Pattison et al., 2003 temperature granulite facies metamorphism

Pattison et al., 2003 temperature granulite facies metamorphism

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Unformatted text preview: JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 PAGES 867±900 2003 Temperatures of Granulite-facies Metamorphism: Constraints from Experimental Phase Equilibria and Thermobarometry Corrected for Retrograde Exchange DAVID R. M. PATTISON1 *, THOMAS CHACKO2 , JAMES FARQUHAR3 AND CHRISTOPHER R. M. MCFARLANE4 1 DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF CALGARY, CALGARY, AB, T2N 1N4, CANADA 2 DEPARTMENT OF EARTH AND ATMOSPHERIC SCIENCES, UNIVERSITY OF ALBERTA, EDMONTON, AB, T6G 2E3, CANADA 3 DEPARTMENT OF GEOLOGY AND EARTH SYSTEM SCIENCE INTERDISCIPLINARY CENTRE, UNIVERSITY OF MARYLAND, COLLEGE PARK, MD 20742, USA 4 DEPARTMENT OF GEOLOGICAL SCIENCES, THE UNIVERSITY OF TEXAS AT AUSTIN, AUSTIN, TX 78701, USA RECEIVED MAY 29, 2002; ACCEPTED NOVEMBER 18, 2002 This study assesses temperatures of formation of common granulites by combining experimental constraints on the P±T stability of granulite-facies mineral associations with a garnet± orthopyroxene (Grt±Opx) thermobarometry scheme based on Al-solubility in Opx, corrected for late Fe±Mg exchange. We applied this scheme to 414 granulites of mafic, intermediate and aluminous bulk compositions. Our findings suggest that granulites are much hotter than traditionally assumed and that the P±T conditions of the amphibolite±granulite transition portrayed in current petrology textbooks are significant underestimates by over 100 C. For aluminous and intermediate granulites, mean corrected temperatures based on our method are 890 Æ 17 and 841 Æ 11 C, respectively (uncertainties reported as 95% confidence limits on the mean), consistent with minimum temperatures for orthopyroxene production by fluid-absent partial melting in these bulk compositions. In contrast, mean temperatures based on Grt±Opx Fe±Mg exchange equilibria, using the same thermodynamic data, are 732 Æ 22 and 723 Æ 11 C, respectively, well below the minimum temperatures for Opx stability. For mafic granulites, the mean corrected temperature using our method is 816 Æ 12 C, similar to the mean temperature of 793 Æ 13 C from Fe±Mg exchange. Reasons for the differences between the mafic granulites and aluminous±intermediate granulites are unclear but may be due to the lower Al concentrations in Opx in the mafic rocks and possible deficiencies in the thermodynamic modelling of these low concentrations. We discuss a number of well-known granulite terrains in the context of our findings, including the Adirondacks, the Acadian granulites of New England, the incipient charnockites of southern India and Sri Lanka, and the Kerala Khondalite Belt. Our findings carry implications for thermotectonic models of granulite formation. A computer program to perform our thermobarometry calculations, RCLC, is available from the Journal of Petrology website at or from the authors at$pattison/drm_pattison-rclc.htm. *Corresponding author. Telephone: 403-220-3263. Fax: 403-2840074. E-mail: Journal of Petrology 44(5) # Oxford University Press 2003; all rights reserved. granulite-facies metamorphism; thermobarometry; garnet; orthopyroxene KEY WORDS: JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 Table 1: Diffusion coefficients and diffusion length scales INTRODUCTION Although thermobarometric studies of granulite terrains have been carried out for more than 30 years, the temperatures reported in many of these studies fundamentally disagree with fluid-absent experimental data delimiting the stability of granulite-facies mineral assemblages. This discrepancy can be attributed to one of two primary causes (e.g. Nair & Chacko, 2002): (1) granulite metamorphism is triggered by the influx of low-aH2O fluids, which allows anhydrous minerals such as orthopyroxene to form at temperatures well below those required under fluid-absent conditions; or (2) peak temperatures of granulite metamorphism are higher than indicated by commonly used granulite geothermometers because of compositional re-equilibration of minerals on cooling. Although low-aH2O fluid infiltration has been inferred in a few granulite terrains (e.g. Janardhan et al., 1982; Chacko et al., 1987; Pattison, 1991; Bradshaw et al., 1989b; Todd & Evans, 1994), most granulites are thought to have formed under fluid-absent conditions (e.g. Harley, 1989; Thompson, 1990; Valley et al., 1990; Clemens, 1992; Spear et al., 1999). Thus, the temperature discrepancy noted above must be attributed to the inability of the minerals to retain their peak-temperature compositions. Resetting of temperatures is related to the diffusional mobility of elements on which conventional granulite thermobarometry is based. Table 1 lists diffusion coefficients and length-scales of diffusion (diffusion penetration distances) for some major elements in common minerals, calculated for 850 C and 10 Ma. The diffusion penetration distances vary over six orders of magnitude in a predictable fashion according to crystal chemical constraints (Dowty, 1980), and suggest that there is little reason to expect simultaneous closure of elements at any point in a high-grade rock's P±T history, let alone at peak P±T conditions. Apart from the obvious implications for temperature, variable closure temperature also affects pressure estimates because these depend on the temperature [the `feedback' effect of Harley (1989)]. Many granulite temperature estimates in the literature are based on Fe±Mg fractionation between coexisting phases such as garnet, biotite, cordierite, orthopyroxene and clinopyroxene [Grt, Bt, Crd, Opx, Cpx; abbreviations of Kretz (1983)]. Several studies have provided evidence that the closure temperature for Fe±Mg exchange between these phases is below peak granulite-facies temperatures (e.g. O'Hara, 1977; Lasaga, 1983; Frost & Chacko, 1989; Harley, 1989; Spear & Florence, 1992; Fitzsimons & Harley, 1994; Pattison & Bgin, 1994a, 1994b; Chacko et al., e 1996). Pattison & Bgin (1994a, 1994b) suggested that, e MAY 2003 Diffusion penetration 10 kbar Reference log10 D (cm2 /s) at 850 C, Element and mineral distance* (mm), 10 Myr À23.0 À19.3 0.001 0.08 À19.2 À16.2 0.09 3.00 Al in plagioclase 1 Al in clinopyroxene 2 Ca±Mg in clinopyroxene 3 Fe±Mg in garnet 4 Fe±Mg in orthopyroxene 5 Fe±Mg in cordierite 6 À15.6 À13.5 Na±K in plagioclase 7 À12.1 5.93 63.1 330.2 *Diffusion penetration distance $(4Dt)0Á5, where D is diffusion coefficient and t is time. References: 1, Grove et al. (1984); 2, Jaoul et al. (1991); 3, Brady & McAllister (1983); 4, Chakraborty & Ganguly (1991); 5, Ganguly & Tazzoli (1994); 6, Lasaga et al. (1977); 7, Foland (1974), Brady & Yund (1983) and Yund (1986). except for touching Fe±Mg minerals that continue to exchange Fe and Mg locally down to low temperatures, closure of intergranular Fe±Mg exchange may be controlled primarily by loss of a pervasive intergranular transport medium (probably melt), rather than by different rates of volume diffusion. In addition to these closure-related issues, a further complication arises if retrograde net-transfer reactions have taken place (Spear & Florence, 1992; Kohn & Spear, 2000). These can cause a variety of changes to the mineral compositions that can lead to erroneously high or low temperature estimates depending on what parts of the minerals are analyzed. Focusing on the primary problem of closure temperature, two important questions are: (1) is the amount of retrograde Fe±Mg exchange below peak granulite conditions generally modest (e.g. 550 C) or substantial (e.g. 4100 C)? (2) If it is substantial, what methods can be used to see through its effects to infer peak granulite P±T conditions? The first question has been addressed by kinetic modelling studies (e.g. Chakraborty & Ganguly, 1991; Spear & Florence, 1992; Ganguly & Tirone, 1999), but the results have been equivocal owing to uncertainties in the diffusion data and their extrapolation to natural conditions (e.g. Pattison & Bgin, 1994a, p. 388). The approach in this e paper is to use constraints from experimental phase equilibria in concert with a geothermobarometry scheme based on Al-solubility in Opx in equilibrium with garnet, corrected for late Fe±Mg exchange, which we have applied to 414 Grt ‡ Opx ‡ Pl ‡ Qtz-bearing 868 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM rocks from 62 granulite terrains worldwide. Whereas this combined approach has been used by Harley (1998a, 1998b) in his studies of ultra-high-temperature metamorphism, our focus is on the less exotic but more widespread `ordinary' granulites of common mineralogy. Our results suggest that commonly reported temperatures of granulite-facies metamorphism are significant underestimates, on average by 4100 C. We discuss several well-known granulite terrains in the context of our findings, including the Adirondacks, the Acadian metamorphic high of Massachusetts, the incipient charnockites of southern India and Sri Lanka, and the Kerala Khondalite Belt. We then explore the implications of these higher temperatures for thermotectonic processes of granulite formation. RELATIONSHIP BETWEEN THERMOBAROMETRY, PHASE EQUILIBRIA AND a H 2 O melting reactions. The dehydration melting reactions shown in Fig. 1 were investigated in vapour-absent experiments, simulating conditions in natural (noninfiltrated) granulites in which any free fluid would have been consumed at the wet solidus down-grade of the dehydration melting reactions. The reaction positions shown in Fig. 1 are for the lowest-temperature occurrence of the diagnostic mineral or mineral association (e.g. Opx), and hence provide lower limits on possible P±T conditions for rocks containing these minerals. To keep the diagrams legible, experimental brackets have been omitted and lines have been drawn by eye approximately through the midpoints of the limiting experiments of the respective studies. At a given pressure, the typical temperature interval between limiting experiments is 25±50 C, so that individual lines should be considered as bands of width Æ25 C or so. Mafic granulites Important first-order constraints on the peak P±T conditions of granulites are provided by a large body of experimental data on the P±T stability of granulitefacies mineral associations (see below). In the absence of low-aH2O fluid infiltration, the characteristic anhydrous mineralogy of granulites (e.g. Grt, Opx, Cpx, sillimanite) results from supersolidus, vapourabsent dehydration melting reactions that consume muscovite, biotite and/or hornblende and produce a melt phase in addition to the anhydrous minerals (e.g. Brown & Fyfe, 1970; Thompson, 1982). In dehydration melting reactions, aH2O is internally controlled by the melting reaction (Powell, 1983; Clemens & Vielzeuf, 1987; Thompson, 1990) and is therefore a function of the bulk composition (mineral assemblage plus melt) and P±T conditions. Recognizing that aH2O is a dependent variable in the absence of fluid infiltration provides the justification for evaluating thermobarometry results against the results from experimental fluidabsent phase equilibrium studies. EXPERIMENTAL CONSTRAINTS ON P±T STABILITY OF GRANULITEFACIES MINERAL ASSOCIATIONS Figure 1 shows experimental and thermodynamically predicted positions of reactions that limit the P±T stability of some key granulite-facies mineral assemblages for the three most common bulk compositions, here defined as mafic, intermediate and aluminous. The relevant wet solidi for each composition are also shown. Except at low pressure (5$3 kbar), the reactions occur