dergone varying degrees of post-peak chemical shifts that track local variation in proportions of ion-exchang- ing minerals (Fig. 1). Thus, adjacent modal ...
Contrib Mineral Petrol (1996) 124: 346–358
C Springer-Verlag 1996
Robert R. Loucks
Restoration of the elemental and stable-isotopic compositions of diffusionally altered minerals in slowly cooled rocks
Received: 28 March 1995 y Accepted: 11 April 1996
Abstract In polymineralic plutonic igneous and metamorphic rocks, slowly cooled crystals seldom retain their initial chemical compositions. This paper introduces a new, simple and widely applicable material-balance method that recovers the former compositions of minerals – regenerating in many rock types their chemical memory of the environment in which they formed – without a priori knowledge of temperature or pressure or diffusion kinetics. Restored stable-isotopic, trace-element, andyor major-element compositions provide a basis for interpretations of petrogenetic processes and conditions, including recovery of peak temperature and pressure, depth, and average diffusion distance during re-equilibration. Case studies illustrate applications of the mineral-restoration technique to regional crustal dynamics, ore metallogeny, and igneous fluid dynamics and petrogenesis. The first illustrative case addresses the controversial origin of decimetre-thick, modally graded rock layers in the Skaergaard intrusion. Layer-wide mineral-chemistry gradients that previously were ascribed a primary origin are here shown to be due to sub-solidus diffusive re-equilibration amongst minerals that initially were chemically uniform. This finding redefines the constraints to be satisfied by fluid-dynamic models of chemical differentiation processes in the magma chamber, and eliminates the basis of prior interpretations of the modally graded layers as products of in situ crystallization on the magma chamber’s floor. In another case, lower crustal olivine-chromite cumulates underwent a long two-stage history of mineral-composition readjustment spanning .5008 C. The technique introduced here removes the effects of the second-stage solid-state diffusion, recovering mineral compositions that represent the igneous solidus temperature at the termination of the
R.R. Loucks Research School of Earth Sciences, Australian National University, Canberra, A.C.T., 0200, Australia Editorial responsibility: R. Binns
metasomatic stage. The third example removes effects of retrograde diffusive ion-exchange from garnet, hornblende, and clinopyroxene in order to restore the rock’s chemical memory of its pressure and depth of crystallization. The depth corresponds to a measure of extreme Cenozoic uplift and erosion (¥58 km) along the Main Mantle Thrust, which juxtaposes the underthrusting Indian Plate and the over-riding Kohistan island-arc terrain in the Pakistani Himalayas.
Introduction An ability to distinguish which of a sample’s chemical features are associable with its conditions of origin and which features developed later, and an ability to assign a sample of the crust or mantle to the temperature-pressure-depth context in which it originated, are crucial to understanding the processes by which rock types are formed and deformed at depth, and to characterizing dynamic phenomena by which materials of deep origin come to be exposed for sampling. In slowly exhumed upper-mantle and lower-crustal rocks, the usual impasse to thermodynamically evaluating the temperature (T) and pressure (P) of crystallization and quantifying crustal uplift and erosional depth is that the diffusion distance exceeded the crystal size during element- and isotope-exchange reactions as the rock cooled, so mineral compositions attained at peak T were subsequently destroyed throughout. Perversely, diffusional resetting is usually greater in rocks of deeper origin that are potentially the most informative regarding crustal dynamics. This paper introduces a versatile, new method that recovers the lost compositions of minerals without prior knowledge of T or P or diffusion kinetics. The restored stable-isotopic andyor elemental compositions provide a basis for inferring the physico-chemical environment of formation and the operative fluid-dynamic and other petrogenetic processes. The T retrieved is usually tens to hundreds of degrees higher than that calculated directly from compositions of grain cores in the same rock; the
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calculated P tends to rise correspondingly by up to several kilobars. This extension of a sample’s memory enables recovery of temporally matched fossil T and P values. Mismatch of the time-temperature conditions at which mineral geothermometers and geobarometers in the same sample became “closed” to further compositional alteration is a major hazard to reconstructing evolutionary P–T paths of mobile terrains (Frost and Chacko 1989; Selverstone and Chamberlain 1990). Mismatch of the closure conditions of geothermometers and geobarometers renders spurious many published interpretations of regional crustal dynamics.
Principle and procedure Many rock outcrops that contain a petrogenetically informative element- or isotope-exchanging mineral suite have centimetre- to decimetre-scale heterogeneities in proportions (modes) of mineral species, e.g. bands, streaks or mottles. In rocks with modal domains (bands, mottles) developed on an appropriate scale, mineral compositions that were uniform amongst chemically equilibrated, adjoining domains at peak T will have undergone varying degrees of post-peak chemical shifts that track local variation in proportions of ion-exchanging minerals (Fig. 1). Thus, adjacent modal domains that were once mutually equilibrated have become mutually disequilibrated by development of different compositions of the same minerals. Measuring the difference in mineral compositions amongst adjacent modal domains and mathematically reversing these chemical shifts is the goal of this analysis. It requires: (1) individually sampling adjacent modal domains that share a mineral assemblage of interest; (2) provisionally adopting the
Fig. 1 A hypothetical rock comprises two modal domains, 1 and 2, that both contain minerals A and B which share exchangeable Fe21 and Mg21 ions. Modal domain 1 has 75 mol% A, 25 mol% B, and domain 2 has 20 mol% A, 80 mol% B. At temperature T0 the mineral compositions were A0 and B0. During cooling to Tf, shortrange diffusive exchange altered their compositions to Af1 and Bf1 in domain 1 and to Af2 and Bf2 in domain 2, whilst maintaining unchanged molar proportions of phases A and B in each domain
testable hypothesis that domains which are now mutually disequilibrated formerly had minerals of the same compositions; (3) judiciously selecting a sample spacing larger than the characteristic distance of retrograde diffusion, but small enough for mutual equilibration amongst domain at peak T (finding a suitable sample spacing and size may entail iterative trials by combining or splitting the initial spacings); (4) selecting samples in which the proportions of ion-exchanging minerals present at peak T are measurable – usually meaning that minerals of interest lack retrograde replacement textures. If these conditions are met, then the chemical and modal bulk composition of each domain’s exchanging assemblage has remained unchanged (closed) as minerals within each domain re-equilibrated by local exchange. Here is the key point: after such intra-domain exchange, a particular mineral’s chemical variation amongst neighbouring, chemically closed domains correlates linearly with modal ratios of ion-exchanging minerals in the domains and is wholly attributable to retrograde re-equilibration. Graphical or algebraic analysis can reconstruct the original mineral compositions from the measured modal proportions and chemical compositions of their diffusionally reconstituted descendants. Figure 1 illustrates the simplest case of two minerals exchanging two components in two domains. A hypothetical banded rock consists of minerals A and B that share an exchangeable chemical constituent – Fe21 and Mg21 ions in this example (although an example entailing exchange of isotopes – e.g. 16O and 18O – would follow an identical graphical procedure). Modal domain (band) 1 has 75 mol% of mineral A, 25 mol% of mineral B, and domain 2 has 20 mol% A, 80 mol% B. The mineral proportions are to be expressed in exchange-equivalent molar units; for Fe–Mg exchange, olivine is (Mg,Fe)Si0.5O2 if clinopyroxene is Ca(Mg,Fe)Si2O6), etc.. At elevated temperature T0 the mineral compositions were A0 and B0. During cooling to Tf, short-range (relative to domain size) diffusive exchange altered their compositions to Af1 and Bf1 in domain 1 and to Af2 and Bf2 in domain 2, whilst maintaining unchanged molar proportions of phases A and B in each domain. We seek a technique to mathematically “reverse” that diffusive alteration. The rock exposure is sampled in its diffusionally retrograded state. Samples 1 and 2 are collected as slices from modal domains 1 and 2. The Ay(A1B) phase ratio is then measured as 0.75 in sample 1 and as 0.2 in sample 2. Mineral compositions are measured as Af1 and Bf1 in sample 1 and as Af2 and Bf2 in sample 2. These compositions are next connected by tie lines Af1–Bf1 and Af2–Bf2, and their rotation pivots are then plotted at 0.75 on sample 1’s tie line, and at 0.2 on 2’s tie line. Then a third tie line is drawn through pivots 1 and 2 to find its abscissa intercepts as A0 and B0, which are the restored, original mineral compositions. The algebraic equivalent of this graphical procedure is based on material balance of each domain’s bulk composition with the abundances and compositions of its constituent minerals:
348 Mg'D5Mg'A ? YA1Mg'B ? YB
(1)
wherein Mg' denotes the molar ratio Mgy(Mg1Fe), D designates the whole modal domain comprising phases A and B in proportions denoted by the molar mode fractions YA5Ay(A1B) and YB5By(A1B); YB512YA. For each of the samples 1 and 2, Eq. (1) identifies the coordinates (YA, Mg'D) of its tie-line-rotation pivot. Rearrangement of Eq. (1) gives YA5(Mg'D2Mg'B)y (Mg'A2Mg'B) and 1 Mg'B ? Mg'D2 (Mg'A2Mg'B) (Mg'A2Mg'B)
YA5
(2)
The slope of any segment of the line is the same as the slope of the whole line, so the slope is
For extension of Fig. 1’s material balance to n exchanging minerals in k .n modal domains, multidimensional linear regression can be used. If the minerals partition oxygen isotopes amongst them, and if minerals A, B, C…n have molar mode fractions YA, YB…Yn (defined using oxygen-equivalent moles – e.g. Fe3O4, Fe1.33Si1.33O4, and Si2O4 for magnetite, ferrosilite, and quartz), then the whole-rock material balance in the jth domain is d18Oj5YAjd18OAj1YBjd18OBj1…. For measured d18Oj in k samples, least-squares multiple linear regression of n–1 mode fractions YAj, YBj… versus d18Oj of each sample gives a best-fit equation from which the restored initial d18On of each mineral n is obtained by consecutively setting the nth mode fraction to 1.0 and the rest to zero.
