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Journal of Petrology Volume 41 Number 9 Pages 1455-1466 2000
© Oxford University Press 2000
Two Disequilibrium Quartz–Feldspar 18O/16O Fractionations within the Aral Granite Batholith, Altay Mountains of China: Evidence for Occurrence of Two Stages of O and H Isotopic Exchange of a Heterogeneous Granite System with Aqueous Fluids
INSTITUTE OF GEOLOGY AND GEOPHYSICS, THE RESEARCH CENTRE OF MINERAL RESOURCES EXPLORATION, CHINESE ACADEMY OF SCIENCES, PO BOX 9701, A-11 DATUN ROAD, BEIJING 100101, P.R. CHINA
Received November 5, 1998; Revised typescript accepted February 9, 2000
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ABSTRACT |
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18O/16O and D/H isotope compositions are reported for coexisting quartz, feldspar and biotite from four lithological units within the Aral granite batholith, Altay Mts of China. The data exhibit a reversed and an anomalously large positive quartz–feldspar 18O/16O fractionation. These two 18O/16O fractionations, together with a marked decrease in the
D values of biotite, are interpreted in terms of a model involving two stages of isotopic exchange with aqueous fluids. The first stage of 18O/16O exchange with an 18O-rich aqueous fluid occurred during subsolidus cooling. Kinetic effects of the first stage of 18O/16O exchange are characterized by the reversed quartz–feldspar 18O/16O fractionation that was recorded in the megacrystic coarse-grained granites of Group I and the megacryst-bearing medium- to coarse-grained granites of Group II. Robust parameters for the first stage of exchange are obtained by modelling on the hypothesis of initial 18O/16O heterogeneity in the granite plus subsolidus 18O/16O exchange. The spread in the measured 18O values of quartz reflects the initial heterogeneity of the granite. Calculated isochrons illustrate that an initially heterogeneous system reacting with an externally buffered fluid can generate arrays that mimic isotherms. The second stage of isotopic exchange with 18O- and D-depleted meteoric water occurred after magma solidification, resulting in the anomalously large positive quartz–feldspar 18O/16O fractionation and a marked decrease in the D values of biotite. These kinetic effects have been recorded in the finer-grained granites of Groups III and IV. However, the coarser-grained granites of Groups I and II have essentially survived the second stage of 18O/16O exchange. KEY WORDS: Altay; granite; initial 18O/16O heterogeneity; pseudoisotherm; quartz–feldspar 18O/16O reversal
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY, PETROLOGY AND...SAMPLE DESCRIPTION, EXPERIMENTAL...DISCUSSIONCONCLUSIONREFERENCES |
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Research on the oxygen isotope exchange kinetics between coexisting minerals and aqueous fluids has been performed on a variety of rock-types (Gregory, 1986
; Gregory & Taylor, 1986a, 1986b; Criss et al., 1987; Gregory et al., 1989; Richards et al., 1996). Kinetic theory has been applied to the oxygen isotope systematics of quartz–feldspar pairs in granites, plagioclase–pyroxene pairs in gabbros and carbonate–quartz pairs in marbles, and to the oxygen isotope systematics of Precambrian siliceous iron formations and low-grade metamorphic terrains (Gray et al., 1991).
Disequilibrium 18O/16O fractionations are characteristically large and positive; for example, between clinopyroxene and plagioclase in layered gabbros, between quartz and feldspar in granites, and between clinopyroxene and olivine in mantle nodules (Gregory & Taylor, 1981, 1986a; Gregory et al., 1981, 1989; Criss & Taylor, 1983; Taylor, 1990; Liu et al., 1999), or between quartz and arfvedsonite in alkali granites (Liu et al., 1999). In these cases, the data in plots of 18O vs 18O exhibit steep arrays that intersect the equilibrium fractionation isotherms. This indicates that these mineral pairs are out of oxygen isotope equilibrium resulting from 18O/16O exchange with an aqueous fluid and, hence, reveals the fossil hydrothermal system.
Another type of disequilibrium fractionation between coexisting minerals, however, is characterized by the 18O/16O reversal. In this situation, the data in a 18O vs 18O diagram form gently sloping or flattened arrays that are approximately parallel to the isotherms and considered to define a ‘pseudoisotherm’ (Richards et al., 1996). Kinetic calculations on the assumption of a homogeneous 18O/16O composition tend to give unrealistically large values of dimensionless time (kt), i.e. the time-duration t of 18O/16O exchange normalized by the exchange rate constant, and the water-to-rock ratio W/R as required by the spread in the measured 18O values of quartz and the flattened arrays.