above the wet solidi and so are dehydration Mafic granulites (Fig. 1a) are characterized by Cpxand/or Hbl-bearing Opx ‡ Pl Æ Grt Æ Bt Æ Kfs Æ Qtz mineral assemblages, and are broadly basaltic in composition. The general reaction that first introduces Opx to Hbl ‡ Pl Æ Qtz Æ Cpx Æ Grt-bearing mafic amphibolites below $10 kbar is Hbl ‡ Qtz Æ Grt ˆ Opx ‡ Cpx Æ Pl ‡ LX …1† At pressures below the solidus, this reaction produces a hydrous vapour phase rather than a melt phase [(1H ) in Fig. 1a]. At pressures above $10 kbar, an analogous reaction introduces Grt ‡ Cpx without Opx: Hbl ‡ Pl ‡ Qtz ˆ Grt ‡ Cpx ‡ L …2† (Pattison, 2003). Quartz-free reactions are also likely, but quartz-bearing reactions are the ones of importance to this study because they occur at lower temperatures (Spear, 1981; Hartel & Pattison, 1996). Figure 1a shows the experimental constraints on reactions (1) and (2), with reaction (1) being of most interest to this study because it produces Opx. The various experimental studies on reaction (1) are in rather good agreement (850 Æ 50 C over the pressure range 3± 10 kbar, with no obvious pressure dependence), even though a fairly wide range of starting mineral compositions is represented [see table 1 of Pattison (2003)]. Intermediate granulites Intermediate granulites (Fig. 1b) are characterized by Cpx- and Hbl-free Opx ‡ Pl Æ Grt Æ Bt Æ Kfs Æ Qtz mineral assemblages, and are broadly representative of metamorphosed psammites, semipelites and intermediate±felsic igneous rocks in which the major hydrous mineral is biotite. The general reaction that 869 JOURNAL OF PETROLOGY VOLUME 44 870 NUMBER 5 MAY 2003 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM combinations of the aluminous minerals Grt, Crd, Al2SiO5 (sillimanite, kyanite or andalusite), spinel Bt ‡ Qtz Æ Pl ˆ Opx ‡ L Æ Grt Æ Crd Æ KfsX …3† (Spl), sapphirine (Spr), corundum (Crn) and osumilite (Os), in addition to Opx and/or Bt (e.g. Harley, 1989, At pressures below the solidus, this reaction produces a 1998a). Some workers consider any rock with stable hydrous vapour phase rather than a melt phase. Figure Kfs ‡ Al SiO to be representative of granulite facies. 2 5 1b shows the experimental data on reactions (3) and The first appearance of this association is by the H (3 ) in addition to two pertinent wet solidi (one invol- general reaction ving K-feldspar and the other not). A wide variety of starting compositions for reaction (3) is represented in Ms ‡ Qtz Æ Pl ˆ Al2 SiO5 ‡ L Æ Kfs …4† Fig. 1b, ranging from end-member phlogopite ‡ quartz to mixtures of different types of natural Bt ‡ whose position is shown in Fig. 1c. At pressures below Pl ‡ Qtz of widely varying Pl composition and Mg/ the solidus, this reaction [(4H ) in Fig. 1c] produces a (Mg ‡ Fe), Ti and F content of biotite. Despite the hydrous vapour phase rather than a melt phase. The wide compositional range, most of the experiments fall experiments of Pati~o-Douce & Harris (1998) using n in an interval ranging from $800 C at 1 kbar to 900 C natural starting materials place reaction (4) rather at 15 kbar. Vielzeuf & Clemens (1992) discussed the higher than the experiments of Pto (1976) and the e lower temperatures found by Peterson & Newton thermodynamically modelled positions of Spear et al. (1989). The two studies showing the highest-tempera- (1999) and White et al. (2001). ture location of reaction (3) (Peterson et al., 1991; Nair Comparing the position of reaction (4) in Fig. 1c & Chacko, 2002) used starting biotite compositions with reactions (1) and (3) in Fig. 1a and b, the first containing significant F and/or Ti, such as found in appearance of Kfs ‡ Al2SiO5 occurs 100±150 C below natural biotites from the amphibolite±granulite transi- the first appearance of Opx in mafic and intermediate tion. The two thermodynamically predicted positions compositions. In regional and contact metamorphic of reaction (3) in Fig. 1b (Spear et al., 1999; White et al., sequences in which the first development of Kfs ‡ 2001) are calculated for F- and Ti-free biotite. These Al2SiO5 in metapelites and the first development of occur at lower temperature and show a stronger pres- Opx in mafic and intermediate composition can both sure dependence than the experiments. be mapped, the former is invariably significantly The position of the reaction curve for dry melting of down-grade of the latter (e.g. Adirondacks: De alkali feldspar solid solution ‡ quartz in the NaAl- Waard, 1969; Bohlen et al., 1985; New England: Si3O8±KAlSi3O8±SiO2 system (Holtz et al., 2001) is Schumacher et al., 1990a; Ballachulish: Pattison & also shown. The position of this curve is a minimum Harte, 1991, 1997; East Ontario Grenville, Carmichael for natural feldspars because addition of Ca results in an in Davidson et al., 1990; Broken Hill: Binns, 1964). up-temperature displacement (e.g. Johannes, 1978). We therefore prefer to define the amphibolite± granulite transition in metapelites by mineral associations that develop closer (both spatially in the field Aluminous granulites and, by extension, in pressure and temperature) to Aluminous granulites (Fig. 1c and d) represent meta- the first appearance of Opx in mafic and intermediate morphosed pelitic bulk compositions and may contain compositions, namely, Grt ‡ Crd ‡ Kfs (at pressures first introduces Opx to these rocks is Fig. 1. (opposite) Experimental and thermodynamically predicted positions of reactions that limit the stability of granulite-facies mineral associations. Numbered reactions are discussed in the text. Reactions (1), (3) and (4) are dehydration melting reactions that occur above the relevant wet solidi, whereas reactions (1H ), (3H ) and (4H ) are the corresponding dehydration reactions that occur below the relevant wet solidi at lower pressure. V, hydrous vapour; L, silicate liquid. Other abbreviations from Kretz (1983). (a) Mafic granulites. B 69, Binns (1969). BL 91, Beard & Lofgren (1991). C 88, Conrad et al. (1988). CW 67, Choudhuri & Winkler (1967). ET 86, Ellis & Thompson (1986). J 78, Johannes (1978). NC 00a, Nair & Chacko (2000), Three Valley Gap composition; NC 00b, Nair & Chacko (2000), Kapuskasing composition. P 68, Piwinskii (1968), tonalite 101. PB 95, Pati~o-Douce & Beard (1995). R 91, Rushmer (1991). R 93, Rushmer (1993). S 81, Spear (1981). SD 94, n Senn & Dunn (1994). WN 91, Winther & Newton (1991). The position of the reaction Grt ‡ Cpx ‡ Qtz ˆ Opx ‡ Pl is dotted in approximately, consistent with thermodynamic modelling. (b) Intermediate granulites. BBWC 83, Bohlen et al. (1983b). HBFJ 01, Holtz et al. (2001); Na±K Fsp (ss) is alkali feldspar solid solution. J 78, Johannes (1978). J 84, Johannes (1984). LJT 64, Luth et al. (1964). NC 02, Nair & Chacko (2002). P 68, Piwinskii (1968), granodiorite 102. PB 95, Pati~o-Douce & Beard (1995). PB 96, Pati~o-Douce & Beard (1996). PN 89, n n Peterson & Newton (1989). PCK 91, Peterson et al. (1991). SCD 97 gw, mp, Stevens et al. (1997), greywacke and metapelite compositions. SKC 99, Spear et al. (1999). TB 58, Tuttle & Bowen (1958). VC, Vielzeuf & Clemens (1992). VM, Vielzeuf & Montel (1994). W 76, Wood (1976). WPH 01, White et al. (2001). The PCK 91 and NC 03 experiments are for biotite compositions enriched in F and/or Ti. (c) Aluminous granulites. CJ 74, Chatterjee & Johannes (1974). J 84, Johannes (1984). LJT 64, Luth et al. (1964). P 68, Piwinskii (1968), granite 104. P 76, Pto (1976). P 92, Pattison (1992). PH 98 ms, mbs, Pati~o-Douce & Harris (1998); muscovite schist and muscovite±biotite-schist composie n tions. SCD 97 mp, Stevens et al. (1997), metapelite composition. SKC 99, Spear et al. (1999). TB 58, Tuttle & Bowen (1958). WPH 01, White et al. (2001). WPHW 00, White et al. (2000). (d) High-grade aluminous granulites. Continuous lines, synthesis of results from (b) and (c); dotted lines, high-grade FMASH equilibria from Harley (1998a). 871 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Fig. 2. (a) Location of the amphibolite±granulite transition from current textbooks compared with the granulite facies-limiting reactions from Fig. 1. Reactions are numbered as in the text and Fig. 1. BT, Blatt & Tracy (1996). BF94, Bucher & Frey (1994). M94, Miyashiro (1994). P90, Philpotts (1990). S93, Spear (1993). T81, Turner (1981). Y89, Yardley (1989). (b) Comparison of the granulite facies-limiting reactions from Fig. 1 with the granulite P±T estimates from table 1 of Harley (1989), Bohlen's (1987) domain of common granulite P±T conditions, and Harley's (1998a) domain of ultra-high-temperature metamorphism. below $9 kbar) or Opx ‡ Al2SiO5 (at pressures above $9 kbar). The lowest-grade reactions by which these assemblages develop are, respectively, Bt ‡ Sil ‡ Qtz Æ Pl ˆ Grt ‡ Crd ‡ L Æ Kfs …5† and Bt ‡ Grt ‡ Qtz Æ Pl ˆ Opx ‡ Al2 SiO5 ‡ L Æ KfsX …6† There are only a few experimental data on reaction (5) in Pl-bearing systems (Stevens et al., 1997) (Fig. 1c). Carrington & Harley (1995) provided constraints on reaction (5) in the Pl-free system that formed the basis for the two thermodynamically predicted positions of reaction (5) for Pl-bearing systems in Fig. 1c (Spear et al., 1999; White et al., 2001). The thermodynamically predicted curves are in close agreement but have a shallower slope than indicated by the experiments. Higher-grade reactions in aluminous granulites involve the minerals or mineral associations sapphirine (Spr), osumilite (Os) and spinel (Spl) ‡ Qtz, with Spr ‡ Qtz indicating particularly high temperatures in excess of 1000 C (e.g. Hensen, 1971; Harley, 1998a, 1998b). Figure 1d shows the estimated positions of some of the key quartz-bearing reactions in the Fe± Mg±Al2O3±SiO2±H2O (FMASH) system and their links to reactions (3), (5) and (6). For simplicity, Qtzabsent equilibria and those involving Os have been omitted, even though these are important in a full evaluation of the phase equilibria of `ultra-hightemperature' metamorphism (Harley, 1998a, 1998b, and references therein). The positions of most of the high-grade FMASH reactions in Fig. 1d are approximate owing to limited and sometimes conflicting experimental data and the strong effect of the water content of cordierite (Harley, 1998a, 1998b). In addition, the stability of Spl ‡ Qtz is dependent on Zn content of Spl and on fO2 (Harley, 1998a, 1998b). Nevertheless, these mineral associations give useful first-order indications of especially high-grade conditions. Comparison of experimental data with the amphibolite±granulite transition in textbooks and P±T estimates from thermobarometry Figure 2a compares the positions of the granulitefacies-limiting reactions for mafic, intermediate and aluminous compositions [reactions (1), (3) and (5)] 872 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM with the amphibolite±granulite boundary from a number of widely used petrology textbooks. The positions of reactions (1), (3) and (5) in Fig. 2 (grey bands) represent the averaged positions from Fig. 1. If the experimental data are accurate simulations of natural granulite-forming reactions, it appears that commonly accepted P±T conditions of the amphibolite±granulite transition are significant underestimates, by 4100 C. The textbook positions are generally close to reaction (4) (minimum stability of Kfs ‡ Al2SiO5; see Fig. 2a), although whether this is just a coincidence is unknown. Most of the textbooks define the amphibolite±granulite transition by the incoming of Opx in common mafic and intermediate bulk compositions, and place the incoming of Al2SiO5 ‡ Kfs by reaction (4) in the upper amphibolite facies as discussed above. Figure 2b compares the position of reactions (1), (3), (4) and (5) with the array