YA12YA2 YA(max)2YA(min) dYA 5 5 dMg'D Mg'D12Mg'D2 Mg'D(max)2Mg'D(min) 5
(120) (Mg'A2Mg'B)
How modal domains mutually equilibrate and close (3)
Equation (2) is the one that corresponds to the reconstructed tie line that gives the restored, initial mineral compositions as its abscissa intercepts Mg'A0 and Mg'B0. The upper and lower abscissa intercepts of the tie line are found by consecutively setting each of the mode fractions to unity and the other to zero in Eq. (2). When YA51, Eq. (2) reduces to Mg'D5Mg'A (i.e. Mg'A0). When YB51, then YA50, and Eq. (2) gives Mg'D5Mg'B (i.e. Mg'B0). If a mineral pair in more than two distinct domains can be analysed to provide a mathematically over-determined set of tie lines and rotation pivots, then leastsquares linear regression through the tie-line-rotation pivots rigorously tests, by collinearity of the pivots, whether the sampled domains were once mutually equilibrated and have remained chemically closed since then. The principle of this procedure is analogous to the whole-rock-isochron method of radiometric dating, which assumes that mineral compositions may have been reset, but whole-rock samples of, say, hand-specimen size have remained compositionally closed. In an isochron plot, the goodness of fit of the whole-rock data points to a straight line serves as a check of the working hypothesis that all of the analysed specimens have remained compositionally closed with respect to the elements of interest. The quality of the linear least-squares fit to an over-determined data set is the test of the working hypothesis listed as item (2) in the first paragraph of this section. If the number of constraints (analysed domains) equals the number of unknowns (mineral compositions) to be evaluated (as in Fig. 1), then hypothesis (2) is an a priori assumption that is accepted on faith as the basis for a solution. However, if the data set has more analysed domains than mineral compositions to be restored, then item (2) is not an assumption, but rather is a working hypothesis that is tested (true or false) and evaluated (quality of fit) by the statistical analysis. In the over-determined case, the restored mineral compositions rest on a foundation containing no assumptions at all.
The procedure described in the preceding section allows restoration of mineral compositions to some prior state in the sample’s evolutionary history. The state represented by the restored compositions may or may not correspond to the highest T the rock has experienced. Interpretations of the geologic processes that produced the condition represented by the restored compositions must vary with the geologic context of the sample and are the responsibility of each investigator. However, by way of introduction to some illustrative examples of the method’s applications in the following sections of the paper, this section briefly considers some of the most common processes or states that are likely to be represented by the restored mineral compositions. During heating, prograde metamorphism of hydrous minerals generates intergranular fluid that provides a medium for grain-boundary diffusion, creates crack permeability, and advects solutes down the fluid pressure gradient (Walther and Orville 1982). Advection and fast diffusion via an intergranular vapour or melt film promote homogenization of mineral compositions in neighbouring bands of gneisses having cm-scale band widths. At T*650–7008 C, millimetre-scale diffusion erases prograde compositional growth zoning in garnet and less refractory minerals (Yardley 1977; Tuccillo et al. 1990). After attaining peak T in the regional granulite or eclogite facies, devolatilization reactions cease in and below the rock volume of interest. Crack permeability closes by annealing. Intergranular melt solidifies, and vapour is resorbed as hydroxyl- and carbonate-bearing minerals, etc.. Sluggish, short-range solid-state diffusion generally supplants advection as the transfer process by which minerals in modal layers undergo retrograde shifts in composition. This transition at or near peak T in the dominant mode of chemical transfer – from longer range fluid advection to shorter range solid-state diffusion – tends to preserve in high-grade rocks those prograde features that developed on a scale larger than the typical distance of retrograde diffusion. This advection-to-diffusion transition T at which the rock became composition-
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ally closed at the sampled scale is recoverable by the technique described in the preceding section. In relatively dry, high-grade rocks, the restored transition T is likely to correspond to peak T. In contrast, strongly hydrated, lower grade metamorphic rocks are generally less promising candidates than higher grade, dehydrated ones for restoration of peakgrade mineral compositions by the method introduced here, because hydrous rocks are more prone to H2Ocatalysed retrograde replacement reactions that alter mineral proportions. Hydrous amphibolite- and greenschist-facies rocks are more likely than dehydrated deep granulites and eclogites to undergo retrograde chemical alteration by fluids released as prograde reactions continue at greater depth (Yardley et al. 1991), or by invasion of surface-derived fluids (Ferry 1992). Chemical exchange also affects minerals in cooling plutonic igneous rocks. During solidification of modally banded igneous crystal cumulates, advection or convection of interstitial melt through the compacting crystal pile continues until permeability is eliminated by precipitation of post-cumulus rinds on primocrysts and by pressure-solution lengthening of grain contacts (Irvine 1980b; Morse 1986). After advection has locally ceased, material transfer is limited to diffusion via trapped pore liquid on a scale commensurate with pore size. Hence, adjacent modally uniform bands or modally graded cycles developed at, say, the decimetre scale become closed to chemical transport beyond centimetre or millimetre distances. Thereafter, evolution of mineral compositions within cooling layers reflects very local modal heterogeneities perceptible at distances typical of the diffusion distance. In layered adcumulates (having little or no trapped liquid), the T at which the rock composition attains local closure is likely to correspond closely to primocryst precipitation Ts, i.e. the ambient liquidus T (Irvine 1980b). However, in meso- and orthocumulates (having more trapped melt), closure of local bulk-rock composition at the hand-specimen or thin-section scale may occur tens or hundreds of degrees below primocryst precipitation Ts. This condition at which the rock’s bulk composition ceased to evolve at the sampled scale is the condition to which the mineral compositions are restored by the technique introduced in Fig. 1. Recovery of the Ts of closure of the sample’s bulk compositions is of value in modelling stages in the evolution of magmatic cumulates, but, as always, the investigator must be alert in identifying the process or stage that is represented by the closure T. The rest of the paper comprises three case studies illustrating diverse applications of the mineral restoration method to deciphering igneous petrogenesis, crustal dynamics, and ore metallogeny.
Case 1: Skaergaard’s graded rhythmic layers Field-based studies encounter riddles that guide or misguide most of the experimental and theoretical research in geophysics and geochemistry. As the classic example of extreme chemical differentiation of magma by crystal-liquid segregation, East Greenland’s Skaergaard layered intrusion is a landmark in explorations of the origins of the chemical diversity of igneous rocks. It is the paradigm to which most laboratory- and field-based studies of tholeiitic magma differentiation are explicitly or implicitly referred. After half a century of intensive study, petrologists are not fine-tuning details of its solidification processes; they are still arguing over its most elementary aspects.