The mechanisms of oxygen isotope exchange have provided an important aspect of kinetic research. Clouding and turbidity are common phenomena in feldspars from plutonic rocks (e.g. Taylor, 1990) and have been attributed to the presence of myriads of micro-fluid inclusions (Worden et al., 1990). Recent scanning electron microscopy (SEM) and transmission electron microscopy (TEM) studies have shown that exsolution and microtextural reorganization of alkali feldspar are mutually facilitated with feldspar–water interaction (Parsons, 1978; Worden et al., 1990; Witt-Eickschen et al., 1996; Martin et al., 1997). These studies show that solution–reprecipitation is a much more effective mechanism than diffusion in the microtextural reorganization of alkali feldspar and presumably the accompanying 18O/16O exchange with an aqueous fluid. On the basis of research on oxygen isotope systems of granites, Barnett & Bowman (1995) have demonstrated that differences in the oxygen isotopic exchange rates between feldspar and quartz can be as large as orders of magnitude and result from exchange via diffusion in quartz and via surface reaction in feldspar.
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REGIONAL GEOLOGY, PETROLOGY AND PETROGRAPHY |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY, PETROLOGY AND...SAMPLE DESCRIPTION, EXPERIMENTAL...DISCUSSIONCONCLUSIONREFERENCES |
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As shown in Fig. 1, the Altay mountain region in the northern Xinjiang Uygur Autonomous Region, China, has been attributed to a Palaeozoic orogeny that occurred on the southwestern continental margin of the Siberian platform (Zou et al., 1988
; Liu, 1990, 1993). During the Late Silurian to Early Devonian the region was folded, and large-scale autochthonous–semiautochthonous gneissic biotite granites were produced concurrently with regional metamorphism. The metamorphic mineral assemblages garnet + andalusite indicate that the metamorphism is of the low-pressure–high-temperature type (Qu & He, 1992).
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43, late Hercynian granite. The circle and the nearby number show the sampling location and sample number with the prefix XG omitted. The continuous line and dotted line stand for, respectively, the geological boundary and facies interface. The bold continuous line represents a fault (F5).During the late Carboniferous, thrusting occurred in the Altay orogen and stacked the crust in the region (Qu & He, 1992
). The thrust–décollement system F1, F2, F3, F4 and F5 (Fig. 1) induced partial melting that led to the production of granites. Uplift of the Altay Mts region occurred during the Permian; the megacrystic biotite granites and two-mica granites were emplaced concurrently.
Figure 2 shows the Aral granite pluton, a composite S-type batholith with a surface area >2000 km2 (Liu et al., 1997). The batholith was intruded into the schist, gneiss and phyllite of the Middle to Upper Ordovician Habahe group and the Lower Carboniferous Hongshanzui group. The batholith is dated at 250·9 Ma by the whole-rock Rb–Sr isochron method (Liu, 1990), with an initial 87Sr/86Sr ratio of 0·7085 ± 0·0005. U–Pb dating of zircon gave a discordant age of 279·4 Ma. On the basis of these data the Aral batholith is considered to be late- or post-orogenic emplaced during the late Hercynian.
As shown in Fig. 2, four lithological units, which are termed Groups I, II, III and IV in this paper, have been recognized within the eastern half of the Aral batholith. Group I, which occurs in the southern part and the northern part of the batholith, consists of megacrystic coarse-grained biotite granite. The abundance of megacrysts in Group I varies from 30 to 60% by volume, whereas the groundmass minerals are generally coarse grained. Group II, which occurs in the eastern part and the central part of the batholith, consists of a megacryst-bearing medium- to coarse-grained biotite granite. The lithological difference between Groups I and II lies in the texture. Megacrysts become less abundant (<30%) and sizes of groundmass minerals are generally decreased in Group II compared with Group I. The megacrysts, which are dominantly composed of microcline, are rectangular shaped with grain sizes ranging from 1·5 cm x 3 cm to 4 cm x 7 cm. The elongation of the megacrysts is aligned at 
320–335°, with the trend of the regional structure. Group III occurs in a narrow zone along the eastern and the northern margins of the batholith and consists of a fine- to medium-grained biotite granite. Group IV, which occurs in the central part of the batholith, comprises a fine-grained two-mica alkali–feldspar granite. Groups I, II and III were intruded concurrently, whereas Group IV is considered to be slightly younger.