of granulite-facies P± T estimates from table 1 of Harley (1989), Bohlen's (1987) P±T box of `common' granulite P±T conditions, and Harley's (1998a, 1998b) domain of ultrahigh-temperature metamorphism. Most of the P±T results are derived from conventional thermobarometry. Assigning meaningful uncertainties to the thermobarometry results is virtually impossible owing to the unknown effects of retrograde cation exchange during cooling and the variety of methods used. Arbitrary `ballpark' uncertainties of Æ50 C and Æ1 kbar are commonly used (e.g. Harley, 1989). The majority of these P±T estimates fall below the incoming of Opx [reactions (1) and (3)]. Three possibilities to explain this discrepancy are: (1) the experimental data do not accurately simulate the P± T conditions of the reactions in nature, tending to overestimate the minimum conditions required to develop the characteristic granulite-facies mineral associations; (2) the P±T conditions have been generally underestimated; (3) low-aH2O fluid infiltration is a much more widespread granulite-forming process than previously thought. Accepting the view that fluid infiltration did not occur in most cases, then one of the first two explanations or some combination of them seems most likely. Whereas retrograde cation exchange is perhaps the most obvious possibility, sluggish reaction kinetics of vapour-absent experiments in the crucial temperature range 700±800 C might result in no apparent reaction in experimental run times even when, for the same P±T conditions but at geological time scales, reaction might have proceeded. A few of the experimental studies shown in Fig. 1 addressed this possibility by reversing their experiments (e.g. Peterson & Newton, 1989; Peterson et al., 1991; Vielzeuf & Clemens, 1992; Nair & Chacko, 2002). GARNET±ORTHOPYROXENE Al-SOLUBULITY BASED THERMOBAROMETRY CORRECTED FOR LATE Fe±Mg EXCHANGE We attempt to address the above situation by developing and applying a thermobarometry scheme based on Al-solubility in Opx in equilibrium with Grt that corrects for the effects of late Fe±Mg exchange. Al concentrations in Opx are expected to be preserved from peak granulite conditions because of extremely slow diffusion of Al (e.g. Anovitz, 1991) (see data in Table 1). The rationale and approach of our scheme is not new, having been developed by Fitzsimons & Harley (1994) and Pattison & Bgin (1994a), and refined by e Chacko et al. (1996). The method reported in this paper is largely the same as reported in Chacko et al. (1996). Given knowledge of the pressure, the effects of late Fe±Mg exchange are corrected for by adjusting the Fe±Mg ratios of Grt and Opx to bring the Al-solubility and Fe±Mg exchange equilibria involving the two minerals into agreement (i.e. into internal equilibrium). The above workers found that the corrected Al-solubility temperatures were up to 200 C higher than those based on Fe±Mg exchange. We have refined the approach, using more recent thermodynamic data, and applied it to a large number of granulites worldwide to see if the predicted higher temperatures pertain to granulites in general. A key facet of our approach is the evaluation of the thermobarometry results against the limiting P±T conditions of the granulite-facies mineral associations described above. A computer program that performs the calculations, called RCLC (for `recalculation'), is described below and in Electronic Appendix A. The program and accompanying explanatory notes are available for downloading from the Journal of Petrology website at from the authors at http://$pattison/drm_pattison-rclc.htm. Thermodynamic data Our thermobarometry scheme uses the TWQ2.02b thermodynamic database, based on the thermodynamic data of Berman & Aranovich (1996) with modifications to incorporate the experiments of Aranovich & Berman (1997) on Al-solubility of Opx in equilibrium with Grt in the Fe-end-member system. Use of a single internally consistent database for all phases instead of a range of individually calibrated equilibria (e.g. Harley, 1998a, 1998b) eliminates problems of inconsistency between calibrations. The Grt±Opx experiments of Lee & Ganguly (1988) were part of the experimental dataset used by Berman & Aranovich 873 JOURNAL OF PETROLOGY VOLUME 44 Table 2: Difference between calculated Al-solubility temperatures using TWQ2.02b and experimental temperatures Study No. of NUMBER 5 MAY 2003 Table 3: Definition of thermodynamic system for RCLC Phases: System components: 5 Thermodynamic end-members: samples 8 Garnet 35 À13 23 À57 90 Nair & Chacko (2002) 6 À35 Alm 54 Orthopyroxene Ca3Al2Si3O12 Fs FeSiO3 En MgSiO3 Al-Op (1996) in the derivation of the thermodynamic data. As a test of the precision and accuracy of the TWQ2.02b database, Grt±Opx Al-solubility temperatures [based on equilibrium (8); see below] were calculated for 64 Grt ‡ Opx-bearing experiments in Fe±Mg Æ Cabearing systems from three studies (Harley, 1984; Lee & Ganguly, 1988; Nair & Chacko, 2002) (Table 2). Despite the variable quality of some of the experimental Al-solubility data [see discussion in Berman & Aranovich (1996) and Aranovich &Berman (1997)], the mean differences between the calculated temperatures and known experimental temperatures are 560 C, and in two of the studies are 535 C (overall mean for 64 samples is ±45 C). In all cases, the mean calculated temperatures are lower than the mean experimental temperatures, suggesting that TWQ-based P±T results if anything provide slight underestimates. Fe3 Al2 Si3 O12 ‡ 3MgSiO3 ˆ Mg3 Al2 Si3 O12 in Grt in Opx in Grt …7† SiO2 6 1 Alm ‡ 3 En ˆ 1 Prp ‡ 3 Fs (7) 1 Alm ˆ 3 Fs ‡ 1 Al-Op (8) 2 Alm ‡ 1 Grs ‡ 3 Qtz ˆ 6 Fs ‡ 3 An (9) 1 Prp ˆ 3 En ‡ 1 Al-Op (10) 2 Prp ‡ 1 Grs ‡ 3 Qtz ˆ 6 En ‡ 3 An (11) 1 Grs ‡ 2 Al-Op ‡ 3 Qtz ˆ 3 An Number of independent equilibria: (12) 3 Fe3 Al2 Si3 O12 ˆ 3FeSiO3 ‡ Al2 O3 in Grt in Opx …8† in Opx 2Fe3 Al2 Si3 O12 ‡ Ca3 Al2 Si3 O12 ‡ 3SiO2 in Grt in Grt ˆ 3CaAl2 Si2 O8 ‡ 6FeSiO3 in Pl in Opx Thermodynamic system The thermodynamic system used in RCLC is summarized in Table 3. The rocks of concern contain the Grt ‡ Opx ‡ Pl ‡ Qtz. Eight end-members in the five-component system Ca±Fe±Mg±Al±Si (CFMAS) account for the compositional variability in these phases. The Al-component of Opx is described by the `orthocorundum' end-member, Al2O3 (Aranovich & Berman, 1997), here given the abbreviation `Al-Op'. Six possible thermodynamic equilibria result, any three of which are independent. The three independent equilibria chosen are: CaAl2Si2O8 Qtz Quartz Total number of equilibria: Al2O3 An Plagioclase *Experiments in graphite capsules only. ‡ 3FeSiO3 in Opx Fe3Al2Si3O12 Mg3Al2Si3O12 Grs 38 Lee & Ganguly (1988) CaO±FeO±MgO±Al2O3±SiO2 Prp SD Harley (1984)* Garnet, Orthopyroxene, Plagioclase, Quartz T(TWQ) ± T(exp) Mean 4 Qtz …9† in which Opx is represented by a three-oxygen formula. Equilibria (7) and (8) both involve only Grt and Opx, equilibrium (7) being the Fe±Mg exchange between the two phases and equilibrium (8) being a net-transfer equilibrium controlling the Al-solubility of Opx in equilibrium with Grt. Equilibrium (9) involves all four phases and is the well-known net-transfer barometer in the Fe-end-member system (e.g. Bohlen et al., 1983b). Rationale for Fe±Mg correction Figure 3 shows two diagrams calculated using the TWQ software (Berman, 1991) for granulite sample PCFM-1 (mineral compositions listed in Table 4). All six equilibria from Table 3 are shown, with the three independent equilibria listed above shown by bold lines. Figure 3a is a typical result for a natural granulite in which the equilibria show considerable scatter, whereas 874 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Fig. 3. (a) TWQ plot (Berman, 1991) showing the typical spread in equilibria for a natural Grt ‡ Opx ‡ Pl ‡ Qtz granulite, sample PCFM-1 (mineral composition data in Table 4; modes of Grt:Opx:Bt:Crd ˆ 10:10:0:0). The phases, system components, phase components and equilibria are listed in Table 3. Of the six equilibria shown, only three are independent. The three chosen independent equilibria are indicated in bold. See text for further discussion of equilibria. (b) Same plot as in (a) except that the Fe±Mg ratios of Grt and Opx have been adjusted according to mass-balance constraints so that all equilibria intersect at a point. (See text for further discussion.) The light grey lines are the positions of the curves prior to adjustment. Table 4: Sample PCFM-1Ðmineral compositions used for calculations in Fig. 3 and Tables 5 and 6 Grt Opx Crd Bt Pl No. of oxygens: 12 6 18 11 8 Si 3.01 0.00 1.90 0.01 5.00 0.00 2.76 0.24 2.69 0.00 2.00 1.55 0.20 0.66 4.00 0.34 1.30 0.78 1.31 0.00 0.20 1.14 0.01 1.23 0.01 1.66 0.01 1.78 0.00 0.00 0.11 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.31 0.67 0.00 8.01 0.00 4.01 0.00 11.01 0.97 7.85 0.02 5.00 Ti Al Fe* Mn Mg Ca Na K Total à All Fe assumed to be Fe2 ‡. Fig. 3b shows the result for the same rock after adjusting the Fe±Mg ratios of the phases as described below. In Fig. 3a, the intersection of Grt±Opx Fe±Mg exchange equilibrium (7) with the Grt±Opx±Pl±Qtz barometer expression (9) (point A) represents the `conventional' P±T estimate of the rock (hereafter referred to as the uncorrected Fe±Mg P±T estimate). Aranovich & Berman (1997) advocated use of equilibrium (8), based on Al-solubility in Opx in the Fe-endmember system, as a good method to retrieve peak temperatures because it is relatively robust to late exchange of Fe±Mg. Intersection of equilibrium (8) with equilibrium (9) (point B in Fig. 3a; hereafter referred to as the uncorrected Fe±Al P±T estimate) gives a P±T estimate that is significantly higher than the `conventional' P±T estimate, consistent with the expected higher closure temperature of Al than of Fe± Mg. However, even this P±T estimate must be in error to some degree because it involves Fe end-members and so must have been affected by the late Fe±Mg exchange. The same comment applies to all of the equilibria in Table 3 and Fig. 3, each of which involves at least one Fe- or Mg-end-member or is affected by Fe±Mg ratio through activity±composition relations. RCLC corrects for the effects of late Fe±Mg exchange by adjusting the Fe±Mg ratios of the phases according to mass-balance constraints so that all the equilibria intersect at a point (point C in Fig. 3b). This intersection is hereafter referred to as the corrected Fe± Mg±Al P±T estimate. The implicit assumption is that this P±T point represents the condition of equilibrium before the onset of late Fe±Mg exchange (Fitzsimons & Harley, 1994; Pattison & Bgin, 1994a). As is the case e 875 JOURNAL OF PETROLOGY VOLUME 44 with all methods of thermobarometry, this assumption is impossible to prove and may not be correct in all situations (for example, garnet and orthopyroxene may have grown at different times and therefore P±T conditions, or may not have equilibrated with respect to Ca and Al). It nevertheless provides a rationale to correct for the obvious and in some cases substantial effects of late Fe±Mg exchange. The displacement of each equilibrium for the same amount of Fe±Mg change depends on the standard state free energy (mainly enthalpy) change and the effect of the Fe±Mg change on the equilibrium constant. Whereas equilibria (7), (10) and (11) show substantial displacement (compare Fig. 3a and b), equilibria (8) and (9) show more modest displacement. The corrected Fe±Mg±Al P±T estimate (point C in Fig. 3b) is about 1Á4 kbar and 170 C higher than the uncorrected Fe±Mg P±T intersection (point A in Fig. 3a), and about 70 C higher than the uncorrected Fe±Al P±T intersection (point B in Fig. 