The keystone question The origin of decimetre- to metre-thick layers across which minerals are graded in chemical compositions and modal proportions is a key issue in the continuing debate about how cumulus (liquidus) crystals accumulate at the magma chamber’s floor and eliminate interstitial liquid. McBirney and Noyes (1979), McBirney (1985), McBirney and Naslund (1990), and McBirney (1995) have argued that crystallization along the roof, walls, and floor released silica-depleted, iron-enriched, increasingly dense residual liquids that sank to the floor and produced the Skaergaard Layered Series as a crystallization front rose gradually against non-convecting magma. During in situ crystallization at the floor, differential diffusion of rejected chemical components and of latent heat released at the crystallization front led to oscillatory oversaturation and nucleation pulses of different phases in different levels of the basal boundary layer (McBirney and Noyes 1979). In all zones of the 2.5 km-thick Layered Series of floor cumulates, this diffusive process allegedly produced modally graded layers from Fe-rich liquids so dense that plagioclase crystals would have floated if they had not nucleated and grown attached at the floor. In this section of the paper, the technique of reconstructing diffusively altered mineral compositions is deployed to contest that conclusion. It demonstrates that the olivine, augite, and plagioclase that presently are graded in composition across the layers were initially of uniform composition throughout the graded layers when they were deposited. This finding has major implications regarding issues central to understanding the processes of chemical differentiation of mafic magmas by crystalliquid segregation – viz., (1) the trends of density and chemical composition followed by the evolving magma in the Skaergaard and other layered intrusions having the same cumulus-mineral crystallization sequence; (2) the magma chamber’s fluid-dynamic regime, including; (3) the relative importance of compaction versus porousmedium convection as the main process by which interstitial pore space is eliminated from floor cumulates.
350 Fig. 2 a Solid tie lines link molar Mg' [;1003Mgy(Mg1Fe21)] of coexisting augite (Aug) and olivine (Ol) measured by Conrad and Naslund (1989) (open circles) in an upwards succession of samples (1–6) spanning modally graded layers L1 and L2A in Upper Zone A of the Skaergaard intrusion. The weighted mean Mg' of augite1olivine in each numbered sample is identified by a solid dot that has functioned as the pivot for rotation of its tie line during diffusive Mg-Fe21 exchange between olivine and augite. The collinear pivots identify a “fossil” tie line (dashed) that was formerly shared by all samples domains within the modally graded layer. The abscissa intercepts (*) of the dashed tie lines identify restored mineral compositions coexisting with silicate melt at the temperature of the layer’s transition from chemical transport dominated by advection to transport by volume diffusion on a length scale smaller than the sample dimensions and inter-sample spacing. The temperatures on measured and restored tie lines are obtained from an experimental calibration of the olivine-augite FeMg-exchange geothermometer (Loucks in press a). Using subscripts to represent the Mg' of olivine (O) and of augite (A), the relative molar percentages [1003Oy(O1A)] and compositions plotted and linearly regressed are the following: layer L1, samples 1–6 (no 5): 1 61.1% O39.13, 38.9% A55.51; 2 48.0% O38.98, 52.0% A55.32; 3 23.1% O37.56, 76.9% A54.65; combined 416 9.6% O37.16, 90.4% A54.16; layer L2A samples 1–6 (no 5): 1 52.8% O37.38, 47.2% A54.24; 2 53.9% O37.10, 46.1% A54.50; 3 31.6% O36.24, 68.4% A53.73; 4 4.7% O33.86, 95.3% A53.46; 6 15.7% O35.61, 84.3% A53.74. b In a trace-element restoration, tie lines link Sr contents of augite (A) and plagioclase (P) reported by McBirney and Noyes (1979) in an upwards succession of samples (1–6) from a 10 cmthick modally graded olivine-diorite layer in Upper Zone A. For visual clarity of pivot collinearity, only the central segments of tie lines are shown (ordinate interval 35–75% out of the range 0– 100%). Each numbered dot represents the weighted mean ppm Sr in the augite1plagioclase mixture, which contains virtually all the Sr in the rock. The dots step systematically from one tie line to the next to form a collinear array (r250.9998) that represents a “fossil” tie line (not drawn) of flatter slope than measured tie lines. The fossil tie line links the restored Sr contents of augite (5.1+2.3 ppm) and plagioclase [596+2 ppm (1s)] that initially were uniform throughout the graded layer. Counterclockwise rotation of tie lines during re-equilibration is in the sense experimentally demonstrated (Sun et al. 1974) to occur during cooling
is (Mg,Fe)1,285Si0.6425O2.570. Using this mole to convert from volumetric to molar units, Fig. 2a shows the molar proportions and compositions of olivine (Ol) and augite (Aug) in layers L1 and L2A. Results
Data The key data set is that of Conrad and Naslund (1989), who give exact descriptions of the variations in mineral proportions, compositions, and grain size in two thin-section series spanning an 11.5 cm-thick layer (L1) and a 16 cm-thick layer (L2A) of modally graded, adcumulate, olivine-ferrodiorite within the 270 m-thick Upper Zone A unit of the Layered Series. The layers grade from olivine- and Fe-Ti-oxide-rich bottoms to plagioclase-rich tops. In each layer, ferromagnesian minerals become Mg-depleted upwards. The Fe-Ti oxides and plagioclase can be ignored in the Fe-Mg-exchange material balance, as the proportions and low Fe-Mg-ion-exchange capacities of Fe-Ti oxides and feldspar give them trivial leverage on the post-cumulus shift of MgyFe21 in olivine and augite. Broad-beam electron-microprobe analysis was used to reintegrate augite-pigeonite exsolution intergrowths. Throughout layers L1 and L2A, the reintegrated augite has 1.285+0.013 (1s) atoms of Mg1Fe21 per 6-oxygen formula unit (Conrad and Naslund 1989), so the exchange-equivalent olivine mole
The important result in Fig. 2a is that collinearity of tie-line rotation pivots proves that all samples in each layer formerly shared the same tie line – i.e. formerly shared identical mineral compositions (found at the ends of the reconstructed fossil tie line). The collinearity of the rotation pivots also means that each sampled domain has remained a chemically closed system (at thin-section scale) during retrograde diffusive re-equilibration. For layer L1, least-squares linear regression of mol% Ol versus molar Mg' [;1003Mgy(Mg1Fe21)] of each Ol1Aug mixture yields r250.9993 for the recovered fossil tie line Mg'553.8720.13563(mol% Ol). The restored igneous mineral compositions at its abscissa intercepts are olivine Mg'540.31+0.10 and augite Mg'553.87+0.07 (1s). These may be recast as initial KD;(FeyMg)Oly(FeyMg)Aug51.729+0.001 at closure of modal-domain bulk compositions. For layer L2A, regression gives r250.9983 for the recovered tie line Mg'553.1920.15033(mol% Ol). Its abscissa intercepts give the restored initial mineral compositions olivine Mg'538.16+0.19 and augite Mg'5
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53.19+0.12 (1s), and KD51.841+0.003 at closure to advection of intercumulus melt. Although geothermometry is tangential to the current discussion topic of the origin of intra-layer zonation of mineral compositions, confidence in the restored olivine and augite compositions may be reinforced by using the restored KDs for L1 and L2A to calculate igneous crystallization temperatures, according to the precisely calibrated olivine-augite Mg-Fe21-ion-exchange geothermometer (Loucks, in press a). Application to the Upper Zone (UZ) A samples yields 1083+68 C for L1 and 1071+7 (1s)8 C for L2A located 25 m farther up section. (The standard-error figure corresponds to the standard error of KD that arises from the standard error of the abscissa intercepts on Fig. 2a; the geothermometer’s calibration also has a standard error of +68 C). These temperatures may be compared with 1068+¥408 C from plagioclase thermometry (Morse et al. 1980), and 1064+268 C from diffusionally retrograded olivine 1pigeonite1magnetite (Williams 1971) lower in UZA. McBirney and Noyes (1979) described another olivine-diorite graded layer in UZA. Layer-wide gradients in major- and trace-element composition of each mineral were interpreted as primary and ascribed to in situ crystallization from a diffusively zoned liquid boundary layer at the magma chamber’s floor. However, Fig. 2b demonstrates collinearity of the rotation pivots of tie lines linking the measured Sr contents of plagioclase and augite. That collinearity means that layer-wide chemical zoning of each mineral is not primary, but instead is due to later diffusive exchange of Sr among neighbouring crystals. The Fe, Sc, and Al contents of augite and plagioclase show the same collinearity of rotation pivots, but are not plotted in Fig. 2. One of the important conclusions to be drawn from Fig. 2 is that its demonstration that the principal cumulus minerals were initially of uniform composition across the modally graded layers eliminates McBirney and Noyes’ (1979) key item of evidence that such graded layers formed by in situ crystallization at the floor. From their observation that minerals in the modally graded layers are sorted by density, Conrad and Naslund (1989) inferred that primocrysts in those layers were sedimentary and therefore were initially chemically uniform across each layer. They interpreted the layer-wide gradients in mineral chemistry as having developed during cementation of pore space, due to gradation of the volumetric ratio of each kind of primocryst to its overgrowth rind that grew from chemically evolving trapped interstitial liquid – e.g. they inferred that in parts of the layer where olivine primocrysts were sparse, the proportion of Fe-enriched overgrowth olivine was large. In their interpretation, subsequent solid-state diffusion homogenized the cores and rims of the crystals to produce unzoned crystals, but preserved layer-wide gradients in the composition of each mineral. Conrad and Naslund’s interpretation that the intra-layer chemical zonation developed by an igneous process (above the solidus) is refuted by Fig. 2’s demonstrations that: (1) all (not just pri-