In Groups I and II the granites comprise mainly K-feldspar (44–57%), quartz (30–35%), plagioclase (10–20%), biotite (1–6%) and minor muscovite. Accessory minerals are chiefly magnetite, ilmenite and apatite, with minor zircon and sphene. In Group III the granites are composed of K-feldspar (30–50%), quartz (37–45%), plagioclase (10–15%), biotite (3–10%) and some opaque minerals. In Group IV the granites are composed of K-feldspar (50–57%), quartz (30–34%), plagioclase (5–10%), muscovite (2–10%) and biotite (3–5%). The K-feldspar is microcline and perthite; the plagioclase is rich in sodium.
The Aral batholith has not been significantly weathered. Biotite is relatively unaltered, but locally shows weak chloritization. Feldspars, especially plagioclase, show clouding or turbidity in Group II and the northern part of Group I. Plagioclase is weakly sericitized with locally developed zoning.
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SAMPLE DESCRIPTION, EXPERIMENTAL PROCEDURES AND ANALYTICAL RESULTS |
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Most of the samples were collected along newly built roads or in quarries. A total of 16 unweathered samples, each weighing >1 kg, were collected from the four petrographic units. Sampling locations are shown in Fig. 2. Samples collected from the Group I, II, III and IV lithological units were XG17–XG23; XG5, XG7 and XG8; XG2 and XG3; and XG9–XG12, respectively.
Rock samples were crushed into grains with sizes ranging between 0·5 and 0·2 mm. Quartz, feldspar and biotite were separated using the standard separation procedure of panning, flotation, magnetic dressing, separation by heavy liquid, corrosion in hot dilute HCl solution and hand-picking under the binocular microscope (Liu et al., 1996). Quartz grains with feldspar inclusions or intergrowths were further crushed into grains of 0·2–0·1 mm or finer sizes and then were repeatedly separated. Fine feldspar intergrowths and iron stains on the margins of quartz grains were dissolved in hot dilute HCl. The final purity of mineral separates exceeded 99%, as checked by X-ray diffraction (XRD).
Determination of the oxygen and hydrogen isotope compositions of quartz, feldspar and biotite was performed in the Isotope Laboratory, Institute of Mineral Ore Deposit Geology, Geological Academy of China. Oxygen and hydrogen were removed from the mineral separates by standard BrF5 fluorination (e.g. Taylor & Epstein, 1962) and zinc reduction techniques (e.g. Qian, 1986; Guo & Qian, 1997). Isotope analyses were performed on a MAT-251 ratio mass spectrometer. Both 18O/16O and D/H ratios are expressed by the conventional 18O and D value in permil relative to V-SMOW. Reproducibilities of 18O and D values are ±0·2
and ±1, respectively. The analytical results are shown in Table 1.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY, PETROLOGY AND...SAMPLE DESCRIPTION, EXPERIMENTAL...DISCUSSIONCONCLUSIONREFERENCES |
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Two stages of isotopic exchange
Figure 3a–c shows three oxygen isotope domains defined by the lithological Groups I and II, Group III and Group IV. Overall, these domains are out of the range of magmatic isotopic compositions as defined by the isotherms; this indicates inconsistent, unreasonable, temperatures. Strikingly, two kinds of disequilibrium quartz–feldspar fractionations are shown in Fig. 3 and Table 1: samples of Groups I and II exhibit negative or reversed
18OQtz–Fsp values that range from -0·69 to -2·51; whereas samples of Group IV except XG9, and XG3 from Group III exhibit anomalously large positive 18OQtz–Fsp values. It is noteworthy that XG9, and XG2 from Group III, exhibit similar, but less extreme, 18OQtz–Fsp fractionations to Groups I and II.