3a). Calculation method To bring the equilibria to convergence, Fe±Mg ratios of Grt and Opx are adjusted according to their modal abundance (assuming to begin with that Grt and Opx are the only two Fe±Mg phases in the rock). Implicit in this approach is that intergranular and intragranular transport of Fe and Mg is efficient enough to effect complete Fe±Mg exchange between the two phases. The sequence of steps to achieve convergence is as follows (refer to Fig. 3): (1) calculate the P±T position of the intersection of equilibria (8) and (9) (uncorrected Fe±Al P±T estimate; point B in Fig. 3a); (2) assuming constant bulk-rock composition and using the constraints provided by the modal abundance of Grt and Opx, change their Fe±Mg ratios so that Fe±Mg exchange equilibrium (7) coincides with point B; (3) recalculate a new position for the intersection of equilibria (8) and (9) using the new Fe±Mg ratios; (4) repeat several times to obtain convergence. In rocks that contain Fe±Mg phases in addition to Grt and Opx, such as Crd and Bt, Fe±Mg exchange is assumed to occur amongst all of the phases until Grt and Opx, the slowest Fe±Mg diffusers (Table 1), close to further exchange. Step (2) can therefore be expanded to include Crd and Bt by incorporating their modal abundances into the mass-balance equation, and simultaneously solving for each phases's Fe±Mg ratio so that each of the Grt±Opx, Grt±Bt and Grt±Crd Fe±Mg exchange equilibria coincide with point B. NUMBER 5 MAY 2003 In the situation that the uncorrected Fe±Mg P±T estimate is higher than the uncorrected Fe±Al P±T estimate (i.e. the positions of points A and B in Fig. 3a are reversed), the correction scheme works in the same way but results in a downward estimate in temperature. This situation probably indicates mineral compositions that are significantly out of equilibrium, perhaps as a result of the effects of retrograde net-transfer reactions, or of closure of one of Grt or Opx to Fe±Mg exchange and continued exchange of the other with Crd and/or Bt as the rock cooled. We can envisage no situation in which Al would exchange to lower temperatures than Fe±Mg. In these situations, the uncorrected Fe±Al P±T estimates are probably more reliable than the corrected Fe±Mg±Al P±T estimates. Effect of varying mineral modes Table 5 shows effects of varying mineral modes on the corrected Fe±Mg±Al P±T estimates, using the mineral compositions in Table 4. The effects are rather small, even with extreme variations in modes. Thus, a reasonable P±T estimate can be obtained by performing the convergence technique with Grt and Opx alone, regardless of the other Fe±Mg minerals in the rock. Furthermore, the simplifying assumption that the Fe± Mg adjustment can be taken up wholly by either Grt or Opx introduces relatively little error to the final calculated P±T conditions [see fig. 17 of Pattison & Bgin e (1994) for a graphical demonstration of this point]. These considerations become useful if modes or compositional information on Fe±Mg phases in the rock other than Grt and Opx are lacking. On the other hand, the amount by which each phase changes its Fe±Mg ratio varies significantly according to the mode, becoming most extreme for rocks with unequal modal proportions (Table 5). Sensitivity of the method Despite the obvious theoretical advantages of Grt±Opx Al-solubility-based thermobarometry, Al concentrations in Opx are typically rather low (1±12 wt % Al2O3, mostly in the range 1±5 wt %) and the method is sensitive to small changes in Al concentration. Table 6 shows changes in inferred P±T conditions for different Al concentrations in Opx, with all other mineral compositions held constant and the other elements in Opx varied proportionately (the mineral compositions are those listed in Table 4). For ease of comparison with other opx studies, XAl in Table 6 is the amount of octahedral Al in Opx assuming a six-oxygen formula [rather than the three-oxygen formula used in Table 3 and reactions (7)± (9)]. For uncorrected Fe±Al P±T estimates, the sensitivity increases as the absolute Al concentration drops, opx opx going from $20 C/0Á01 XAl at XAl ˆ 0Á13 ($6 wt % 876 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Table 5: Variation in recalculated temperatures with respect to modal variations for sample PCFM-1 (Table 4) P±T estimates before correction: P (kbar) T ( C) Uncorrected Grt±Opx Fe±Mg estimate [intersection 6.28 703 7.63 807 5.81 666 6.16 694 of equilibria (7) and (9); point A in Fig. 3a] Uncorrected Grt±Opx Fe±Al estimate [intersection of equilibria (8) and (9); point B in Fig. 3a] Intersection of Grt±Bt Fe±Mg exchange and equilibrium (9) Intersection of Grt±Crd Fe±Mg exchange and equilibrium (9) Mineral modes Corrected P±T Changes in Mg/(Mg ‡ Fe) Grt Opx Crd Bt P (kbar) T ( C) DGrt DOpx DCrd DBt 10 10 Ð Ð Ð Ð Ð Ð À0.02 À0.05 Ð Ð Ð Ð 0.03 0.01 Ð Ð 1 7.63 7.47 868 10 Ð Ð Ð Ð 1 10 Ð Ð Ð Ð 857 10 10 10 Ð Ð 7.76 7.69 0.05 0.04 0.00 À0.01 Ð Ð À0.06 10 10 Ð Ð 10 10 10 10 7.72 7.76 860 10 0.04 0.05 À0.01 0.00 Ð Ð À0.05 10 1 1 1 7.52 7.78 877 855 0.01 0.05 À0.04 0.00 À0.09 À0.05 À0.08 À0.03 7.93 7.90 844 0.08 0.07 0.02 0.02 À0.03 À0.03 À0.01 1 10 1 1 1 1 10 1 1 1 1 10 881 863 857 841 Ð Ð Ð Ð À0.04 À0.04 À0.01 DGrt, DOpx, DCrd, DBt: corrected Mg/(Mg ‡ Fe) ± initial Mg/(Mg ‡ Fe). Corrected P±T: point C in Fig. 3b. Table 6: Variation in uncorrected Fe±Al and corrected Fe±Mg±Al temperatures as a function of Al concentration in Opx for sample PCFM-1 (Table 4) XAl wt % Al2O3 Uncorrected Fe±Al Corrected Fe±Mg±Al P (kbar) T ( C) P (kbar) T ( C) 9.7 8.9 1050 8.1 7.3 920 750 480 0.15 0.13 6.8 5.9 9.6 8.9 920 0.11 0.09 5.0 4.1 8.1 7.3 840 0.07 0.05 3.2 2.2 6.4 5.5 730 670 6.4 5.5 0.03 1.3 4.3 570 4.3 880 790 XAl ˆ Al/2 for six-oxygen Opx. Uncorrected Fe±Al: point B in Fig. 3a. Corrected Fe±Mg±Al: point C in Fig. 3b. 877 990 840 650 JOURNAL OF PETROLOGY opx opx VOLUME 44 Al2O3) to $50 C/0Á01 XAl at XAl ˆ 0Á03 (1Á3 wt % Al2O3). For the corrected Fe±Mg±Al P±T estimates, the effects are stronger ($30 C and 60 C, respectively) because of the magnifying effect of the Fe±Mg correction. The above considerations show that relatively subtle factors become important to P±T estimation using Grt±Opx Al-solubility methods. These include analytical precision, assumptions about Al partitioning between the tetrahedral and octahedral sites in Opx, and the effects of minor elements such as Na ‡, Fe3 ‡, Cr3 ‡ and Ti4 ‡. Droop (1987) and Carson & Powell (1997) discussed strategies for assigning Al between the octahedral and tetrahedral sites of Opx and for determining Fe3 ‡ stoichiometrically from microprobe analyses. Regardless of sophistication, no approach can avoid the primary source of error, namely the accuracy and precision of Si and to a lesser degree Fe, Mg and Al. Ignoring initially the issue of Fe3 ‡ and other minor elements, opx calculating XAl assuming ideal Tschermak exchange vi [(Fe,Mg) ‡ Siiv ˆ Alvi ‡ Aliv ] gives rise to the scheme opx XAl ˆ AlM1 ˆ Altotal /2 (for a six-oxygen Opx formula unit). In our processing of a large body of data from the literature (see below), we have found that this approach results in significantly less scatter and fewer opx obviously erroneous values than calculating XAl by the opx M1 stricter site occupancy scheme XAl ˆ Al ˆ Altotal ± (2 ± Si). With regard to Fe3 ‡, in many cases the magnitude of inferred Fe3 ‡ is similar to or smaller than the combined analytical uncertainty on Si, Fe, Mg and Al, perhaps accounting for the sometimes unpredictable pattern of calculated Fe3 ‡ where there is no independent indication for significant variations in fO2 from oxides or other minerals in the rock. In addition, it is likely that the experimental Opx compositions from which the thermodynamic data were obtained contained at least a small component of Fe3 ‡ that is not accounted for (all Fe is assumed to be Fe2 ‡). Unless there is good evidence to indicate otherwise, we are therefore of the opinion that the dangers of overcorrection are as great as the dangers of undercorrection. Nevertheless, RCLC provides a opx variety of options for estimating XAl , including one 3‡ that takes account of Fe and other minor elements. For the above reasons and in the interests of uniformity (many studies did not analyse for minor elements and/or assumed all Fe was Fe2 ‡), we have followed Fitzsimons & Harley (1994), Aranovich & Berman (1997) and Berman & Bostock (1997) in assuming that opx XAl ˆ Al/2 (for a six-oxygen formula). This approach opx provides a maximum estimate for XAl , which in some cases, particularly high-Al metapelites, may result in an overestimate of calculated temperature (see Appendix A and discussion below on Kerala Khondalite Belt). For low concentrations of Al in Opx (e.g. less than $2 wt % Al2O3), such as in many mafic granulites, the accuracy of NUMBER 5 MAY 2003 the Al analysis may become an issue in itself (see discussion below). We therefore encourage reporting the standards used for Opx analysis along with the weight percent oxides. Scatter in pressure estimates and the use of RCLC-P Application of RCLC to a few natural datasets results in widely scattered pressure estimates beyond what is reasonable geologically, suggesting that in some samples the analyzed plagioclase and garnet compositions were substantially out of equilibrium. A modified version of RCLC, termed RCLC-P, allows pressure to be input as a known variable and calculates corrected Fe±Mg±Al temperatures by convergence of equilibria (7) and (8) at that pressure [essentially the method of Fitzsimons & Harley (1994) and Pattison & Bgin (1994a)]. RCLC-P e is also useful if the Grt ‡ Opx-bearing rocks have no plagioclase or the plagioclase compositions were not reported. For example, in estimating the temperatures of the Nain granulites of Berg (1977a, 1977b), a pressure of 5 kbar was assumed in the absence of information about Pl composition. Testing of the method We analyzed three datasets to test the above method. The first is that of Pattison & Bgin (1994a), which was e concerned with the variation with grain size of the composition of Grt and Opx in two regional granulites from the Minto granulite terrain of northern Quebec. Fe±Mg ratios of core compositions of Grt and Opx show increasing degrees of retrograde exchange as grain size decreases, whereas Al shows little pattern with respect to grain size. Figure 4 shows temperature vs grain size for the uncorrected Fe±Mg, uncorrected Fe±Al and corrected Fe±Mg±Al methods for the two samples. The strong dependence on grain size of the Fe±Mg exchange temperatures is absent from both of the Al-solubility-based methods. In addition, the absolute values of the two Al-solubility methods are in better agreement with the phase equilibrium constraints (estimated from Fig. 1b as 850 Æ 25 C at the $7 kbar pressure of the two samples), with the corrected Fe±Mg±Al temperatures showing the best agreement for B69E. The second dataset comes from the aureole of the Makhavinekh intrusion, Labrador (McFarlane et al., 2003). In these rocks, regional metamorphic garnet is progressively replaced by Opx ‡ Crd as the contact is approached. Figure 5 shows a profile through the aureole comparing uncorrected Fe±Mg, uncorrected Fe±Al and corrected Fe±Mg±Al temperatures (all assuming a uniform pressure of 5 kbar) as a function of distance from the contact. Whereas the Fe±Mg 878 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Fig. 4. Plots of temperature vs grain size for two granulites from Pattison & Bgin (1994a), using the three methods discussed in the text and e shown in Fig. 3. The grain size axis refers to grain size of garnet. The temperatures were