Journal of Petrology Advance Access originally published online on December 10, 2004
Journal of Petrology 2005 46(2):355-375; doi:10.1093/petrology/egh079
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Journal of Petrology vol. 46 issue 2 © Oxford University Press 2004; all rights reserved
A Low-Variance Mineral Assemblage with Talc and Phengite in an Eclogite from the Saxonian Erzgebirge, Central Europe, and its P–T Evolution
1 INSTITUT FÜR MINERALOGIE UND KRISTALLCHEMIE, UNIVERSITÄT STUTTGART, AZENBERGSTRASSE 18, D-70174 STUTTGART, GERMANY
2 LANDESAMT FÜR GEOWISSENSCHAFTEN UND ROHSTOFFE BRANDENBURG, STAHNSDORFER DAMM 77, D-14532 KLEINMACHNOW, GERMANY
RECEIVED JUNE 26, 2003; ACCEPTED SEPTEMBER 13, 2004
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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A light-coloured, fine-grained eclogite sample from near the village of Hammerunterwiesenthal in the Erzgebirge (NW Bohemian Massif) preserves the low-variance mineral assemblage of garnet, omphacite, phengite, talc, amphibole, clinozoisite, quartz, rutile, and accessory phases. Porphyroblasts of amphibole, clinozoisite, and phengite formed during a late stage (III) of metamorphism. Paragonite joined the assemblage late in this stage (IIIb). The chemical zonation of the minerals was carefully studied. Various geothermobarometric methods were applied, especially involving phengite and talc. The constrained P–T path for the eclogite starts at about 480°C and 25 kbar (stage Ib), followed by a significant temperature rise (stage II) at slightly increasing pressure. At the peak P–T conditions of 720°C and 27 kbar, blastesis of amphibole, clinozoisite, and phengite was caused by infiltrating hydrous fluids. The resulting density reduction may have allowed buoyant uplift of the eclogite. Subsequently, significant cooling occurred at high pressures. Stage IIIb is characterized by P–T conditions around 520°C and 18 kbar at reduced water activities. This unusual late P–T evolution might explain the freshness of the eclogite, including the preservation of chemical zonation on the micrometre scale.
KEY WORDS: eclogite; Saxonian Erzgebirge; P–T evolution; talc; phengite
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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Eclogites showing textural equilibrium in high-pressure (HP) orogenic belts are commonly characterized by assemblages of minerals that rarely exceed three phases in addition to garnet, omphacite, and rutile. However, numerous other minerals can, in principle, coexist in metabasites with the above assemblage at pressure–temperature (P–T) conditions of the eclogite facies. For example, in low- to medium-T eclogites the following minerals can occur in more than accessory quantities: amphibole, carbonates (aragonite, calcite, and dolomite), chloritoid, clinozoisite (including epidote and zoisite), gedrite (including related orthoamphiboles), lawsonite, kyanite, paragonite, phengite, phlogopite, quartz (or coesite), staurolite, and talc. Some of these, such as amphibole, phengite and quartz, are frequent constituents of eclogites, whereas others appear only in exceptional cases. Examples for these cases are staurolite in eclogites of NW Spain (Gil Ibarguchi et al., 1991
), chloritoid in eclogites from various localities of the Alps (Miller, 1977; Kienast & Pognante, 1988; Messiga et al., 1999), and gedrite in eclogites from the Dabie Shan (Wang & Massonne, 2000). Another Fe–Mg-silicate, talc, may be more common but is, nevertheless, a rare constituent of eclogites. Because talc plays an important role in our study, we cite subsequently various occurrences of talc-bearing eclogites around the world: Austria (Miller, 1986; Miller & Thöni, 1995; Hoschek, 2001), China (Zhang & Liou, 1994; Zhang et al., 1995; Liu et al., 1998), Italy (Pognante, 1989), Switzerland (Zack et al., 2001), Togo (Menot & Seddoh, 1985), and Venezuela (Sisson et al., 1997), which we will also consider for comparison later. Here we report an eclogite from the Saxonian Erzgebirge, central Europe, with a rare low-variance assemblage involving amphibole, clinozoisite, paragonite, phengite, quartz, and talc. We evaluate whether this assemblage together with omphacite, garnet, and rutile represents a true equilibrium assemblage. In addition, we apply several geothermobarometers to this assemblage, and demonstrate its utility for P–T estimates. The significance of the resulting P–T path for understanding the evolution of extended areas with HP metamorphism in the Bohemian Massif, of which the Erzgebirge is a part, will be considered.