mocryst cores) of the olivine and augite now present in the rock formerly had uniform chemistry across each layer; (2) all of the layer-wide chemical zonation in augite, olivine, and plagioclase composition is accounted for by sub-solidus ion-exchange by diffusion on a “modal-domain” scale that is smaller than the thin sections. Implications Figure 2 implies that pore space was eliminated by development of grain overgrowths that are compositionally indistinguishable from the olivine and augite primocryst cores at all levels of each layer. (Even in plagioclase, which does preserve compositionally zoned post-cumulus overgrowths, the core-to-rim composition range in UZA is usually only 4 mol% An, and the mean composition of the grains is within 1 mol % An of the maximum value found in each sample; Maaløe 1976; Conrad and Naslund 1989.) The correspondence of grain core and rim compositions means that interstitial melt was exchanged for material of liquidus composition whilst crystals growing throughout the 16 cm-thick layer were in good chemical communication with the main body of magma. This observation is of the most fundamental importance to establishing what was the chemical differentiation trend of the Skaergaard magma, and what were the fluid-dynamic mechanisms by which that differentiation trend was achieved. The efficient chemical communication between pore liquid and bulk magma that is required for development of isochemical adcumulates implies that most cementation occurred near the top of the crystal pile. This efficient exchange of components between pore space and bulk magma effectively eliminates McBirney’s (1995) hypothesis that compaction by overlying cumulates was the main mechanism of consuming pore space and expelling interstitial melt. Figure 2’s implication that most adcumulate cementation occurred near the top of the crystal pile is corroborated by Irvine’s (1980c, Fig. 5) illustration of a 436 metre gabbroic xenolith that fell from the roof and hit gently sloping, layered floor cumulates of the Middle Zone. The impact scraped up a small “bow wave” of unconsolidated crystals at the site of impact, but it damaged the cumulate layer beneath the xenolith to a depth of only a few tens of centimetres. The shallow penetration implies that the depth of unconsolidated crystal mush was thin indeed. The results obtained in Fig. 2 accord nicely with a model of the Skaergaard’s chemical evolution (Hunter and Sparks 1987, 1990; Sparks 1988; Snyder et al. 1993; Toplis and Carroll 1995) in which liquidus magnetite saturation atop Lower Zone B initiated a monotonic trend of iron depletion and silica enrichment in the magma, producing residual, hydrous, buoyant liquids of low enough density for feldspar to sink. This trend mimics the Fe-Ti-basaltic → ferroandesitic → dacitic liquid sequence in H2O-free experiments that reproduce advanced differentiation of mid-ocean-ridge basaltic magma (Juster et al. 1989). This differentiation trend can ex-
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plain growth of adcumulate textures in modally graded layers by porous-medium convection (Kerr and Tait 1986). Buoyant liquid also accords with evidence that mafic → felsic density-graded crossbeds and parallel layers represent winnowing by currents reworking the chamber floor, and gravitational sorting of crystals settling during recurrent magma-stagnation episodes (Sparks et al. 1993) or settling from avalanche slurries of hydrated, 16O-enriched magma that swept down the walls and across the floor during (seismic?) disturbances that also dislodged swarms of low-d18O, hydrothermally altered roof-rock fragments that are concentrated in graded layers (Wager and Brown 1968; Irvine 1980a; Taylor and Forrester 1979). That these graded layers formed as gravity-sorted, settled sediments (Wager and Brown 1968; Irvine 1980a; Conrad and Naslund 1989), not by in situ floor crystallization (McBirney and Noyes 1979; McBirney 1985, 1995), is clear from four lines of evidence that jointly are compelling: (1) Fig. 2’s demonstration that mineral-chemistry gradients are not a primary crystallization feature, but developed by nearsolidus or sub-solidus re-equilibration of initially uniform minerals; (2) strong grading of layers by mineral density (Conrad and Naslund 1989); (3) in the four similar graded layers in UZA in which feldspar oxygen isotopes were analysed, d18Oplag is systematically lower than in intervening ungraded layers (Taylor and Forrester 1979), which means that primocrysts in the graded layers nucleated and grew elsewhere in the magma chamber before transport to their deposition site; (4) graded layers are selectively enriched in xenolithic fragments of the Marginal or Upper Border Series (Irvine 1980a, plates 1 and 2), where circulation of low-d18O meteoric waters altered the rocks during solidification of the intrusion (Taylor and Forrester 1979). The major conclusion from these four observations is that buoyant escape of interstitial liquid from loose, settled plagioclase crystals at the tops of quickly deposited, decimetre-thick graded layers – leaving adcumulates – requires drastic revision of widely accepted models of the Skaergaard magma’s chemical and density evolution (Wager 1960; McBirney and Noyes 1979; McBirney and Naslund 1990; McBirney 1995) and fluid dynamics (McBirney 1985, 1995). The buoyant-liquid density trend requires accumulation of more SiO2 and Al2O3, and enough volatiles to explain the late pegmatites and exsolved magmatic brine (Sonnenthal 1992) and to accommodate the compelling evidence cited above that anorthositic xenoliths and plagioclase crystals that originated near the roof were deposited on the floor under the influence of gravity. McBirney (1995) advocates compaction as the only important agent of removing pore liquid from the floor cumulates, because he is committed a priori to a model of chemical evolution of Skaergaard magma that produced the entire Layered Series from increasingly dense residual liquids that could not spontaneously depart the pore space by buoyancy. His model entails liquid 6% denser than plagioclase (¥An65) at the base of the Lower
Zone, 12% denser than plagioclase (¥An49) in the Middle Zone, 13% denser than plagioclase (An43) at the base of the Upper Zone (UZA), and 17% denser than plagioclase (An34) in the Sandwich Horizon at the top of the Upper Zone. McBirney and Naslund (1990) and McBirney (1995) cite as the basis of their model of Skaergaard liquid density evolution a series of partial-melting experiments by McBirney and Nakamura (1974) that attempted to regenerate and analyse the trapped interstitial melt in Skaergaard cumulate rocks. Among the defects of those experiments are two fundamental conceptual flaws in their design: Firstly, if the primocryst assemblages that developed throughout the Middle and Upper Zones were more mafic than their parental liquids (as suggested in the preceding paragraph), then the buoyancy of the interstitial melt would have allowed it to convect away and be replaced by liquid that replenished the supply of femic components needed for continued adcumulus growth of mafic minerals – thus assuring that the pore space cement, as well as primocrysts, would have an average composition more mafic than its parental liquid. If McBirney and Nakamura’s experiments had actually succeeded in their intent to melt only the post-cumulus portion of the rock, they would have generated a liquid more mafic than any natural liquid that had occupied the interstices. Their experimental design dismissed a priori the possibility of buoyant residual liquids and of cumulate cementation by the scenario presented above. Secondly, documented evidence of near-solidus brine exsolution at P$1 kbar (Sonnenthal 1992) requires $3 wt% H2O in the melt by that stage (Burnham 1979). McBirney and Nakamura (1974) did only anhydrous partialmelting experiments on samples from the Middle and Upper Zones. Their exclusion of H2O biased the composition of the partial melts to inappropriately low Al2O3 and SiO2, and high FeO and high density, by suppressed melting of plagioclase and pyroxene and enhanced melting of olivine, relative to appropriately hydrous melts (Spulber and Rutherford 1983; Holloway and Burnham 1972; Helz 1976). Conclusion from case 1 Of the models proposed to the present date for magmatic differentiation processes in the Skaergaard intrusion, the only one consistent with the results in Fig. 2 and the other evidence enumerated as items 1–4 above is the fluid-dynamic and chemical-differentiation model of increasingly siliceous, hydrous, buoyant residual liquids that was proposed by Hunter and Sparks (1987) on the basis of comparisons with differentiated theoleiitic volcanic suites, and proposed by Snyder et al. (1993) on the basis of extensive and well controlled experiments on a widely accepted parental liquid composition found in the Skaergaard’s quenched margin.
