.
.
.
.
.
o in t h e
Strontium
mineral
r:
Parent minerals
~!i!!!~ NeoformeO minerels
Fig. 3. Schematic representation of the mechanism of neogenic mineral differentiation from the isotopic point of view (see text for discussion).
5. Implications to natural fluids
Changes in fluid chemical composition and the isotopic ratios result mostly from reactions occurring at an early stage during dissolution. Lasaga (1984) estimated that the time required to completely dissolve a 1-mm crystal of muscovite (at 298 K and p H = 5 ) is 2.7 Myr, 0.52 Myr for K-feldspar, 8 kyr for albite, and 112 yr for anorthite. Natural plagioclases being solid solutions can dissolve even more rapidly (Brantley, 1992). McNutt et al. (1990) suggested that high Sr concentrations in fluids associated with Canadian basement rocks (in an anorthositic environment) originate from the dissolution of feldspars, primarily the plagioclases. Edmunds et al. (1984, 1987) and Kay and Darbyshire (1986) concluded that feldspars, and especially the plagioclases, are the main source of Sr in brines associated with the Carnmenellis granite from Cornwall, southwestern England, although Sr from fluorite and carbonates may also be significant. Alternatively, Fritz et al. (1987) and Clauer et al. (1989) have identified chlorite, an alteration product of biotite, as a main Sr donor. Our experimental data indicate that the contribution of biotite in determining the fluid isotopic composition is important in the initial stage of the interaction whereas as a result of "long-term" interaction in near-equilibrium systems fluids reach the same isotopic signature as the rock. We believe that when natural fluids have a different isotopic signature compared to the host rock, the assumption of a closed system is not correct for isotopic considerations even if it is valid for the chemistry of major and minor elements. The precipitation of neogenic phases is intimately connected to the evolution of the fluid compositions during the alteration of the granitic rocks. Grimaud et al. (1990) showed that the evolution of major cations in Fennoscandian Stripa waters cannot always be correctly represented by a classical equilibrium (saturation) model between water and neogenic minerals. In fact, the same Ca 2 ÷ concentration can be buffered, if the Pco2 partial pressure of the system is fixed by a carbonatic equilibrium, by laumontite or prehnite or Ca-smectite (Zuddas and Michard, 1993). Due to the successive precipitation and dissolution of neogenic phases, being quantitatively minor with respect to the bulk rock, steady-state fluid compositions do not
P. Zuddas et al. /Chemical Geology 121 (1995) 145-154
always c o r r e s p o n d to a true t h e r m o d y n a m i c equilibrium state.
6. Conclusions In this study, we w e r e able to demonstrate that the isotopic d o p i n g t e c h n i q u e is suitable for the experimental investigation o f f l u i d - g r a n i t e interaction under near-equilibrium conditions. O w i n g to the different Sr isotopic c o m p o s i t i o n o f constituent minerals in granite, doping initial fluid with 84Sr a l l o w e d us to trace the initial dissolution o f individual minerals and detect the precipitation o f n e o g e n i c phases. Our double spiked e x p e r i m e n t (fluid and rock) shows that the initial stage o f the granite dissolution is characterized by the release o f Sr having a higher 87Sr/ 865r ratio than the a v e r a g e o f the rock. W e evaluated that biotite dissolves at the s a m e rate than plagioclase in the first stage, w h i l e in the later one plagioclase dissolves 2 times faster than biotite. Initial dissolution o f biotite and plagioclase and conc o m i t a n t precipitation o f n e o g e n i c phases capturing strontium ( c a l c i t e + c l a y - m i n e r a l s ) are f o l l o w e d by steady-state [Sr 2 ÷ ] concentration. Our experimental data a l l o w e d us to s h o w that this steady state is determ i n e d by s u c c e s s i v e dissolution and precipitation o f n e o g e n i c phases and does not correspond to the end of the reaction. W e suggest that the concentrations o f natural g e o t h e r m a l fluids are g o v e r n e d by a re-equilibration o f n e o g e n i c m i n e r a l s and only a strong opening of the system can c h a n g e the c h e m i c a l c o m p o s i t i o n of the fluids.
Acknowledgements The authors wish to thank D. G r i m a u d , B. Dupr6 and J.-L. B i r c k for helpful discussions during the various stages o f this work. W e wish to thank S. W h i t t a k e r and A. M u c c i for reading an earlier version of the manuscript. Financial support was p r o v i d e d by the E u r o p e a n E c o n o m i c C o m m u n i t y ( E E C ) through research grant to P.Z. as well as funds m a d e available through the " D y n a m i q u e et Bilan de la T e r r e ( D B T ) " p r o g r a m o f the C N R S . J.I. D r e v e r and R.H. M c N u t t p r o v i d e d helpful and e n c o u r a g i n g r e v i e w s .
153
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