calculated using the different garnets and a single Opx grain [see Pattison & Bgin (1994a) for further details]. `Minimum T of formation of Opx' is estimated from Fig. 1b for a pressure of 7 kbar, the e estimated pressure of formation of the two granulites. Fig. 5. Makhavinekh aureole, Labrador (McFarlane et al., 2003). Comparison of calculated temperatures with modelled thermal profiles [see McFarlane et al. (2003) for details of the modelling]. RCLC-P was used to calculate the temperatures at a fixed pressure of 5 kbar. exchange temperatures are around 600 C and show no variation within the aureole, the temperatures based on Al-solubility show a pattern of rising temperature as the contact is approached. The Al-solubility-based temperatures fit best with the thermal modelling of McFarlane et al. (2003). The third dataset comes from the ultra-hightemperature rocks of the Napier Complex, Enderby Land, Antarctica (Sheraton et al., 1980). Harley (1985) provided analyses of Grt ‡ Opx-bearing rocks that are interlayered with rocks containing coexisting Spr ‡ Qtz, mesoperthite and metamorphic pigeonite, 879 JOURNAL OF PETROLOGY VOLUME 44 together indicative of temperatures above $1000 C at 8±11 kbar (Harley, 1998a, 1998b; Fig. 1d). The mean uncorrected Fe±Mg temperature of 750 Æ 130 C (error reported as one standard deviation) is a gross underestimate. The uncorrected Fe±Al estimate (880 Æ 90 C) is better, whereas the corrected Fe±Mg±Al temperature (950 Æ 80 C) is close to the P±T conditions indicated by the assemblages. APPLICATION TO GRANULITES Based on the success of the three pilot studies, we applied RCLC to 414 Grt ‡ Opx-bearing rocks from 62 granulite terrains using analyses in the literature, and evaluated the results against the phase equilibrium constraints. An obstacle to applying this analysis to more granulite terrains was the relatively common absence of data on the Al content of Opx, particularly in papers written in the 1970s and 1980s, perhaps reflecting the view at that time that Fe±Mg exchange thermometry was all that was needed to estimate peak granulite temperatures. A more subtle factor influencing our results is our lack of control over, and many cases lack of knowledge of, the rationale used in selecting mineral analyses to be used for P±T calculations. Different analysis points from the same minerals in the same rock can result in P±T estimates that vary by several kilobars in pressure and hundreds of degrees Celsius in temperature (Spear & Florence, 1992; Pattison & Bgin, 1994a; Kohn & Spear, 2000). This e factor may account for a good deal of the scatter in the results described below. The Appendix lists the terrains, literature references, sample numbers and type of granulite (mafic, intermediate or aluminous) used in the compilation, the last being based on the mineralogical criteria given earlier. Of the 414 granulite samples examined, 80 are aluminous, 201 are intermediate and 133 are mafic. Electronic Appendix B lists key compositional parameters and calculated P±T results for each sample using the uncorrected Fe±Mg, uncorrected Fe±Al and corrected Fe± Mg±Al methods, and is available from the Journal of Petrology website or directly from the authors at http:// www.geo.ucalgary. ca/$pattison/drm_pattisonresearch.htm#publications. Table 7 summarizes the results of Electronic Appendix B by listing, for the three types of granulite, the mean, standard deviation, 95% confidence limit on the mean, and range of the compositional parameters and P±T estimates. What these statistics actually mean is somewhat unclear, given that the data almost certainly do not represent a Gaussian distribution arising from numerous independent, random perturbations. Systematic perturbations probably cause most of the scatter in the data, most of which result in a lowering of the temperature estimate (see discussion below). NUMBER 5 MAY 2003 We caution that the results in Table 7 and related figures below show broad trends only and are influenced by the samples and terrains included in the analysis. For example, the ultra-high-temperature Enderby Land terrain and the high-temperature regional contact aureoles around Nain, Labrador, are represented only by intermediate and aluminous samples. In addition, we cannot apply RCLC to Opx-free Grt ‡ Crd-bearing aluminous granulites or to Grt-free Opx‡Cpx-bearing mafic granulites. Nevertheless, we believe that our sample base is broad enough that the general trends are meaningful. opx Grt Figure 6a and b shows trends of XCa vs XAl , and Mg/ opx (Mg ‡ Fe)opx vs XAl , grouped according to type of opx Grt granulite. Compositional clustering in XCa vs XAl is obvious, with mafic granulites containing the most Carich Grt and Al-poor Opx, aluminous granulites containing the most Ca-poor Grt and Al-rich Opx, and intermediate granulites in between. There is a gap in Grt Grt composition between approximately XCa of 0Á12 and 0Á16, with values above the gap largely restricted to mafic mineral assemblages containing either or both of the calcic mafic phases Hbl and Cpx, and values below the gap restricted to intermediate and aluminous mineral assemblages lacking these phases. With respect opx to Mg/(Mg ‡ Fe)opx vs XAl , there is a trend to more Ferich compositions in the most Al-poor Opx compositions and a weak trend to higher Mg/(Mg ‡ Fe)opx as Opx becomesmorealuminous,similartothatobservedexperimentally (e.g. Harley, 1984; Lee & Ganguly, 1988). Figure 7 shows differences in temperature and pressure between the uncorrected Fe±Al and uncorrected Fe±Mg estimates (points B and A in Fig. 3) and between the corrected Fe±Mg±Al and uncorrected Fe±Mg estimates (points C and A in Fig. 3), grouped according to the three compositional types of granulite. Pressure differences between the two sets of estimates are largely the same (Table 7) and so only the difference between corrected Fe±Mg±Al and uncorrected Fe±Mg pressures is displayed. The temperature and pressure differences are strongly correlated because of the dependence of the pressure estimate on the temperature. Figure 8 illustrates compositional dependence of the results by plotting differences in temperature between the corrected Fe±Mg±Al and uncorrected Fe±Mg opx Grt estimates against XCa , XAl and Mg/(Mg ‡ Fe)Grt , grouped according to type of granulites. Figure 9a plots the mean results from Table 7 with respect to the granulite-facies-limiting reactions from Figs 1 and 2. Aluminous and intermediate granulites In aluminous and intermediate granulites, the mean difference between both the uncorrected and corrected Al-solubility-based P±T estimates and the uncorrected 880 Aluminous granulites (n ˆ 80) Mean SD Intermediate granulites (n ˆ 201) 95% CL* Range Mean SD 95% CL* Mafic granulites (n ˆ 133) Range Mean SD 95% CL* Range PATTISON et al. Table 7: Compositional parameters and P±T results for granulites; summary of Electronic Appendix A Compositional parameters Mg/(Mg ‡ Fe) Mg/(Mg ‡ Fe)opx X opx Al ( ˆ Al/2 for 6-O formula) 0.01 0.13 0.002 0.02 0.01±0.09 0.10±0.58 0.06 0.31 0.03 0.08 0.003 0.01 0.01±0.18 0.12±0.66 0.18 0.27 0.02 0.10 0.004 0.01 0.07±0.25 0.03±0.78 0.52 0.12 Grt 0.13 0.04 0.02 0.008 0.26±0.76 0.05±0.22 0.55 0.08 0.09 0.04 0.01 0.005 0.25±0.80 0.02±0.21 0.53 0.04 0.12 0.02 0.02 0.002 0.10±0.89 0.01±0.09 P±T estimates using different methods 881 Uncorrected Fe±Mg T ( C) (point A in Fig. 3a) P (kbar) Uncorrected Fe±Al T ( C) (point B in Fig. 3a) P (kbar) Corrected Fe±Mg±Al T ( C) (point C in Fig. 3b) P (kbar) 732 5.6 854 6.4 890 6.4 118 1.9 22 79 2.2 15 90 2.2 17 88 1.2 16 48 0.2 9 132 1.2 24 0.3 0.4 0.4 475±1156 2.0±14.1 723 672±1059 3.0±13.5 807 700±1093 3.0±13.5 841 À138 to 286 À1.7 to 3.8 84 À95 to 224 À1.1 to 0.9 34 À233 to 510 À1.2 to 4.3 118 6.7 7.9 8.0 98 1.6 88 1.4 98 1.5 11 421±987 1.5±11.2 793 629±1009 3.0±12.1 806 644±1077 3.0±12.2 816 0.2 8 0.1 À107 to 337 À1.5 to 5.3 13 4 0.02 À37 to 177 À0.3 to 1.0 0.2 10 0.2 11 92 1.6 13 70 1.7 10 84 1.7 12 75 1.3 11 0.2 9 0.0 35 0.1 9.6 9.8 9.9 0.2 0.2 0.3 518±1034 4.8±14.1 670±982 6.5±13.8 656±1081 6.6±14.9 P±T differences between methods Fe±Al À Fe±Mg T ( C) P (kbar) Fe±Mg À Al±Fe±Al T ( C) P (kbar) Fe±Mg À Al±Fe±Mg T ( C) P (kbar) 121 0.8 45 0.0 158 0.9 0.2 0.04 0.2 1.1 0.1 1.2 72 1.0 35 0.2 105 1.1 12 0.1 À144 to 481 À1.4 to 6.3 23 0.2 109 1.4 *95% confidence limit on mean ˆ t  standard error ˆ t  standard deviation/Hn, where t is Student's t statistic for 95% confidence interval. 0.2 5 0.02 16 0.2 À187 to 195 À3.6 to 3.4 À81 to 111 À0.3 to 0.5 À231 to 292 À3.9 to 4.0 TEMPERATURES OF GRANULITE-FACIES METAMORPHISM 0.03 0.31 X Grt Ca JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Fig. 6. Mineral composition trends in granulites, grouped by compositional type (aluminous, intermediate, mafic; see text for discussion). opx opx opx Grt (a) X Grt vs X Al . X Ca ˆ Ca/(Ca ‡ Fe ‡ Mn ‡ Mg). X Al ˆ Al/2 for a six-oxygen orthopyroxene formula. (b) Mg/(Mg ‡ Fe)opx vs XAl . Ca Fe±Mg P±T estimates is substantial: 121 Æ 16 and 158 Æ 24 C, respectively, for aluminous granulites, and 84 Æ 8 and 118 Æ 12 C, respectively, for intermediate granulites (uncertainties reported as 95% confidence limits on the mean). These values are of a similar magnitude to those found by Fitzsimons & Harley (1994), Pattison & Bgin (1994) and Chacko e et al. (1996). Pressure differences in both cases are $1 kbar higher. Regarding absolute temperatures, for aluminous granulites the mean uncorrected Fe±Mg temperature (732 Æ 22 C) is substantially below the minimum stability of Opx (Fig. 1), whereas the uncorrected Fe±Al and corrected Fe±Mg±Al temperatures (890 Æ 17 and 854 Æ 15 C) are consistent with Opx stability (Fig. 9a). The rather high mean corrected Fe±Mg±Al temperature for the aluminous granulites may reflect a combination of sample bias (several samples from ultra-high-temperature localities) and possible temperature overestimation for samples in which there is non-negligible Fe3 ‡. For intermediate granulites, the mean uncorrected Fe±Mg temperature (723 Æ 11 C) is substantially below the minimum stability of Opx (Fig. 1), whereas the mean corrected Fe±Mg±Al temperature (841 Æ 11 C) satisfies this constraint, with the mean uncorrected Fe±Al temperature (807 Æ 10 C) falling in between. We consider that the overall agreement between the Al-solubility-based thermobarometry and the experimental constraints on Opx stability lends support to both approaches, and points to retrograde exchange as the most likely 882 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Fig. 7. Temperature or pressure differences between methods vs absolute temperatures or pressures, grouped according to compositional type (aluminous, intermediate, mafic). (a, d, g) Uncorrected Fe±Al temperatures ± uncorrected Fe±Mg temperatures. (b, e, h) Corrected Fe±Mg± Al temperatures ± uncorrected Fe±Mg temperatures. (c, f, i) Corrected Fe±Mg±Al pressures ± uncorrected Fe±Mg pressures. (See text for discussion.) explanation for the discrepancy between the phase equilibria and geothermobarometry discussed in the Introduction. We caution that the P±T estimates in Table 7 and Fig. 7 show considerable scatter. This scatter is difficult to attribute to any single factor, arising from some combination of: differences in peak P±T conditions of the samples; varying degrees of retrograde Fe±Mg exchange; retrograde net-transfer reactions; mineral compositions that may not have been in equilibrium before late Fe±Mg exchange; analytical issues relating to the relatively small concentrations of Al in Opx; and 883 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Fig. 8. Temperature differences between corrected Fe±Mg±Al temperatures and uncorrected Fe±Mg temperatures vs mineral compositional opx Grt parameters, grouped according to compositional type (aluminous, intermediate, mafic). (a) X Al . (b) X Ca . (c) Mg/(Mg ‡ Fe)Grt . (See caption to Fig. 6 for definition of parameters.) possible deficiencies in the thermodynamic modelling of Al in Opx, especially at low concentrations (see section below on mafic granulites). Figure 8a shows that temperature differences between the corrected Fe±Mg±Al and uncorrected opx Fe±Mg estimates are positively correlated with XAl . This correlation is probably due to the fact that, for a given pressure, higher Opx Al contents indicate higher temperatures. The higher the peak temperature is above the closure temperature for Fe±Mg exchange, the greater the expected difference between the calculated Al-solubility and Fe±Mg exchange temperatures. 