353
Case 2: uncovering chemical effects and temperature regimes of metasomatism overprinted by solid-state diffusion The Chilas mafic igneous complex in the Kohistan terrain, Pakistani Himalayas, formed as sub-batholithic cumulates from hydrous basaltic magmas in the lower crust of a Cretaceous Tethyan intra-oceanic island are (Miller et al. 1991). Its Thurly Gah olivine-cumulate unit precipitated upon sporadically copper-mineralized gabbronorite atop the preceding magmatic-differentiation cycle. Magma-chamber replenishment by the batch parental to the Thurly Gah unit accompanied deformation of the magma-chamber floor by large-scale folding, faulting, and debris slides of semi-consolidated cumulates over an area .200 km2 (R.R. Loucks, D.J. Miller, and M. Ashraf, unpublished work). Field relations suggest that anticlinal folds in the permeable gabbronoritic cumulates focussed migration of residual, hydrous, buoyant interstitial melt up through widely separated 1– 2 km2 portions of the magma-chamber floor during and after deposition of Thurly Gah unit. The syn- and postcumulus flux of interstitial melt (cf. Irvine 1980b, Kelemen 1990) is postulated to have produced regions of dunite-hosted, large-tonnage, low-grade, uniformly disseminated chalcopyrite1bornite1gold1platinum-palladium-telluride mineralization in portions of the Thurly Gah unit that overlie the crests of anticlinal folds in underlying cumulates. (Evidence of platinum group elements mineralization by analogous post-cumulus magmatic-infiltration metasomatism is well documented in the Merensky Reef (Mathez 1995) and platiniferous dunite pipes (Rudashevsky et al. 1992) of the Bushveld Complex.) The Thurly Gah sulphide assemblage has Cu*Fe and Cu:Ni .4 that are very atypical of cumulus sulphides in ultramafic rocks (Naldrett 1989). The ore has 1–3% hornblende, trace phlogopite and diopside, no plagioclase, and has variable amounts of chromite as disseminations and discontinuous thin layers. The undeformed, fresh, mineralized dunite is pervasively recrystallized to granoblastic textures. As shown below, infiltration of exotic intercumulus melt reset the compositions of olivine and chromite. The problem at hand is to employ olivine-chromite Fe-Mg-exchange geothermometry to characterize the thermal regime in which the putative infiltration metasomatism altered the cumulus Fe-Ni-rich sulphides in the dunite to produce mineralogically and chemically peculiar Cu-Au-Pt-Pd ore.
Data Figure 3a shows two bands of olivine-chromite cumulates. The microprobe traverse shows flat plateaux for mineral compositions within bands. Steep gradients in mineral compositions between bands are inferred to represent the limits of solid-state diffusion across the band contact. Table 1 lists whole-rock wet-chemical analyses
Fig. 3 a A thin section 30 mm long shows isomodal bands of chromite dunite (C12-805Ad, left) and olivine chromitite (C12805Ac, right). The electron-probe traverse shows relatively flat profiles of local mean (grain cores1rims) molar Mgy(Mg1Fe21) in olivine (Mg'Ol) and spinel (Mg'Sp) within the layers, and steep gradients in mineral chemistry at the layer contact. b Using the coordinate frame introduced by Irvine (1965), solid lines link mean compositions of spinel and olivine in each of three layers differing in modal spinel: olivine ratio at field locality C12-805 and in a pair of layers at locality C10-100. A dot on each tie line identifies the spinel modal proportion (1003Spy(Sp1Ol), in M21 cation-equivalent moles) and weighted mean Mg' of combined olivine1spinel in each analysed modal domain (i.e. the bulk-rock composition in Table 1, projected from clinopyroxene and hornblende having the abundances and compositions tabulated). Linear regression through these tie-line-rotation pivots yields the dashed tie line having abscissa intercepts that are the recovered spinel and olivine compositions that formerly were shared by all the adjacent samples analysed. For C12-805, standard errors of the abscissa intercepts are from the regression (Till 1974). The recovered igneous mineral compositions, pertaining to closure of the modal domains to melt advection, yield calculated palaeo-temperatures hundreds of degrees higher than the average temperatures of closure of grain compositions to diffusive exchange, which are identified on the solid tie lines
6.37 0.03 41.1 45.4
40.7 41.1 40.0
41.3 43.0 41.0 50.9
2 15 21 2
2 19 7
2 16 10 4
WR SP OL HB
WRe HBf GA
WR HBg GA CP
J2-6800H
J2-6800G
0.40 0.89 0.05 0.78
1.52 1.04 0.06
0.89 1.18 0.00 1.02
0.05 0.96 0.00 0.17
0.52 0.77 0.00 1.20
0.28 0.70 0.00
0.11 0.64 0.00 1.02
TiO2
19.7 15.5 22.8 5.34
17.1 16.9 22.6
18.5 21.1 0.00 11.3
0.65 20.5 0.00 1.51
12.8 19.0 0.00 11.3
8.08 19.8 0.00
2.63 19.5 0.00 11.1
Al2O3
0.02 0.01 2 0.03
0.05 0.06 0.07
27.9 32.1 0.01 1.17
0.96 30.2 0.01 0.31
24.9 39.2 0.01 1.75
15.6 37.8 0.02
4.57 36.3 0.01 1.56
Cr2O3
0.04 0.08 0.01 0.07
0.07 0.12 0.04
0.21 0.25 0.00 2
0.00 0.27 0.00 2
0.13 0.20 0.00 0.05
0.09 0.21 0.00
0.03 0.21 0.00 0.14
V2O3
1.39 0.88 1.13 1.40
1.05 0.81 1.72
11.6 14.3 2 1.34
0.38 15.1 2 0.14
6.56 10.1 2 2
4.46 10.7 2
1.42 11.5 2 2
11.3 6.26 14.6 2.81
8.48 7.12 14.3
19.0 21.3 9.37 4.25
14.2 25.5 14.2 2.78
15.5 18.8 9.09 3.61
13.5 20.9 8.40
11.6 23.2 10.1 4.25
Fe2O3b FeO
b
Number of crystal analysed Fe31yFe21calculated from mineral stoichiometric charge balance and mode c H2O not included d Weight percent mode obtained by least-squares mass balance of mineral and whole-rock compositions. Sample C12-805Ac has 2.2 wt% HB, assumed to be same composition as in C12-805Ad
a
C10-100C
39.5 0.03 40.1 53.9
2 11 19 8
WR SP OL CP
C10-100D
14.4 0.04 41.2 45.7
2 12 12 2
WR SP OL HB
C12-805B
24.2 0.04 41.1
2 42 48
WR SP OL
C12-805Ac
35.6 0.04 40.8 45.3
2 30 32 2
WR SP OL HB
C12-805Ad
SiO2
Phase
Sample
na
0.01 0.01 2 2
0.01 0.01 2
0.09 0.11 0.01 2
0.03 0.30 0.01 2
2 0.12 2 2
0.06 0.11 0.02
0.03 0.12 0.01 2
ZnO
0.19 0.03 0.28 0.03
0.04 0.03 0.25
0.28 0.31 0.12 0.07
0.42 0.44 0.22 0.12
0.27 0.34 0.15 0.05
0.19 0.37 0.13
0.20 0.42 0.16 0.03
MnO
12.9 15.8 12.1 14.4
14.2 14.6 11.1
15.3 9.48 49.6 17.5
43.6 6.14 45.6 16.9
23.8 10.5 49.8 18.4
33.0 9.30 50.2
43.1 7.60 49.3 18.3
MgO
11.0 12.5 9.9 24.5
12.0 12.3 10.5
0.22 0.00 0.01 12.9
0.70 0.00 0.01 24.3
0.24 0.00 0.01 12.5
0.27 0.00 0.01
0.20 0.00 0.01 13.0
CaO