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and • represent the samples of Group I, II, III and IV, respectively. Each data point is specified by the sample number whose prefix XG is omitted. The lines stand for the isotherms, which are calculated according to the mineral–water oxygen isotope fractionation equations (O’Neil & Taylor, 1967
Figure 4a–c shows four oxygen and hydrogen isotope domains defined by Groups I, II, III and IV, respectively. As shown in Fig. 4a–c and in Table 1, D values of biotites from Group II are distinctly lower than those from Group I, which display the highest D values (–57 to -77). The 18O values of feldspars and biotites from Group III are distinctly lower than those from Group I. Biotites from Group IV exhibit the lowest D values, varying between -88 and -99. These shifts are represented by arrows in Fig. 4a–c. It is interesting that, although D values of biotites of Group II exhibit a marked decrease from those of Group I, 18O values of biotites and feldspars of Group II have not been shifted from those of Group I. This discoupling between the D/H and 18O/16O shifts of silicate minerals in granite depends on the molar fractions of hydrogen and oxygen in silicate minerals. The molar fraction of hydrogen in silicate minerals is so small that only a relatively small cumulative flux of meteoric water or W/R ratio is required to considerably shift their D/H isotopic composition. However, a much larger cumulative flux of meteoric water or W/R ratio would be expected to reset the 18O/16O isotopic compositions of silicate minerals.
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The two contrasting disequilibrium quartz–feldspar 18O/16O fractionations, together with the oxygen and hydrogen isotope shifts shown in Fig. 4a–c, suggest that two stages of isotopic exchange with aqueous fluids have occurred. The less extreme
18OQtz–Fsp reversals of samples XG2 and XG9 imply that Groups III and IV also underwent the isotopic exchange that caused the 18OQtz–Fsp reversals of Groups I and II before the extensive overprinting of the isotopic exchange that shifted the 18OQtz–Fsp reversals to smaller values or resulted in the anomalously large positive 18OQtz–Fsp fractionations. Also, the temperatures of the two stages of isotopic exchange provide constraint on their relative sequence. The reversed 18OQtz–Fsp fractionations are so remarkable that it is reasonable to infer that the responsible isotopic exchange must have extensively influenced the megacrystic coarse-grained or medium- to coarse-grained granites of Groups I and II. Consequently, a relatively high-temperature condition during subsolidus cooling is required to effectively drive the exchange into the interior of megacrysts. However, another stage of isotopic exchange that brought about the anomalously large positive 18OQtz–Fsp fractionations influenced only the finer-grained granites of Groups III and IV. Thus, it is plausible that this stage of isotopic exchange occurred at low temperatures post-dating magma solidification. On the basis of the above arguments, the relative sequence of occurrence of these two stages of exchange can be established. Samples of Groups I and II provide evidence for the first stage of 18O/16O exchange with an aqueous fluid, recording marked quartz–feldspar 18O/16O reversals. However, they have essentially survived, or have been only slightly overprinted, in terms of their D/H isotopic compositions, by a second stage of isotopic exchange with 18O- and D-depleted meteoric water, owing to their coarser grain sizes and the lower temperatures of exchange. Samples of Groups III and IV largely reflect the second stage of isotopic exchange, which resulted in the anomalously large positive quartz–feldspar 18O/16O fractionations and the isotopic shifts as shown by the arrows in Fig. 4a–c.
It could be inferred that quartz was inert to the second stage of exchange because the 18O values of quartz from all four groups, although showing a large spread, coincide or overlap.
The grain sizes of minerals become smaller in sequence from Group I to Group II to Group IV; correspondingly, the extent of D/H and 18O/16O exchange between biotite and feldspar and meteoric water becomes greater from Group I to Group II to Group IV. For mineral–fluid systems, the change in the oxygen isotopic composition of a mineral is directly proportional to the surface area of the mineral (Cole et al., 1983) and, therefore, is in reverse proportion to the effective grain size.
On the external aqueous fluid of the first stage of exchange
Samples from Groups I and II have essentially preserved the evidence for the first stage of isotopic exchange with an aqueous fluid. On the basis of the observed effects, the 18O/16O exchange kinetics of the first stage can be explored. In Fig. 3a the array formed by the samples of Groups I and II forms a weak positive correlation. Therefore, 18O/16O exchange must have occurred under open-system conditions in the presence of an external aqueous fluid.
The following equations (Criss et al., 1987) describe the kinetic effects of oxygen isotope exchange between coexisting minerals and an external aqueous fluid in a single-pass open system:
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iw to mineral–water isotope fractionation factors; kiw to the mineral–water isotopic exchange rate constants; xi and xw to oxygen mole fractions; Ri and Rw to the 18O/16O composition of the phase indicated by the subscript; u to the normalized infiltration rate in s-1; and t to time in seconds.