884 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Fig. 9. (a) Comparison of mean P±T results of mafic, intermediate and aluminous granulites containing Grt ‡ Opx with the granulite facieslimiting reactions from Fig. 1. *, uncorrected Fe±Mg method; *, corrected Fe±Mg±Al method. Reactions are numbered as in the text and Fig. 1. (b) Comparison of mean P±T results of individual granulite terrains with the limiting granulite facies-limiting reactions from Fig. 1. The label for each terrain is placed beside the corrected Fe±Mg±Al estimate (*). *, uncorrected Fe±Mg estimates. Abbreviations for the terrains are listed in Table 8. Corrected Fe±Mg±Al temperatures for the three contact metamorphic localities (SC, MA, NA) were calculated for fixed pressure using RCLC-P. This explanation probably also accounts for the greater spread in corrected Fe±Mg±Al temperatures compared with uncorrected Fe±Al temperatures, especially at higher absolute temperature (compare Fig. 7a and b, and 7d and e, respectively). Figure 8b shows a negative correlation between the above temperature Grt differences and XCa , probably a secondary effect opx Grt owing to the fact that XAl decreases as XCa increases (Fig. 6a). Figure 8c shows that there is no significant dependence of temperature difference on Mg/(Mg ‡ Fe)Grt . The mean difference between the corrected Fe±Mg± Al and uncorrected Fe±Al temperatures (points B and C in Fig. 3) is 45 Æ 9 and 34 Æ 4 C for aluminous and intermediate granulites, respectively (see Table 1). It could be argued that this difference is so small as to be immaterial, supporting Aranovich & Berman's (1997) cautionary view of applying recorrection schemes. Our view is that if demonstrable and systematic errors can be corrected for it is worth while to do so, especially when in a number of cases the difference becomes significant (e.g. Enderby Land examples given above). About 10% of the aluminous and intermediate samples have uncorrected Fe±Mg temperatures that are higher than the uncorrected Fe±Al temperatures. As discussed above, we consider the uncorrected Fe±Al P±T estimates to be more reliable for these samples. Mafic granulites In mafic granulites, the mean corrected Fe±Mg±Al estimate (816 Æ 12 C) is lower than in the intermediate and aluminous granulites but is still in agreement within error with the phase equilibrium constraints (Fig. 9a). In contrast, the mean uncorrected Fe±Mg exchange estimate (793 Æ 13 C) is considerably higher than in the intermediate and aluminous granulites. The higher mean pressure for the mafic granulites ($10 kbar) compared with the intermediate and aluminous granulites (6±8 kbar) is a result of the fact that garnet is a stable phase in mafic granulites only at relatively high pressure (e.g. Pattison, 2003). The temperature difference between the corrected Fe±Mg±Al estimate and the uncorrected Fe±Mg exchange estimate (23 Æ 16 C) is so small as to be insignificant (Table 7). Just under half of the mafic 885 JOURNAL OF PETROLOGY VOLUME 44 samples show Fe±Mg temperatures that are higher than Al-solubility-based temperatures (Fig. 7g±i). Figures 6 and 8a and b show that these results corresopx pond to the generally low XAl in the more Ca-rich mafic samples. The reasons for these patterns are unclear. Assuming that Fe±Mg always closes at lower temperature than Al, the most likely explanations are: (1) Fe±Mg diffusion is slower in Ca-rich garnets than in Ca-poor garnets, resulting in a smaller temperature gap between closure of Fe±Mg and Al; (2) the rocks experienced retrograde net-transfer reactions (Spear & Florence, 1992), leading to spuriously high Fe±Mg temperatures; (3) the parts of the Grt and Opx analyzed were not in equilibrium before late Fe±Mg exchange; (4) some of the analytical data for the generally low Al concentrations in Opx in these rocks are in error (too low); (5) the thermodynamic model for Al solubility in Opx loses opx accuracy at low XAl . Although slower Fe±Mg diffusion may account for the higher mean Fe±Mg exchange temperatures, it does not account for the many samples showing Fe±Mg temperatures that are higher than Al-solubility temperatures. Analytical inaccuracy seems unlikely as a general explanation because in studies in which several samples were analyzed using the same procedure (e.g. Adirondack Highlands, Furua Complex: Appendix and Electronic Appendix B), some samples indicate Al-solubility temperatures higher than Fe±Mg temperatures whereas others show the opposite. We see no reason why retrograde net-transfer reactions should be more prevalent in mafic granulites than in aluminous and intermediate granulites. We therefore think that thermodynamic inaccuracy is the most likely single explanation, perhaps augmented in some cases by selection of analysis points on minerals that were not in equilibrium and/or analytical errors. Additional experimental data and attendant thermodynamic modelling bearing on this question are needed. Comparison with other refractory methods of thermobarometry The results obtained with our Grt±Opx Al-solubility method are comparable with those obtained with other thermobarometry methods based on refractory cation systems, such as reintegrated feldspar thermometry [Kroll et al. (1993) and references therein] and reintegrated Fe±Ti-oxide±olivine±pyroxene thermometry (Frost & Lindsley, 1992; Lindsley & Frost, 1992). For example, reintegrated compositions of mesoperthitic alkali feldspar grains in sample 45-84 from the Kerala Khondalite Belt of south India (Chacko et al., 1987) indicate a temperature of 975 C [feldspar model of Fuhrman & Lindsley (1988)] compared NUMBER 5 MAY 2003 with uncorrected Fe±Mg and corrected Fe±Mg±Al temperatures of 821 and 926 C, respectively, for the same sample. Mean corrected Fe±Mg±Al temperatures from the Enderby Land granulites ($950 C; see above) are similar to temperatures calculated from reintegrated mesoperthitic feldspars [950±1050 C, using the analyses of Ellis et al. (1980) and Sandiford (1985) and the feldspar model of Fuhrman & Lindsley (1988)], both of which show good agreement with the ultra-high temperatures (4950 C) indicated by metamorphic pigeonite (Sandiford & Powell, 1986) and the Spr ‡ Qtz mineral assemblages. Unfortunately, exsolution features in feldspars and pyroxenes are easily destroyed during later deformation (e.g. Frost & Chacko, 1989), and are therefore considerably less widespread in granulite terrains than Grt±Opx assemblages. Our results are also comparable with those calculated with the oxygen isotope thermometry method of Farquhar et al. (1993), which corrects for the effects of retrograde isotope exchange. In applying this scheme to the high-temperature Enderby Land granulites and Spl ‡ Qtz-bearing Taltson granulites (Chacko et al., 1994; Berman & Bostock, 1997; Grover et al., 1997), Farquhar et al. (1996) retrieved oxygen isotope temperatures 4900 C that are consistent with the mineral assemblage stabilities. Strategies for effective use of the method The scatter of P±T results using our method calls into question the reliability of any individual P±T estimate, and suggests that to be confident of a P±T estimate for a given area, many samples need to be analyzed. We also advocate element (X-ray) mapping of minerals before analysis points for thermobarometric calculations are selected so that zoning patterns can be interpreted and compositions that are obviously out of equilibrium can be avoided (see Pattison & Bgin, e 1994a; Kohn & Spear, 2000). Even with X-ray maps, non-central sectioning of Opx grains is an important factor to consider, given the generally small core±rim variation in Al content of individual Opx grains (typically $1±2 wt %) and the strong temperature dependence on these small changes. An unanswered question of fundamental importance to the use of Al solubilitybased thermobarometry is the nature and controls of zoning of Al in Opx (e.g. McFarlane et al., 2003). IMPLICATIONS FOR P±T CONDITIONS OF GRANULITEFACIES METAMORPHISM The central conclusion of our study is that a significant number of thermobarometry-based temperature estimates for granulites over the past 30 years are too low 886 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Table 8: P±T results for different terrains No. of Type Author(s) estimates Uncorr. Fe±Mg Uncorr. Fe±Al Corr. Fe±Mg±Al T ( C) Terrain and abbreviation in Fig. 9 T P T T 10.2 5.2 840 samples P (kbar) Adirondack Highlands AD 11 M 750±800 7.5±8.0 900 Ashuanipi AS 10 I 700±835 3.5±6.5 7.0±10.0 690 Enderby Land EL 25 A 900±950 English River ER 7 I 700±750 Furua Complex FC 16 M 800 840 760 7.5 880 620 5.0 10.0 4.8 11.8 750 850 890 Karnataka KA 7 I/M 700±800 5.0±7.0 720 7.7 800 Kasai KS 6 I/A 720 680 KE 34 I/A 700±800 6.0 6.5 770 Kerala Khondalite Belt 6.7 5.0±7.0 780 880 Ketilidian Belt KB 9 A/I 650±800 2.0±4.0 720 3.7 890 Madras MD 8 I/M 770±830 750 MA 11 A 700±900 7.4 5.0 830 Makhavinekh 6.0±8.0 5.0 600 760 Minto MI 7 I/A 750±950 6.0±9.0 700 5.3 800 Nain NA 16 A 645±915 760 NHs 11 I/M 730±750 5.0 7.8 850 Nilgiri Hills (Srik.)