1.00 2.72 2 0.34
2.58 2.87 2
0.05 0.00 0.00 2.15
0.01 0.00 0.00 0.09
0.06 0.00 0.00 2.40
0.03 0.00 0.00
0.04 0.00 0.00 1.8
Na2O
P 2O 5
0.01 2 2 2
0.10 0.23 2 2
0.30 0.34 2
0.03 2 2 2
0.04 2 2
0.02 ,0.01 0.00 2 0.00 2 0.24 2
0.01 ,0.01 0.00 2 0.00 2 0.00 2
,0.01 0.00 0.00 0.16
0.01 ,0.01 0.00 2 0.001 2
,0.01 ,0.01 0.00 2 0.00 2 0.43 2
K2O
99.4 98.1 101.8 100.8
99.8 97.4 100.6
100.7 100.4 100.5 97.5
100.8 99.8 100.5 100.2
99.5 99.2 100.6 97.1
100.0 100.1 100.2
99.8 99.7 100.7 97.1
Sumc
f
Includes 0.03 wt% SrO; mode has 1.55 wt% ilmenite Il0.84Hm0.13Gk0.03 Includes 0.03 wt% SrO, 0.11 wt% F, 0.03 wt% Cl g Includes 0.04 wt% SrOin HB and 0.06 wt% SrO in CP
e
0.01 0.03 0.02 0.04
0.02 0.03 0.01
0.24 0.24 0.33 0.10
0.32 0.28 0.31 0.04
0.22 0.18 0.32 0.07
0.25 0.16 0.33
0.28 0.15 0.31 0.09
NiO
2 34.9 63.6 1.5
2 92.2 6.3
2 84.6 13.6 1.7
2 2.7 94.2 3.0
2 64.9 33.2 1.9
2 41.1 56.7
2 12.4 86.0 1.5
Moded
Table 1 Composition of rock samples and constituent minerals from isomodal layers. (WR whole rock, SP spinel, OL olivine, HB hornblende, CP clinopyroxene, GA garnet)
354
355
of slabs C12-805Ac and -805Ad cut from the plateau zones in each of the isomodal bands, and lists averages of microprobe analyses of the three main Mg-Fe minerals in each band. Sample C12-805B was collected ¥2 m farther along the same chromitite unit. The &2 wt% accessory hornblende is a trivial Fe-Mg reservoir (&1% of the total), so its compositional shift can be ignored in treating the post-cumulus compositional evolution of chromite and olivine. Figure 3b shows tie lines linking plateau compositions of olivine and chromite within the three domains, which collectively span the maximum range of chromite: olivine modal ratios at field locality C12-805. Results The ion-exchange reaction is FeSi0.5O21Mg(Cr,Al)2O4 5MgSi0.5O21Fe(Cr,Al)2O4. During ion exchange that accompanies cooling, olivine becomes more magnesian and the spinel less so, causing each tie line to rotate to its measured orientation about a pivot corresponding to the weighted average Mg' of spinel1olivine (weighted using a two-oxygen olivine mole). The pivots in Fig. 3b are collinear, so all three modal domains once shared one tie line and have remained chemically closed as mineral compositions subsequently evolved to their present values. Linear regression of the spinel modal proportion 1003Spy(Ol1Sp) versus weighted mean Mg' (through the pivots) yields r250.9995 for the restored tie line having abscissa intercepts spinel Mg'552.4+0.4 and olivine Mg'588.9+0.1 (1s). These are spinel and olivine compositions formerly shared by all three sampled modal domains at locality C12-805. The restored spinel’s mean Mg'552.4 is substantially more magnesian than the highest preserved Mg'548.3 encountered amongst the 72 electron-microprobe spot analyses. At a P of 7.5 kbar (from the equilibrium 2Mg2SiO4 1CaAl2Si2O85CaMgSi2O61Mg2Si2O61MgAl2O4 in green-spinel-bearing olivine-gabbronorites precipitated above the dunite in the same Thurly Gah cyclic unit), the restored olivine and spinel compositions and experimental calibrations of the olivine-spinel geothermometer at high pressure (O’Neill and Wall 1987; Ballhaus et al. 1991) give 8438 C as the effective T of closure to exchange amongst the sampled domains by “long-range” melt advection. Table 1 and Fig. 4b show another pair of conjugate bands, C10-100C (chromitite) and C10-100D (dunite), from a different level in the same Cu-Pt-Au orebody. These yield the reconstructed compositions olivine Mg'584.9 and spinel Mg'546.8 that correspond to ¥8848 C at that locality. Conclusions from case 2 Temperatures in the range 835–8908 C for apparent closure to interstitial melt migration through the ore deposit resemble Ts recovered in late hornblende-diorites high in
Fig. 4 a A 7-cm-long sample of banded garnet-clinopyroxene hornblendite shows locations of analysed slices J2-6800G along a pale, garnet-rich layer, and of J2-6800H from a hornblende-rich layer. Hornblende is varied shades of grey. Black disseminations are ilmenite. b Measured values (Table 1) of the Mg' of garnet and hornblende and of bulk domain J2-6800H (a hornblende-rich layer; open circle symbols) and bulk domain J2-6800G (a garnetrich layer; squares) are plotted on the abscissa. The modal proportion of hornblende in the hornblende (Hb)1garnet (Ga) mixture is represented on the ordinate axis. Solid tie lines link the measured mineral and domain compositions in each sample. As described in Fig. 1, a dashed tie line is drawn through the two bulk compositions to find the original hornblende and garnet compositions at the upper and lower abscissa intercepts (diamond symbols): hornblende Mg'H576.4+1.2 (1s) and garnet Mg'G559.4+1.4. The temperature at which the two modal domains became closed to mutual chemical exchange is TC5896+98 C, which is obtained from the value of the garnet-hornblende Fe-Mg exchange G H H KD;(XG FeyXMg)y(XFeyXMg)52.217+0.055 (1s) and from the correlation of KD with temperature (Graham and Powell 1984). The standard error of the restored KD (Bevington 1969, p. 62) yields the sample-derived standard error (98 C) shown for TC; the additional T uncertainty of the thermometer’s calibration (Graham and Powell 1984) is omitted as extraneous to the diagram’s mineral restoration message. The retrograded temperatures 793 and 8358 C that pertain to measured garnet and hornblende compositions within each domain are calculated in the same way and shown on their respective tie lines
356
the differentiated intrusion. So these Ts seem to reflect cessation of ore-forming magmatic infiltration metasomatism near the hydrous but H2O-undersaturated solidus T appropriate for residual interstitial liquids of andesitic composition issuing from underlying gabbroid cumulates up through the sulphidic dunite (cf. Irvine 1980b). Case 2 has demonstrated the capability of the mineral-composition restoration technique to remove the effects of solid-state diffusional re-equilibration and thereby recover mineral compositions and crystallization temperatures produced by another process at an earlier evolutionary stage that terminated at temperatures dramatically hotter (843–8848 C) than palaeo-temperatures (range 623–7418 C) represented by the mineral compositions that presently occur in these deep-seated, slowly cooled plutonic rocks.