According to equation (1), the change in the 18O value of a mineral depends upon the exchange rate constant kiw and the term (iwRw – Ri). The latter represents the free energy difference that drives the exchange reaction (Gregory et al., 1989). When the external aqueous fluid is rich in 18O, the increase in the 18OFsp (Fsp:feldspar) value should be far greater than that in the 18OQtz (Qtz: quartz) value during the initial stages of exchange as kFsp–water is far greater than kQtz–water, so that 18OQtz–Fsp values decrease gradually until they became negative, i.e. reversed. The 18OQtz–Fsp values of the coexisting quartz and feldspar pairs from Groups I and II are all negative (Table 1), indicating that the external aqueous fluid is rich in 18O. An 18O-rich aqueous fluid could be released through dehydration reactions during regional metamorphism or 18O/16O exchange with country rocks. Thus, this 18O-rich aqueous fluid has acquired the nature of metamorphic waters. The 18O values of metamorphic waters range between +3 and +20 (Sheppard, 1986). The value of +10, which roughly represents the median value between +3 and +20, is assumed to be the 18O value of the 18O-rich aqueous fluid of the first stage of 18O/16O exchange that infiltrated the Aral batholith, for modelling purposes.
On the temperature of the first stage of exchange
For modelling purposes, 300°C was chosen as the temperature of the first stage of 18O/16O exchange, for the following reasons:
- the temperature of exchange should be slightly higher than 200°C, for otherwise intense chloritization and/or sericitization would be encountered (Taylor, 1990).
- The magma was initially rich in water because OH-bearing minerals, such as biotite and muscovite, are abundant in the Aral batholith. Definitive evidence for the influx of external water into the granitic pluton at temperatures higher than 400°C is lacking, and at these higher temperatures magmatic fluids appear to be dominant (Sheppard et al., 1971) and would form a barrier to the external fluid infiltration (Sheppard et al., 1971; Taylor & Forester, 1979). As a result, the exchange temperature should be lower than 400°C.
- Both feldspar and biotite are sensitive to 18O/16O exchange, and have similar exchange rates. With an increase in the cumulative flux of the external fluid, the values of 18OFsp and 18OBt (Bt: biotite) should tend to reflect equilibration near the isotherm of the exchange temperature. In Fig. 3c the array of data of Groups I and II has already overstepped the 400°C isotherm, and is approaching the 300°C isotherm.
Kinetic modelling of the first stage of exchange on the assumption of a single initial starting point
Modelling of the first stage of isotopic exchange was performed on the assumption of a single ‘initial starting point’, i.e. a homogeneous initial 18O/16O composition of the granitic magma. Although the exact initial 18O values of quartz cannot be defined, some broad constraints can be given. The upper limit of the initial 18O values of quartz should be near but lower than the measured 18O values of quartz, which vary from 11·17 to 13·60, whereas the lower limit value should not be too far from the measured 18O values of quartz because quartz is inert to the 18O/16O exchange and the measured 18O values of quartz must not have been considerably shifted from their initial values. Hence, initial 18O values of quartz are assumed ranging from 8 to 10. By the lower limit value of 8, a marked shift magnitude from 3 to >5 to the measured 18O values of quartz is implied. Hence, the above assumption should have covered all the possible initial 18O values of quartz. As a result, two limit values 8 and 10 and their median value 9 are chosen as representing the initial 18O values of quartz, for modelling purposes.
Chemical compositions of biotites, alkali feldspars and coexisting plagioclases were determined by electron microprobe (EPMA). With the relations between the Fe/(Fe + Mg) molar ratio, the molar fractions of annite in biotite and the temperature (Wones & Eugster, 1965), biotite geothermometers were estimated. With the equation Kab = (Ab value of alkali feldspar)/(Ab value of plagioclase), where Kab is the equilibrium constant (Barth, 1962), the plagioclase–alkali feldspar geothermometers were calculated. On the basis of geothermometry calculations, biotite, plagioclase and alkali feldspar within the Aral batholith appear to have equilibrated between 580 and 700°C. A temperature of 600°C is chosen as approximately representing the equilibrium temperature of major minerals.