* 3.7±6.6 7.0±10.0 670 730 Nilgiri Hills (Raith)y NHr 51 M/I 730±750 7.0±10.0 700 8.7 760 Prydz Bay PB 16 I 860 780 QN 6 I 700 6.2 5.1 860 Quetico NE 6.0 4.0±6.0 630 810 Scottish aureolesz SC 7 A 700±850 3.0±5.0 740 4.1 830 Sri Lanka Highlands SL 10 M/I 820 740 TA 15 I/A 920±1045 7.1 5.9 780 Taltson 8.0 6.0±7.8 840 910 P 9.2 7.3 8.9 6.5 12.5 9.0 7.2 7.8 5.8 8.4 5.0 6.7 5.0 8.7 9.7 7.3 7.6 4.2 7.7 6.7 820 880 950 810 900 840 810 920 940 860 830 840 870 740 800 890 880 840 800 930 P 9.2 7.8 8.9 6.5 12.6 9.3 7.2 7.9 5.8 8.5 5.0 6.7 5.0 8.9 9.9 7.2 8.0 4.2 7.7 6.7 Uncorrected Fe±Mg: point A in Fig. 3a. Uncorrected Fe±Al: point B in Fig. 3a. Corrected Fe±Mg±Al: point C in Fig. 3b. M, mafic; I, intermediate; A, aluminous. à Srikantappa et al. (1992) dataset (see Appendix). yRaith et al. (1990) dataset (see Appendix). zCombination of Ballachulish and NE Scotland aureoles from Appendix and Electronic Appendix B. References to localities are given in the Appendix. and are therefore misleading. Many of these estimates are inconsistent with the stability of the mineral assemblages in the rock. A higher temperature for the amphibolite±granulite transition compared with traditional estimates (Fig. 2a) spreads out the P±T range of the upper amphibolite facies. Concomitantly, it reduces the P±T interval between `ordinary' granulite-facies metamorphism and ultra-high-temperature metamorphism (Harley, 1998a; Fig. 2b), making the latter the high-temperature end of a continuum rather than a thermally distinct anomaly. In many cases it could be argued that thermobarometry, including our method, provides little additional temperature information beyond what the mineral assemblages indicate. Where our method of thermobarometry may be most useful is for bulk compositions that maintain the same mineral assemblage over large ranges of elevated P and T, such as in the 800±1000 C range for intermediate and mafic bulk compositions (e.g. Fig. 1). P±T ESTIMATES OF GRANULITE TERRAINS Table 8 provides a summary of mean P±T results for 24 terrains with six or more samples. The uncorrected Fe±Mg P±T estimates and corrected Fe±Mg±Al P±T estimates are plotted in Fig. 9b with respect to the granulite-facies-limiting reactions from Figs 1 and 2. In terrains in which the corrected Fe±Mg±Al P±T estimates are lower than the uncorrected Fe±Al P±T estimates (e.g. Adirondack Highlands), we have plotted the corrected Fe±Mg±Al P±T estimates to maintain consistency, even though we favour the higher estimates. In some of the terrains, such as the Kerala Khondalite Belt and the Nilgiri Hills, the mean P±T estimates are to some degree meaningless because of significant P±T variations across the region from which the samples were collected. The main purpose of Fig. 9b is to show that the corrected Fe±Mg±Al P±T estimates 887 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Fig. 10. (a) Map of isograds and metamorphic isotherms in the Adirondacks, from Bohlen et al. (1985). The various Opx-in isograds have been discussed by Bohlen et al. (1985) and Valley et al. (1990). The isograd labelled 1 and 5 represents the close coincidence of the incoming of Opx in metabasites and Crd ‡ Grt in metapelites according to De Waard (1969). (b) Revised isotherms in NW Adirondacks according to Kitchen & Valley (1995). (c) Comparison of P±T conditions along transect A±B in (a) with the granulite facies-limiting reactions and the Kfs ‡ Sil-in reaction from Fig. 1. Reactions are numbered as in the text and Fig. 1. Ovals show the inferred P±T conditions in the vicinity of the Opx-in isograd according to Bohlen et al. (1985) and Kitchen & Valley (1995). largely fall in or close to the granulite-facies stability field, in contrast to the uncorrected Fe±Mg P±T estimates which typically fall well below the granulitefacies stability field. Terrains in which a significant proportion of the sample suite consists of mafic granulites tend to show the lowest P±T estimates [e.g. the Nilgiri Hills datasets of Raith et al. (1990) and Srikantappa et al. (1992)], and may be unreliable for the reasons discussed above. Some well-known granulite terrains are discussed separately below. Adirondacks The Precambrian rocks of the Adirondack region of upper New York State (Fig. 10a) comprise one of the best-known granulite-facies terrains in the world. In their summary papers on the granulite-facies metamorphism of the Adirondacks, Bohlen et al. (1985) and Valley et al. (1990) presented a pattern of isotherms based on a variety of geothermobarometers that they considered to represent peak or near-peak P±T conditions. Kitchen & Valley (1995) modified the distribution of isotherms in the NW part of the Adirondacks (Fig. 10b). Valley et al. (1990) concluded that the granulite-facies metamorphism was largely driven by the magmatic processes of intrusion and partial melting, based on low calculated values of aH2O, abun- dance of migmatitic features associated with the granulite mineral assemblages, and absence of evidence for large-scale fluid infiltration. The positions of the Kfs ‡ Sil-in, Grt ‡ Crd-in and Opx-in isograds, discussed by Bohlen et al. (1985), are shown on the map in Fig. 10a. These isograds are represented approximately by reactions (4), (5) and (1), respectively. Figure 10c shows the range of P±T conditions along transect A±B in Fig. 10a compared with the limiting P±T stability fields for Kfs ‡ Sil, Grt ‡ Crd and Opx. The Bohlen et al. (1985) P±T conditions are lower by up to 200 C than the minimum stability ranges of Grt ‡ Crd and Opx, and, in the vicinity of the Grt ‡ Crd-in and Opx-in isograds, are below the minimum stability of Kfs ‡ Sil. The somewhat higher temperatures deduced by Kitchen & Valley (1995) are consistent within error of Kfs ‡ Sil stability, but are still considerably lower than those necessary to produce Grt ‡ Crd- or Opx-bearing assemblages. Three possible explanations are: (1) the peak P±T conditions have been significantly underestimated, especially in the vicinity of the isograds, as a result of retrograde cation exchange from peak conditions; (2) the isotherms record a later cryptic, broadly amphibolite-grade, metamorphic event that reset some element systematics but did not modify the peak mineral 888 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM assemblages; (3) low-aH2O fluid infiltration has been widespread. Locally variable aH2O in the absence of fluid infiltration is not a tenable explanation if the peak mineral assemblages were developed by partial melting because aH2O is internally buffered by the mineral ‡ melt assemblage. We accept the evidence of Valley et al. (1990) against widespread fluid infiltration and therefore favour an explanation involving P±T underestimation as a result of either retrograde exchange from peak conditions or the effects of a cryptic, lower-grade overprint. An independent indication of higher peak temperatures in the Adirondack Lowlands in the vicinity of the isograds comes from a recalculated Fe±Mg±Al solubility temperature of 820 C for the one Grt ‡ Opxbearing sample (RS-34) reported by Edwards & Essene (1988). In the Adirondack Highlands, several recent studies have suggested peak temperatures higher than $850 C, including those by Spear & Markussen (1997) and Alcock & Muller (1999). A puzzling aspect of the Adirondacks results in Table 8 is the higher mean temperature from Fe±Mg exchange than from the Al-solubility-based methods, which may be due to some or all of the factors discussed above for mafic granulites in general. Acadian metamorphic high One of the classic prograde amphibolite±granulite transitions is represented by the Acadian metamorphic high in central Massachusetts. The higher-grade parts consist of the following zones, defined by mineral assemblages in pelitic compositions: Zone IIIÐMs ‡ Sil zone; Zone IVÐMs ‡ Sil ‡ Kfs zone; Zone VÐ Kfs ‡ Sil zone; Zone VIÐGrt ‡ Crd ‡ Sil ‡ Kfs zone (Schumacher et al., 1990a). Within Zone VI, Opx ‡ Cpx ‡ Pl assemblages occur in metabasites. The Zone III±IV transition reaction corresponds approximately to reaction (4), the Zone V±VI transition reaction corresponds approximately to reaction (5), and the development of metabasic Opx ‡ Cpx ‡ Pl assemblages in Zone VI implies P±T conditions above reaction (1). Schumacher et al. (1990a, fig. 9.9) provided the following approximate temperature boundaries between the zones, based on Grt±Bt Fe±Mg exchange thermometry assuming a pressure of 6 kbar: III±IV, 640 C; IV±V, 670 C; V±VI, 690 C. Thomson (2001) estimated peak temperatures of 700±750 C for Grt ‡ Crd-bearing metapelitic granulites in Zone VI. Whereas Schumacher et al.'s Zone III±IV estimate is in reasonable agreement with the minimum stability of Kfs ‡ Sil at 6 kbar ($680 C), the Zone V±VI estimate is 450 C below the minimum stability of Grt ‡ Crd in metapelites [reaction (5)] and 4100 C below the minimum stability of Opx ‡ Cpx ‡ Pl in metabasites [reaction (1)] (Fig. 9). Lack of reported Grt ‡ Opxbearing assemblages in Zone VI does not permit estimation of peak temperatures by our recorrection method. Incipient charnockites of southern India and Sri Lanka The incipient charnockite localities of southern India and Sri Lanka are characterized by the development of green-weathering, Opx-bearing assemblages in discrete planar and linear networks within white, grey and pink Opx-free gneisses (e.g. Janardhan et al., 1982). These localities were the focus of intense interest in the 1980s and early 1990s because their formation was ascribed to channelled infiltration of low-aH2O (carbonic) fluids (e.g. Newton et al., 1980; Janardhan et al., 1982; Hansen et al., 1987; Perchuk et al., 2000), leading to a debate over the relative importance of infiltration-driven carbonic metamorphism vs thermally driven partial melting in the generation of granulites. We have applied RCLC to 18 incipient charnockite localities in the literature (six of the seven Karnataka± Tamil Nadu samples in the Appendix, eliminating one outlier, and samples 83-123, 4-10a, 121-166, 141-201, M-4, 23, 25, K18-6a, K18-17, 147-214, TN3-1 and TN21-4 from the Kerala Khondalite Belt). Two of the samples are mafic granulites with the rest being intermediate granulites. The mean and 95% confidence limit on the mean of the temperature estimates is 827 Æ 18 C for a pressure range of 6±8 kbar, not significantly different from the equivalent values for all intermediate granulites (841 Æ 11 C; Table 7). Even though low-aH2O fluid infiltration appears to have triggered the production of Opx in these localities, the amount by which aH2O in the fluid was lower than ambient values in the host gneisses might have been rather modest if the host gneisses were close to a temperature where they would produce Opx by closedsystem dehydration melting. This observation supports the suggestion of Frost & Frost (1987) and Clemens (1992) that intrusion and degassing of mafic±charnockitic magmas, partial melting and local low-aH2O fluid infiltration are expected to be intimately related processes in granulite formation, and may occur in close proximity to one another in individual terrains rather than occurring as terrainscale end-members. The southern Indian incipient charnockites may therefore represent sporadically developed, slightly lower-temperature, fluid-triggered granulite `fronts' that develop locally a little downgrade of the main expanse of granulite, the latter controlled largely by magmatic and partial melting processes. A counter-argument to the generality of 889 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Fig. 11. Generalized geological map of southern India (Chacko et al., 1996) showing the dominantly metasedimentary Kerala Khondalite Belt and adjacent Kodaikanal±Cardamon Hills and Nagercoil charnockite massifs. Samples used for thermobarometry have been described by Chacko et al. (1996) and are listed in the Appendix and Electronic Appendix B. (a) Uncorrected Grt±Opx Fe±Mg exchange temperatures. Note the absence of any spatial pattern in the temperatures. (b) Corrected Grt±Opx Fe±Mg±Al temperatures. Note the higher temperatures on the margins of the belt adjacent to the charnockite massifs. (c) Corrected Grt±Opx Fe±Mg±Al temperatures taking account of stoichiometrically calculated Fe3 ‡ in Opx. Note the somewhat lower temperatures compared with (b) but the same overall regional pattern. this inference comes from the experimental study by Nair & Chacko (2002) on dehydration melting of the host gneisses to some of the southern India incipient charnockite localities. As is typical of many of the south Indian localities, the gneisses contain biotite with high Ti and F contents. The gneisses did not undergo fluidabsent melting until temperatures in excess of 900 C (at P ! 