Case 3: geobarometric constraint on extreme Himalayan uplift and erosion Most geothermometers are based on crystal-lattice-conserving ion-exchange reactions that are kinetically faster and effectively cease later in a rock’s cooling history than sluggish lattice-reconstructive reactions that are the basis of most geobarometers (Frost and Chacko 1989). Case 3 illustrates the large errors in P and depth estimates that often arise by using retrograded crystal-compositional closure T, rather than rock-compositional closure T in geobarometry. The lowermost unit exposed in the upper-mantleylower-crustal section of the Kohistan island arc terrain, Pakistani Himalayas, is the Jijal Complex. It comprises an ultramafic series ¥4 km thick (mainly olivine and clinopyroxene) overlain by garnet-hornblende clinopyroxenite and a series of kindred rocks ¥7 km thick having plagioclase1quartz1garnet1clinopyroxene1hornblende, with intermittent ultramafic layers or lenses (Miller et al. 1991; Miller and Christensen 1994). At one locality in a mafic-to-ultramafic transition, the rock shows cm-scale banding and consists chiefly of garnet and pargasitic hornblende, with interstitial low-Na clinopyroxene, and no feldspar (Fig. 4a). Table 1 gives wet-chemical whole-rock analyses of a slab ca. 1 cm thick cut along a hornblende-rich band and another slab along an adjacent garnet-rich band. Table 1 also gives the mean compositions of bulk crystals (grain cores1rims) of garnet, hornblende, and pyroxene analysed by electron probe in each slab. Hornblende grains are relatively homogeneous (sMg'51.1%), but garnet rims and cores vary by up to 3.6% in Mg'. Restoration of diffusionally altered mineral compositions If each analysed band’s bulk chemical composition and proportions of Mg-Fe-exchanging minerals have not changed during cooling, the retrograde diffusive shift in
bulk-crystal garnet and hornblende Mg's can be assessed by material balance. The 1.5% ilmenite concentrated in the hornblende-rich band has a trivial Mg-Fe21exchange capacity with respect to retrograde perturbation of the garnet1hornblende Mg-Fe21 reservoir. Clinopyroxene present mainly in the garnet-rich band accounts for only 1.1% of that band’s whole-rock Mg1Fe, so pyroxene’s retrograde shift in Mg' likewise has trivial leverage on garnet-hornblende Mg' systematics. The material balance on Mg' is essentially a binary lever given by Mg'GH5Mg'H2YG(Mg'H 2Mg'G), wherein YG;GyG1H is the molar proportion of garnet in the garnet-hornblende mixture (ion-exchange-equivalent moles), and Mg'GH is the weighted mean of the hornblende Mg'H and garnet Mg'G (equivalent to the domain’s whole-rock composition projected from augite and ilmenite). Figure 4b represents this mass balance as tie lines in a coordinate frame like the previous cases. The restored garnet’s mean Mg'G559.4 is significantly higher than the most magnesian grain core composition, Mg'G558.3, preserved in the garnet-rich band, as sampled by 17 electron-microprobe spot analyses. Figure 4b’s recovered TC58968 C for modal-domain compositional closure is obtained from a P-insensitive ion-exchange equilibrium. It can be combined with a Psensitive equilibrium using restored mineral compositions to evaluate depth at the time the domains closed. Geobarometry using restored mineral compositions A P-sensitive equilibrium relating activities of molecular components of hornblende, augite, and garnet in domain J2-6800G is magnesiohornblende12 diopside5tremolite12y3 grossular11y3 pyrope. Thermodynamic activities of molecular components in hornblende are modelled by a Margules-type calibration of pairwise ion-interaction energies describing non-ideal mixing of different elements occupying the same kind of lattice site (Mäder et al. 1994). The restored hornblende has Mg'H576.4 and other element concentrations as given in Table 1. The restored composition yields the ratio of tremolite and magnesiohornblende activities as RT ln (atremyamhb)527.13 kJymol at 8968 C. The restored garnet composition has Mg'G559.4 and pyrope and grossular activities corresponding to RT ln apyro 5217.99 kJymol and RT ln agros5231.72 kJymol (Table 1; garnet mixing properties from Berman, 1990). The augite in Table 1 has Mg'590.1, but the restored augite in Fe-Mg exchange equilibrium with the garnet at 8968 C should have Mg'587.1 (Krogh 1988). That composition, in conjunction with published information on the non-ideal energies of mixing elements on lattice sites in augite (Newton et al. 1977; Holland et al. 1979; Holland 1990) yields a diopside activity corresponding to RT ln adiop522.24 kJymol. Stoichiometric multiples of the preceding activity terms sum to RT ln K5229.79 kJ for the equilibrium constant K. The enthalpy of formation of magnesiohornblende from the elements is taken from a
357
comprehensive compilation and thermodynamic analysis of relevant experimental phase equilibria (Loucks in press b). Other standard thermodynamic data for the amphibole-augite-garnet equilibrium are form Holland and Powell (1990). These give the reaction’s changes of Gibbs energy and volume as DrG01169K 1b541.61 kJ and 0 DrVTP 520.881 kJykbar. From the relation 2RT ln 0 KT,P5DrG01169K 1b1PDrVTP , a pressure of 13.4 kbar is obtained (palaeo-depth ¥48–50 km), in excellent agreement with independent evidence presented by Miller and Christensen (1994). [Note: Carroll and Wyllie (1989) synthesised the assemblage hornblende1garnet1augite in three experiments on an andesitic bulk composition at high P. Their analysed mineral compositions can be used with the procedure above to calculate an apparent equilibration P, for comparison with their experimental P, in order to test reliability of the geobarometer utilized above. For their runs 49 and 50 at 950+158 C and 15.0+0.75 kbar, calculated apparent Ps are 14.4 and 15.0 kbar, respectively – both within their stated uncertainty of experimental pressure. Their run 65 produced the same assemblage at 850+158 C, but typographical errors in the augite analysis preclude calculation of apparent P.] The retrograded mineral compositions in J2-6800G (Table 1) yield estimates of T57938 C (Fig. 4b) and P512.1 kbar that are much lower than the values calculated above by the same procedures from the restored mineral compositions. The pressure discrepancy of 1.3 kbar underestimates by ¥5 km the amount of crustal uplift and Cenozoic erosion along the Main Mantle Thrust at the southern boundary of the accreted Kohistan island-arc terrain in the western Himalayas. Sample locality J2-6800 occurs at a structural height of 4.9 km above the stratigraphic base of the Jijal Complex (Miller et al. 1991, Fig. 1) which now crops out at elevations to ¥4000 m above sea level, so 57–59 km of uplift is indicated in that region in the ¥53 Ma since the onset of Kohistan-derived, collision-related molasse sedimentation (Murree Formation; Garzanti et al. 1987) on the northwestern Indian Plate. Concluding remarks Plutonic igneous and metamorphic rocks commonly are mottled, streaked, or banded on the scale of centimetres to decimetres, due to variation in proportions of minerals in an assemblage held in common. By judicious selection and sampling of heterogeneities developed on the appropriate scale (or by iteratively combining or splitting intervals), a researcher can apply linear correlation statistics to chemically analysed samples in order to identify rock domains that formerly were in chemical-exchange equilibrium (via an aqueous fluid or melt phase), but have subsequently become mutually disequilibrated by local inter-granular solid-state diffusion on a scale smaller than sampled modal heterogeneities. In a suite of such samples, material balance – using measured pro-
portions and compositions of diffusionally altered minerals in individual domains – permits reconstruction of mineral compositions that formerly were shared. The reconstructed mineral compositions have broad-ranging applications to deciphering the physical and chemical conditions and processes by which the rocks were formed or deformed, and to characterizing dynamic phenomena by which materials of deep origin came to be accessible for sampling. Acknowledgements Work on the Chilas and Jijal complexes in Pakistan was supported by the U.S. National Science Foundation grant EAR 88-17353 to the author. Frank Spear, Roberta Rudnick, John Mavrogenes, Sue Kesson, David Ellis, Ian Campbell, Ray Binns and an anonymous reviewer are thanked for comments that clarified the presentation.