Corresponding initial 18O values of feldspar and biotite were calculated according to the 18OQtzinitial values and the equilibration temperature. The mineral–water fractionation equations that were used are:
- 103ln Qtz–water = 3·55 x 106T-2 – 2·57(195–573°C) (Shiro & Sakai, 1972);
- 103ln
Qtz–water = 3·23 x 106T-2 – 2·94(573–1000°C) (Shiro & Sakai, 1972);
- 103ln
Fsp–water = 2·91 x 106T-2 – 3·41(300–800°C) (O’Neil & Taylor, 1967);
- 103ln
Bt–water = 0·03 x 106T-2 – 2·59 (Bottinga & Javoy, 1973).
- 103ln
Quantitative modelling is performed to constrain the parameters, which include kQtzt, u/kQtz and kQtz/kFsp/kBt. kQtz, kFsp and kBt are the 18O/16O exchange rate constants of quartz, feldspar and biotite, respectively. kQtzt stands for the kQtz-normalized time-duration of 18O/16O exchange; u/kQtz the kQtz-normalized infiltration rate of the aqueous fluid; kQtz/kFsp/kBt the relative exchange rate constants of quartz, feldspar and biotite. Values of kQtzt are varied from 0·1 to 5·0 with a step of 0·1; values ofu/kQtz from 0·1 to 10·0 with a step of 0·5; values of kQtz/kFsp/kBt from 1/5/5 to 1/50/50 with a step of 1/1/1. When the data sets of Groups I and II are consistently enveloped in both Fig. 5a and 5b by the same trajectories with the same u/kQtz values and the same isochrons with the same kQtzt values, the quantitative modelling gives the best fit. The modelling results based on the assumption of 18OQtzinitial = 10 are shown in Fig. 5. kQtzt and u/kQtz are converted to real values t and u on the assumption of kQtz = 10-14 s-1 (Gregory et al., 1989). These parameters obtained through quantitative modelling and their real values are summarized in Table 2. To fit the array of data of Groups I and II with their large spread (2·4) in the measured 18OQtz values, the u/kQtz values of the exchange trajectory and the kQtzt values of the isochron would be required to vary in the range 1–5 and 0·1–1, respectively. These variations are unlikely to be true for Groups I and II within the Aral batholith for the following reasons. Granites within Groups I and II have similar grain sizes. This homogeneity in grain size implies similar permeabilities, which should sustain similar infiltration rates for the external aqueous fluid, and similar cooling rates within Groups I and II, which should allow the exchange with the infiltrating fluid during similar time-durations. Moreover, the large spread in the measured 18O values of quartz have to be assumed as being purely brought about by the 18O/16O exchange, and thus unrealistically large values of kQtzt, and W/R values that range from 0·12 to 5·12 are required to model the data array (Table 2). Still larger values of kQtzt and W/R would be required on the assumption of 18OQtzinitial = 8 or 9. The cumulative flux of metamorphic water or meteoric water within metamorphic terrains is typically small to moderate, with W/R ratios 
1 (e.g. Criss & Taylor, 1986; Ferry, 1986; Gray et al., 1991; Richards et al., 1996). Because the Aral batholith is of kilometre scale, large values of u, t and W/R mean large cumulative fluxes of metamorphic fluids associated with the batholith. Additionally, we have to invoke a large reservoir for the 18O-rich fluid. Therefore, the assumption of a homogeneous initial 18O/16O composition for the granite is not likely.
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In Fig. 6a and b, samples of Groups I and II define a linear array parallel to the 600°C isotherm. The array of data is more likely to have been displaced from an initial set of positions on the 600°C isotherm towards the upper right in Fig. 6a and towards the lower right in Fig. 6b. In other words, before the 18O/16O exchange with the 18O-rich aqueous fluid of the first stage, the data of Groups I and II were already linearly arranged on the 600°C isotherm rather than concentrated at a single initial starting point. Consequently, it is plausible that variations in the 18O values of the coexisting minerals from the samples of Groups I and II depend upon both the initial 18O/16O heterogeneity and later subsolidus 18O/16O exchange with an 18O-rich aqueous fluid.