6 kbar) were reached. For these localities, mean temperatures of $830 C from RCLC are well below the temperatures necessary for fluid-absent melting, suggesting that aH2O in the infiltrating fluid was significantly lower than in the host gneisses. Kerala Khondalite Belt The Kerala Khondalite Belt (KKB) of southern India (Fig. 11) is an example of a terrain in which the results of Grt±Opx Al-solubility-based P±T estimation resulted in a complete reinterpretation of the P±T 890 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM regime and associated thermotectonic evolution (Chacko et al., 1996). Earlier studies based on Fe±Mg exchange methods (e.g. Chacko et al., 1987) suggested a rather uniform P±T regime across the belt of 5±6 kbar and 700±800 C (Fig. 11a). Recalculation of the same samples using RCLC reveals a sharp contrast between a lower P±T ($800±850 C, 6 kbar) central zone with numerous incipient charnockite localities, a northern marginal zone where extreme P±T conditions (4950 C, 9±10 kbar) are found, and a southern marginal zone where less extreme but still elevated temperatures of 850±950 C are found (Fig. 11b). Calculated temperatures taking account of stoichiometrically determined Fe3 ‡ in Opx are on average 44 C lower for the whole sample suite but reveal the same regional pattern (Fig. 11c). The elevated temperatures in the marginal zones were attributed by Chacko et al. (1996) to the intrusion of igneous charnockite in the massifs to the north and south of the KKB. Nandakumar & Harley (2000) came to similar conclusions based on an independent set of samples. The 4950 C temperatures of the northern zone can be confirmed in a limited number of samples with exsolved feldspars. In the southern zone, Braun et al. (1996) found mesoperthites indicating temperatures of 900±1000 C and reported a few occurrences of Spl ‡ Qtz. Interestingly, despite its high temperatures, the northern zone is not characterized by the widespread development of mineral assemblages indicative of ultra-high-temperature conditions (e.g. Spr, Opx ‡ Sil, Spl ‡ Qtz). The paucity of these assemblages may be due to a combination of bulk composition and P±T conditions. The moderate Mg/(Mg ‡ Fe) of most rocks in this region does not favour the formation of the Mg-rich minerals sapphirine or osumulite, and at these relatively high pressures the stability of Spl ‡ Qtz assemblages is restricted to T 4 1000 C except in oxidized or high-Zn rocks (Fig. 1a). On the other hand, in and adjacent to the Kodaikanal±Cardamon Hills massif, the occurrence of Spr- and Opx ‡ Silbearing assemblages (Raith et al., 1997) indicates very high temperatures in the massif as a whole. Thus, in terrains such as the Kerala Khondalite Belt, Grt±Opx thermobarometry may provide the most widely applicable and effective means available for retrieving hightemperature data from granulites that show little mineralogical evidence for such high temperatures. THERMOTECTONIC MODELLING OF GRANULITES The indication that most `ordinary' granulites form at considerably higher temperature than previously assumed carries significant implications for thermotectonic models of granulite formation. In granulites that show isobaric cooling paths and that may have formed along anti-clockwise P±T paths, the heat source for the metamorphism is usually ascribed to mafic magmatic underplating (e.g. Bohlen, 1987; Harley, 1989). In granulites preserving evidence for isothermal decompression and that may have formed along clockwise P±T paths, the metamorphism is commonly ascribed to crustal thickening and associated internal heating in collisional orogens (Bohlen, 1987; Harley, 1989). Using standard mantle heat flow and radioactive heat generation parameters, temperatures in the range 650±800 C can be attained in the middle crust ($20± 30 km depth) by this means (England & Thompson, 1984; Pati~o-Douce et al., 1990; Ashwal et al., 1992; n Jamieson et al., 2002). If, however, mid-crustal granulites typically form at temperatures of $850 C and above (Fig. 1 and Table 8), much higher heat flow and heat generation parameters, or preferential incorporation of high heat-producing material at midto lower-crustal levels (Pati~o-Douce et al., 1990; n Jamieson et al., 2000) is required. Alternatively, it may be that even in collisional settings, advection of heat into the middle crust by mafic or charnockitic magmas (see Bohlen, 1987; Frost & Frost, 1987; Chacko et al., 1996) may be needed for granulite-facies metamorphism. ACKNOWLEDGEMENTS This research was supported by NSERC Discovery Grants 0037233 to D.R.M.P. and 0046751 to T.C. 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United Kingdom Water and Environmental Resources Council ± Progress in Experimental Petrology 6, 17±19. Xishan, L., Wei, J., Shuxun, L. & Xuechun, X. (1993). Two types of Precambrian high-grade metamorphism, Inner Mongolia, China. Journal of Metamorphic Geology 11, 499±510. Yardley, B. W. D. (1989). An Introduction to Metamorphic Petrology. Harlow: Longman, 248 pp. Yund, R. A. (1986). Interdiffusion of NaSi±CaAl in peristerite. Physics and Chemistry of Minerals 13, 11±16. SUPPLEMENTARY DATA Supplementary data for this paper are available on Journal of Petrology online. APPENDIX: LOCALITIES, SAMPLES AND REFERENCES Locality and samples Reference(s) Regional/contact Type* Adirondack Highlands, NY Bohlen & Essene (1979; 1980) R M Edwards & Essene (1988) R I Glassley & Sorensen (1980) R M Griffin et al. (1980) R M Riciputi et al. (1990) R I Lamb et al. (1986) R M Percival (1991) R I Ackermand et al. (1987) R A Pattison (1989 and unpublished) C A Bhattacharya & Mukherjee (1987) R M Hansen & Stuk (1993) R M Bucher-Nurminen & Ohta (1993) R I Loock et al. (1990) R M Harley (1985); Harley et al. (1990) R I/A 11 AS3, BM13, BM2, ET1, ET15, ET24, IN11, N3, S15, SR29, SR31 Adirondack Lowlands, NY 1 RS34 Agto, W. Greenland 3 88589, 91154, 91156 Akia, W. Greenland 3 152735, 152733, 152765 Akia, W. Greenland 3 266935, 341925, 341965 Arendal, Norway 6 B750, C7652, C814, C815, D714, D7606 Ashuanipi, Quebec 10 27, 47, 96, 98, 142, 152, 216, 262, 266, 347 Bahai, Brazil 1 CAR-19 Ballachulish, Scotland 2 568-1, 568-2 Bengal, E. India 2 SM-4b, SM-44b Cone Peak, California 6 82G4t, 87G2, 87G3, 87G4, H2a, H3 Dronning Maud Land, E. Antarctica 1 KBN-702 Eifel, Germany 4 EIF-S1, EIF-S16, EIF-S35, EIF-S37 Enderby Land, Antarctica 25 49659, 3411, 3423, 3468, 3532, 3552, 3597, 3964, 3970, 3973, 4003, 4528, 4547, 4549, 4593, 4833, 49382, 49657, 49734, 49740, 49748, 49749, 49891, 4648, 4652 897 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Appendix: continued Locality and samples Reference(s) Regional/contact Type* English River, Ontario Perkins & Chipera (1985) R I Bradshaw (1989a) R M Coolen (1980) R M Mohan & Windley (1993) R I Anovitz & Essene (1990) R M Wodicka et al. (2000) R A/M Pattison (1991) R M Perkins et al. (1982) R M/I Iyer et al. (1996) R M Komatsu et al. (1994) R A Boullier & Barbey (1988) R M/I Sills et al. (1983); Sills (1984) R I/M Hand et al. (1994) R A Percival (1983) R I Percival & McGrath (1986) R I Janardhan et al. (1982) R I/M Hansen et al. (1987) R I Bingen et al. (1988) R A Chacko et al. (1987, R I/A Harris et al. (1982) R I/A Hansen et al. (1987) R I Nandakumar & Harley (2000) R I/A Kerala Khondalite Belt, S. India Santosh (1987); Santosh et al. R A/I 2 86-2, M-4 (1990) 7 ID1683, G1a, G63B, RW1283, BL1083, CL2283, HS1583 Fiordland, New Zealand 1 51112 Furua Complex, Tanzania 16 1, 2, 4, 5, 6, 11, 22, 23, 25, 26, 28, 33, 49, 55, 91, 94 Ganguvarpatti, S. India 3 97-127, 97125-40, 97125-86 Grenville Province, Ontario 8 BA-9A, 83C6, 83C5, 83C70, 83C9, 85A12, 83C15, 83C19 Grenville Prov. (Georgian Bay), Ont. 2 N131a, N153c Grenville Prov. (Huntsville), Ont. 3 Hu1, P2-6-2b, Hu15a Grenville Prov. (Otter Lake), Quebec 6 AL2, DD11, DD17, DD12, RK1, RK2  Guaxupe, Brazil 2 AC89, AC89s Hidaka, Japan 1 HS412p2 Iforas, Mali 3 IC-114a, IC115, IC115b Ivrea Zone, Italy 4 IVL, IV101, IV346, IV383 Jetty Peninsula, E. Antarctica 1 JT6665 Kapuskasing, Ontario 4 PG13, PG16, PG21, PG22 Kapuskasing (North), Ontario 6 Kap-5, Kap-6, Kap-7, Kap-8, Kap-9, Kap-10 Karnataka-Tamil Nadu, S. India 6 11-1b, 11-3e, 11-1, GN-4a, 4-1c, 10-1a Karnataka-Tamil Nadu, S. India 1 T13-83 Kasai, Zaire 6 12195, 139131, 139160, 73535, 139246, 139146 Kerala Khondalite Belt, S. India 17 76-112, 96-136, 165-241, 77-115, 95-135, 45-84, 55-89, 83-123, 170-246, 1996 and unpublished) 177-260, 4-10a, 121-166, 141-201, 153-225, 154-226, 155-229, 147-217 Kerala Khondalite Belt, S. India 4 732, 758, 766, 770 Kerala Khondalite Belt, S. India 5 K18-6a, K18-7, 147-214, TN3-1, TN21-4 Kerala Khondalite Belt, S. India 3 ED3, ED8, ED9 898 PATTISON et al. TEMPERATURES OF GRANULITE-FACIES METAMORPHISM Locality and samples Reference(s) Regional/contact Type* Kerala Khondalite Belt, S. India Srikantappa et al. (1985) R I Dempster et al. (1991) R/C A/I Speer (1982) C A/I Grant & Frost (1990) C A/I Owen & Erdmer (1989) R I/A Sen & Bhattacharya (1984) R I/M Munyanyiwa et al. (1993) R M McFarlane et al. (2003) C A Perkins & Chipera (1985) R I Minto, Quebec  Begin & Pattison (1994a); R I/A 6 B69E, C10B, P88, B58, B74b, P69  Pattison & Begin (1994) Molodezhnaya, Antarctica Grew (1981) R I Xishan et al. (1993) R M Berg (1977a, 1977b) C/R A Srikantappa et al. (1992) R I/M Raith et al. (1990) R M/I Lal et al. (1987) R A Ravindra Kumar & Chacko (1994) R M Raith et al. (1997) R A Arima & Barnett (1984) R A Fitzsimons & Harley (1994) R I Thost et al. (1991) R M Percival & McGrath (1986) R I 3 23, 25, 36 Ketilidian Belt, S. Greenland 9 A255, A289, A374, A375, A379, I65, Q62, Q64, Q67 Kiglapait, Labrador 2 NC-49, SN-211 Laramie, Wyoming 3 19-4a, 53-2b, 53-2c Long Range Inlier, Nfld 3 86-451-2, 86-111, 86-279 Madras Charnockites, S. India 8 379, 558, 702, 715, 729, 403b, 543a, 79112 Magondi, N. Zimbabwe 3 NY81-3, NY84-1, NY84-2 Makhavinekh, Labrador 11 M02-20, M15-250, M04-450, M05-800, M17-1015, M21-1500, M22-2025, M23-2400, M08-3125, M12-4025, T03-5750 Minnesota River, USA 5 GF3-8, GF3-6, GF5-4, GF5-3, GF5-2 3 322, 322a, 341b Inner Mongolia, China 2 MONGO010, MONGO691 Nain, Labrador 16 LRD72, KI3909, KI3557, 2-893, 2-625, 2-1833, 2-275, RAW437, NU69, NK420b, 74-98, 2-1726, 2-1637, KI3911, 74-16b, 74-18x Nilgiri Hills, S. India 11 78, 85, 88, 95, 110, 182, 304, 314, 337, KU4, KU7 Nilgiri Hills, S. India 51 2, 3, 4, 5, 6, 7, 8, 10, 11, 12, 13, 15, 18, 19, 20, 22, 31, 32, 33, 34, 35, 36, 37, 38, 39, 40, 41, 42, 43, 44, 45, 46, 47, 48, 49, 50, 51, 52, 53, 54, 55, 56, 57, 58, 61, 64, 65, 68, 72, 76, 78 Paderu, S. India 3 252, 259, 320 Palghat Gap, S. India 4 112, 112a, 122, 170 Palni Hills, S. India 1 PH-172b Pikwitonei, Manitoba 2 PIKPO4a1, PIKPO4b Prydz Bay, Antarctica 16 88/118, 89/71, 89/63, 89/59, 89/334, 89/2, 89/108, 88/88, 88/65, 88/50, 88/41, 88/38, 88/333, 88/168, 89/40, 89/106 Prydz Bay, E. Antarctica 1 8813001 North Quetico, Ontario 6 Q-1, Q-2, Q-4, Q-5, Q-6, Q-7 899 JOURNAL OF PETROLOGY VOLUME 44 NUMBER 5 MAY 2003 Appendix: continued Locality and samples Reference(s) Regional/contact Type* S.-Central Quetico, Ontario Pan et al. (1994) R I Harley & Fitzsimons (1991) R I Barth & May (1992) R M NE Scotland Ashworth & Chinner (1978); Droop C A 5 BELHM5, BELSP1d, HP105570, HP82891, HPCB1 & Charnley (1985) Scourie, Scotland Savage & Sills (1980); Rollinson R M/I 2 SCOUR84, 275 (1981) Highland Series, Sri Lanka Schumacher et al. (1990b) R M/I Hansen et al. (1987) R I Johansson et al. (1991) R M Berman & Bostock (1997) R I/A Riciputi et al. (1990) R M Mengel & Rivers (1991, 1997) R M/I St-Onge & Lucas (1995) R M Percival & McGrath (1986) R I 3 YP01B, G409, G2229 Rauer Group, E. Antarctica 3 88-335, 65768c, 88-218a San Gabriel, S. California 4 273b, CB86-1a, CB88-5, CC83-157 10 K130-1, K144-4, K1-5, K200, K242-2, K302, K355-1, K408, K460-1, K50-1 Highland Series, Sri Lanka 1 D4-N2 Sveconorwegian Belt, Sweden 2 HALLANDSAS, TYLOSAND Taltson, NWT 15 106b, 145b, 187c, 2063a, 223b, 258a, 310a, 317a, 353a, 40a, 505b, 59a, 59b, 718a, 96b Tasiusarsuaq, W. Greenland 2 6728-22, 6728-27 Torngat, Labrador 3 F84-110, MZ-194, 25b Ungava, Quebec 1 UNGD-192 Val Rita, Ontario 5 VR-1, VR-2, VR-3, VR-4, VR-5 Total: 414 samples. *M, mafic; I, intermediate; A, aluminous. 900 ...
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