References Ballhaus C, Berry RF, Green DH (1991) High pressure experimental calibration of the olivine-orthopyroxene-spinel oxygen geobarometer: implications for the oxidation state of the upper mantle. Contrib Mineral Petrol 107: 27–40 Berman RG (1990) Mixing properties of Ca-Mg-Fe-Mn garnets. Am Mineral 75: 328–344 Bevington PR (1969) Data reduction and error analysis for the physical sciences. McGraw-Hill, New York Burnham CW (1979) The importance of volatile constituents. In: HS Yoder (ed) The evolution of the igneous rocks: fiftieth anniversary perspectives. Princeton Univ Press, Princeton, New Jersey, USA, pp 439–482 Carroll MR, Wyllie PJ (1989) Experimental phase relations in the system tonalite-peridotite-H2O at 15 kbar; implications for assimilation and differentiation processes near the crust-mantle boundary. J Petrol 30: 1351–1382 Conrad ME, Naslund HR (1989) Modally graded rhythmic layering in the Skaergaard intrusion. J Petrol 30: 251–269 Ferry JM (1992) Regional metamorphism of the Waits River Formation, eastern Vermont: delineation of a new type of giant metamorphic hydrothermal system. J Petrol 33: 45–94 Frost BR, Chacko T (1989) The granulite uncertainty principle: limitations on thermobarometry in granulites. J Geol 97: 435– 450 Garzanti E, Baud A, Mascle G (1987) Sedimentary record of the northward flight of India and its collision with Eurasia (Ladakh Himalaya, India). Geodin Acta 1: 297–312 Graham CM, Powell R (1984) A garnet-hornblende geothermometer: calibration, testing, and application to the Pelona Schist, Southern California. J Metamorphic Petrol 2: 13–31 Helz RT (1976) Phase relations of basalts in their melting ranges at PH2O55 kbar. II. Melt compositions. J Petrol 17: 139– 193 Holland TJB (1990) Activities of components in omphacitic solid solutions. Contrib Mineral Petrol 105: 446–453 Holland TJB, Powell R (1990) An enlarged and updated internally consistent thermodynamic dataset with uncertainties and correlations: the system K2O-Na2O-CaO-MgO-MnO-FeO-Fe2O3Al2O3-TiO2-SiO2-C-H2-O2. J Metamorphic Geol 8: 89–124 Holland TJB, Navrotsky A, Newton RC (1979) Thermodynamic parameters of CaMg-Si2O6-Mg2Si2O6 pyroxenes based on regular solution and cooperative disordering models. Contrib Mineral Petrol 69: 337–344 Holloway JR, Burnham CW (1972) Melting relations of basalt with equilibrium water pressure less than total pressure. J Petrol 13: 1–29 Hunter RH, Sparks RSJ (1987) The differentiation of the Skaergaard intrusion. Contrib Mineral Petrol 95: 451–461
358 Hunter RH, Sparks RSJ (1990) The differentiation of the Skaergaard intrusion: reply to A.R. McBirney and H.R. Naslund. Contrib Mineral Petrol 104: 248–254 Irvine TN (1965) Chromium spinel as a petrogenetic indicator. 1. Theory. Can J Earth Sci 2: 648–672 Irvine TN (1980a) Magmatic density currents and cumulus processes. Am J Sci 280A:1–58 Irvine TN (1980b) Magmatic infiltration metasomatism, doublediffusive fractional crystallization, and adcumulus growth in the Muskox intrusion and other layered intrusions. In: Hargraves RB (ed) Physics of magmatic processes. Princeton University Press, NJ, pp 325–383 Irvine TN (1980c) Observations on layering in the Skaergaard intrusion. Carnegie Inst Washington Geophys Lab Yearb 79: 257–262 Juster TC, Grove TL, Perfit MR (1989) Experimental constraints on the generation of Fe-Ti basalts, andesites, and rhyodacites at the Galapagos Spreading Center, 858W and 958W. J Geophys Res 94: 9251–9274 Kelemen PB (1990) Reaction between ultramafic rock and fractionating basaltic magma. I. Phase relations, the origin of the calc-alkaline magma series, and formation of discordant dunite. J Petrol 31: 51–98 Kerr RC, Tait SR (1986) Crystallization and compositional convection in a porous medium with application to layered igneous intrusions. J Geophys Res 91: 3591–3608 Krogh EJ (1988) The garnet-clinopyroxene Fe-Mg geothermometer – a reinterpretation of existing experimental data. Contrib Mineral Petrol 99: 44–48 Loucks RR (1996) A precise olivine-augite Mg-Fe-exchange geothermometer. Contrib Mineral Petrol (in press a) Loucks RR (1996) Thermodynamic properties of magnesiohornblende-tremolite-magnesiocummingtonite solutions. Am Mineral (in press b) Maaløe S (1976) The zones plagioclase of the Skaergaard intrusion, East Greenland. J Petrol 17: 398–419 Mäder UK, Percival JA, Berman RG (1994) Thermobarometry of garnet-clinopyroxene-hornblende granulites from the Kapuskasing structural zone. Can J Earth Sci 31: 1134–1145 Mathez EA (1995) Magmatic metasomatism and formation of the Merensky Reef, Bushveld Complex. Contrib Mineral Petrol 119: 277–286 McBirney AR (1985) Further considerations of double-diffusive stratification and layering in the Skaergaard intrusion. J Petrol 26: 993–1001 McBirney AR (1995) Mechanisms of differentiation in the Skaergaard intrusion. J Geol Soc London 152: 421–435 McBirney AR, Nakamura Y (1974) Differentiation of the Skaergaard intrusion. EOS 55: 459 McBirney AR, Naslund HR (1990) The differentiation of the Skaergaard intrusion: a discussion of Hunter and Sparks. Contrib Mineral Petrol 104: 235–247 McBirney AR, Noyes RM (1979) Crystallization and layering in the Skaergaard intrusion. J Petrol 20: 487–554 Miller DJ, Christensen NI (1994) Seismic signature and goechemistry of an island arc: a multidisciplinary study of the Kohistan accreted terrane, northern Pakistan. J Geophys Res 99: 11623– 11642 Miller DJ, Loucks RR, Ashraf M (1991) Platinum-group element mineralization in the Jijal ultramafic-mafic complex, Pakistani Himalayas. Econ Geol 86: 1093–1102 Morse SA (1986) Convection in aid of adcumulus growth. J Petrol 27: 1183–1214 Morse SA, Lindsley DH, Williams RJ (1980) Concerning intensive parameters in the Skaergaard intrusion. Am J Sci 280A:159–170 Naldrett AJ (1989) Magmatic sulfide deposits. Clarendon Press, New York
Newton RC, Charlu TV, Kleppa OJ (1977) Thermochemistry of high pressure garnets and clinopyroxenes in the system CaOMgO-Al2O3-SiO2. Geochim Cosmochim Acta 41: 369–377 O’Neill HSC, Wall VJ (1987) The olivine-orthopyroxene-spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth’s upper mantle. J Petrol 28: 1169– 1191 Rudashevsky NS, Avdontsev SN, Dneprovskaya MB (1992) Evolution of PGE mineralization in hortonolitic dunites of the Mooihoek and Onverwacht pipes, Bushveld Complex. Mineral Petrol 47: 37–54 Selverstone J, Chamberlain CP (1990) Apparent isobaric cooling paths from granulites: two counter-examples from British Columbia and New Hampshire. Geology 18: 307–310 Synder D, Carmichael ISE, Wiebe RA (1993) Experimental study of liquid evolution in an Fe-rich, layered mafic intrusion: constraints of Fe-Ti oxide precipitation on the T-fO ❚ and T-q paths of tholeiitic magmas. Contrib Mineral Petrol 113: 73–86 Sonnenthal E (1992) Geochemistry of dendritic anorthosites and associated pegmatites in the Skaergaard intrusion, East Greenland: evidence for metasomatism by a chlorine-rich fluid. J Volcanol Geothermal Res 52: 209–230 Sparks RSJ (1988) Petrology and geochemistry of the Loch Ba ring-dyke, Mull (N.W. Scotland): an example of the extreme differentiation of tholeiitic magmas. Contrib Mineral Petrol 100: 446–461 Sparks RSJ, Huppert HE, Koyaguchi T, Hallworth MA (1993) Origin of modal and rhythmic igneous layering by sedimentation in a convecting magma chamber. Nature 361: 246–249 Spulber SD, Rutherford MJ (1983) The origin of rhyolite and plagiogranite in oceanic crust: an experimental study. J Petrol 24: 1–25 Sun C-O, Williams RJ, Sun S-S (1974) Distribution coefficients of Eu and Sr for plagioclase-liquid and clinopyroxene-liquid equilibria in oceanic ridge basalt: an experimental study. Geochim Cosmochim Acta 38: 1415–1433 Taylor HP Jr, Forrester RW (1979) An oxygen and hydrogen isotope study of the Skaergaard intrusion and its country rocks: a description of a 55-million-year-old fossil hydrothermal system. J Petrol 20: 355–419 Till R (1974) Statistical methods for the earth scientist. Wiley and Sons, New York Toplis MJ, Carroll MR (1995) An experimental study of the influence of oxygen fugacity on Fe-Ti oxide stability, phase relations, and mineral-melt equilibria in ferro-basaltic systems. J Petrol 36: 1137–1170 Tuccillo ME, Essene EJ, van der Pluijm BA (1990) Growth and retrograde zoning in garnets from high-grade metapelites: implications for pressure-temperature paths. Geology 18: 839– 842 Wagner LR (1960) The major element variation of the layered series of the Skaergaard intrusion and re-estimation of the average composition of the hidden layered series and of the successive residual magmas. J Petrol 1: 364–398 Wager LR, Brown GM (1968) Layered igneous rocks. Freeman, New York Walther JV, Orville PM (1982) Volatile production and transport in regional metamorphism. Contrib Mineral Petrol 79: 252– 257 Williams RJ (1971) Reaction constants in the system Fe-MgOSiO2-O2: intensive parameters in the Skaergaard intrusion, East Greenland. Am J Sci 271: 132–146 Yardley BMD (1977) An empirical study of diffusion in garnet. Am Mineral 62: 793–800 Yardley BMD, Bottrell SH, Cliff RA (1991) Evidence for a regional-scale fluid loss event during mid-crustal metamorphism. Nature 349: 151–154