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Modelling kQtz/kFsp/kBt and u/kQtz for the subsolidus 18O/16O exchange on the assumption of initial 18O/16O heterogeneity—internal vs external exchange
In the following quantitative modelling, the exchange trajectories are assumed to start from an array of initial 18O/16O compositions on the 600°C isotherm. However, as illustrated in Fig. 7, the 18O/16O shifts of individual minerals in the Aral granite, from their initial 18O/16O compositions on the 600°C isotherm during the first stage of exchange, may be a composite product of both internal and external exchange. The overall pattern of the exchange trajectory depends on the relative importance of external mineral–water exchange vs internal exchange between coexisting minerals. For feldspar, internal exchanges that are dominated by the quartz–feldspar 18O/16O exchange tend to lower its
18O value because the quartz–feldspar fractionation factor would increase at a lower temperature, whereas feldspar–water external exchange tends to increase its 18O value. For biotite, both biotite–quartz and biotite–feldspar internal exchanges tend to lower its 18O value because both the quartz–biotite and feldspar–biotite fractionation factors would increase at a lower temperature, whereas biotite–water external exchange should increase its 18O value. We estimate the mineral–water external exchange with ki/kQtz and u/kQtz. The greater these values, the more developed the external exchange compared with the internal exchange.
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If the relative exchange rate kFsp/kQtz is <30, the slope of the exchange trajectories on the 18OFsp–18OQtz diagram would not be sufficiently steep to envelope the data. However, if kFsp/kQtz is >30, the trajectories on the 18OBt–18OQtz diagram do not have a concave-up shape, which plunges from the onset of exchange and then rises, but, rather, steepens and rises linearly following the onset of exchange. Therefore, the trajectories would overlie or miss the data, and thus cannot model the data. This means that the feldspar–water external exchange must be too strong, enhanced by the large exchange rate kFsp, to effectively increase the 18O value of feldspar. This increase in the 18O value of feldspar practically raises the 18OFsp–Bt fractionation and, hence, restrains the biotite–feldspar internal exchange. Thus, the decrease in the 18O value of biotite becomes less remarkable. Meanwhile, if the relative exchange rate kBt/kQtz is >10, the trajectories on the 18OBt–18OQtz diagram tend to steepen and rise linearly following the onset of exchange, and again the trajectories would overlie or miss the data. This means that the biotite–water external exchange must be so developed, as enhanced by the large exchange rate kBt, that it predominates over the biotite–feldspar and the biotite–quartz internal exchanges. As a result, the relative exchange rate kQtz/kFsp/kBt must be close to 1/30/10 as shown in Fig. 6a and b.
The large differences in the exchange rates between quartz and feldspar or biotite imply different exchange mechanisms between quartz and feldspar or biotite. Microscopically, clouding and turbidity are developed within feldspars from Groups I and II within the Aral batholith, providing direct evidence for extensive feldspar–water interaction. A perfect cleavage is penetratively developed within biotite. Therefore, it is probable that quartz exchanged via diffusion whereas feldspar and biotite exchanged via surface reaction (Barnett & Bowman, 1995).
With u/kQtz >> 1, exchange trajectories on the 18OBt–18OQtz diagram tend to rise linearly or do not plunge sufficiently following the onset of exchange so as to include the array of data below the 600°C isotherm. The external exchange with the 18O-rich water is so strong, enhanced by the large infiltration rate u/kQtz >> 1, that it effectively increases the 18O value of feldspar. This increase in the 18O value of feldspar again raises the 18OFsp–Bt fractionation and, therefore, restrains the biotite–feldspar internal exchange. As a result, the decrease in the 18O value of biotite is not obvious. With u/kQtz << 1, exchange trajectories on the 18OFsp–18OQtz diagram cannot attain a sufficient slope to include the array of data above the 600°C isotherm. This means that the u/kQtz value is too low to accomplish the disequilibrium 18O-enrichment of feldspar relative to quartz as indicated by the 18OQtz–Fsp reversals. As a result, the u/kQtz value must approach one (Fig. 6a and b).
The hypothesis of initial 18O/16O heterogeneity plus subsolidus 18O/16O exchange kinetics
Values of kQtzt were varied from 0·00 to 0·50 with a step of 0·01 to achieve the best fit of the exchange isochrons to the data set. The best fit is given at kQtzt values of 0·01 and 0·07. The quantitative modelling results are comprehensively shown in Fig. 6a and b. kQtzt and u/kQtz were converted to real values t and u on the assumption kQtz = 10-14 s-1. These factors and their real values are summarized in Table 2. Data of Groups I and II are consistently enveloped in both Fig. 6a and 6b by the exchange trajectories, with a uniform u/kQtz value of one, which start from an array of initial starting compositions on the 600°C isotherm with 18OQtzinitial = 10·5, 11, 12, 13 and 14. The data are consistently limited in both Fig. 6a and 6b by the isochrons with kQtzt = 0·01 and 0·07. Modelling on the hypothesis of initial 18O/16O heterogeneity plus subsolidus exchange kinetics has two advantages over modelling on the assumption of a single initial starting composition. First, a large increase in the 18O values of quartz from the 18OQtzinitial value is not required. Only the disequilibrium shifts in the 18O values of feldspar and biotite are required to displace the array of data from their initial positions on the 600°C isotherm towards the upper right in Fig. 6a and the lower right in Fig. 6b. Second, the fitted values appear to be reliable. A relatively short lifetime of the fluid system, as indicated by t = 0·03–0·22 Ma (Table 2), and a small W/R ratio ranging from 0·03 to 0·09 (Table 2), are also compatible with the disequilibrium 18O/16O relation.
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CONCLUSION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGY, PETROLOGY AND...SAMPLE DESCRIPTION, EXPERIMENTAL...DISCUSSIONCONCLUSIONREFERENCES |
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The Aral granite batholith in the Altay Mts of China underwent two stages of 18O/16O and D/H exchange with aqueous fluids. The second stage of 18O/16O and D/H exchange with 18O- and D-depleted meteoric water occurred at lower-temperature conditions of post-magma solidification, and has probably recorded a near-surface alteration after uplift of the batholith.
Granites within the four lithological units of Groups I, II, III and IV differ in grain size; correspondingly, they show different sensitivities to the isotopic exchange of the two stages. The megacrystic coarse-grained granites of Group I survived the second stage of 18O/16O and D/H exchanges. The megacryst-bearing medium- to coarse-grained granites of Group II, although essentially surviving the second stage of 18O/16O exchange, were slightly overprinted in terms of their D/H isotopic compositions by the second stage of D/H exchange. Consequently, granites of Groups I and II have preserved the kinetic effects of the first stage of 18O/16O exchange with the 18O-rich aqueous fluid. However, the fine-grained granites of Group IV and the fine- to medium-grained granites of Group III were significantly overprinted by the second stage of isotopic exchange. The D/H isotopic composition of biotite is more easily susceptible to the second stage of meteoric water alteration than is its 18O/16O composition.
The first stage of 18O/16O exchange occurred between feldspars and biotites from Groups I and II of the Aral batholith, and an 18O-rich aqueous fluid during subsolidus cooling, resulting in remarkable 18OQtz–Fsp reversals. The 18O-rich aqueous fluid may have derived from the fluid release from meta-sedimentary rocks during the post-orogenic uplift of the Altay orogen. Our model results suggest that only a relatively short time-duration of fluid–rock interaction such as 0·03–0·22 My and, correspondingly, a small W/R ratio such as 0·03–0·09 can bring about remarkable disequilibrium 18O/16O relations between coexisting minerals. The 18OQtz–Fsp reversal explicitly demonstrates that both the quartz–feldspar and the quartz–biotite arrays of the Aral batholith may represent a pseudoisotherm. This once again illustrates the uncertainties of geothermometry using stable isotopes (Richards et al., 1996).
The large-scale peraluminous Aral granite batholith inherits an initial heterogeneity in both chemical and isotopic compositions (Liu et al., 1997). This type of magma is viscous and its liquidus temperature is typically too low to permit a thorough chemical and isotopic homogenization. This study demonstrates the possibility that variations in the measured 18O/16O compositions of coexisting minerals depend on both the initial 18O/16O heterogeneity, as indicated by the spread in the measured 18O values of quartz, and the subsolidus 18O/16O exchange with an external aqueous fluid.
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ACKNOWLEDGEMENTS |
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I thank Professor R. T. Gregory, Professor G. W. Bergantz, Professor M. Wilson and an anonymous reviewer for providing excellent and constructive comments. This research was funded by the Ministry of Science and Technology of China Grants G1999043204, 95-Y-25 and 95-Y-39, and by the National Natural Science Foundation of China Grant 49373164.
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FOOTNOTES |
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*Telephone: +86-1-(010)64889078. Fax: +86-1-(010)64889849. e-mail: liuw{at}mail.c-geos.ac.cn
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