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GEOLOGICAL SETTING AND SAMPLE DESCRIPTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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The crystalline complex of the Erzgebirge is situated in Saxony, Germany, and the Northern Czech Republic (Fig. 1), forming an oval-shaped antiformal structure extending nearly 80 km in a NE–SW direction. This complex is located at the northwestern margin of the Variscan Bohemian Massif. The crystalline complex of the Erzgebirge can be subdivided into three major medium- to high-grade metamorphic units (Willner et al., 2000
) surrounded by a low-grade and low-pressure unit, the Phyllite Unit (Fig. 1). Two of these three units, the Mica-schist–Eclogite Unit (MEU) and Gneiss–Eclogite Unit (GEU), contain abundant eclogite lenses. The metabasic rocks of the GEU have experienced peak metamorphic temperatures (T) around 800°C (Massonne, 1994) or, according to Schmädicke et al. (1992), up to 900°C. In addition, ultrahigh-pressure (UHP) conditions are indicated by the occurrence of coesite (Massonne, 2001). In contrast, eclogites of the MEU were formed at lower P and T. Schmädicke et al. (1992) determined maximum metamorphic T for eclogites from the southwestern Erzgebirge of 630–700°C. Those workers estimated P as high as 24 kbar because of the presence of kyanite in some eclogites and ‘the preservation of paragonite and plagioclase as inclusions in garnet cores’. The latter ‘may indicate that pressure values were not much in excess of 22–24 kbar’ (Schmädicke et al., 1992). Massonne (1992) determined a maximum P of 23 kbar (before reaching the T climax somewhat above 600°C) on the basis of phengite barometry for an eclogite occurrence (sample E25d) situated 3·5 km NW of the town of Marienberg (close to the eclogites of the GEU; see Fig. 1). Klápová et al. (1998) gave a peak P of 25·6–26·1 kbar at 650–700°C for eclogites from the Czech part of the southwestern Erzgebirge. For the surrounding mica-schists, a peak P of 14 kbar at 600°C was estimated (Konopásek, 1998). Somewhat lower P–T conditions (12 kbar, 525°C and subsequently 8 kbar, 560°C) were determined for similar rocks occurring NW of the Czech occurrences on the German side of the Erzgebirge (Rötzler et al., 1998).
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The age of the HP metamorphism in the crystalline complex of the Erzgebirge is believed to be around 340 Ma (Kröner & Willner, 1998
). However, the age dating results of Schmädicke et al. (1995) also suggest some older HP events. Those workers reported an 40Ar–39Ar plateau age of 355 ± 2 Ma for phengite from an eclogite near Mildenau, which they assigned to the same zone as the eclogite described here. However, no 40Ar–39Ar plateau ages older than 341 Ma were reported by Werner & Lippolt (2000) for phengite and amphibole concentrates from eclogites of the MEU.
Our eclogite sample from the western Erzgebirge belongs to the MEU. It was collected close to a rock formation called the Stümpelfelsen (50°26·7'N, 12°59·4'E, Potsdam datum), which is located 1·5 km west of the village of Hammerunterwiesenthal (situated at the Czech border). According to the distribution of abundant loose eclogite blocks around this rock formation, an eclogite lens of some hundred square metres is inferred to be exposed at the surface here. A new quarry has confirmed the presence of a massive eclogite body embedded in garnet-bearing gneisses and mica-schists. This situation is repeated 5 km NW and 6·5 km south of the Stümpelfelsen at the landmark Siebensäure (S in Fig. 1) and at the rock formation Wirbelstein (W in Fig. 1; in Czech, Meluzína), respectively. The eclogite bodies are relatively uniform and similar to each other at all three localities. According to Massonne & Czambor (2003) these eclogites have chemical signatures typical for mid-ocean ridge basalts (MORB). At the Stümpelfelsen lighter (MgO contents around 7 wt %) and darker coloured (MgO contents around 6·5 wt %) eclogite varieties can be distinguished, which are not as foliated as the eclogites at the other sites (see Klápová et al., 1998). A common texture is produced by millimetre-long, non-oriented amphibole porphyroblasts set in a relatively fine-grained matrix of omphacite and garnet (Fig. 2). A few larger single white mica flakes are also discernible with the naked eye. In addition, several millimetre-sized amphiboles and oriented white mica flakes are concentrated on millimetre-thick planes often bordering quartz segregations cross-cutting the eclogite body. Cracks, occasionally running parallel with decimetre spacings, are filled with late retrograde minerals, such as chlorite and actinolite, causing a darker colouring over several millimetres around the cracks.
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The sample selected for detailed study (sample 18333) represents the light-coloured eclogite variety from the Stümpelfelsen. In this fresh eclogite, garnet (
25 vol. %) and omphacite (30 vol. %) show a narrow grain-size distribution with a mean close to 100 µm for garnet (omphacite is smaller). Amphibole, clinozoisite, quartz and the various phyllosilicates, phengite, talc and paragonite, are similar in grain size, forming the rock matrix. However, millimetre-sized porphyroblasts of amphibole and less abundant clinozoisite and phengite are also present, relatively homogeneously distributed in the sample. The ratio of porphyroblasts (+ inclusions of matrix minerals) to matrix is about 1:2. Amphibole and clinozoisite porphyroblasts are interpreted to have replaced omphacite because there is virtually no omphacite enclosed in these porphyroblasts, whereas the amount and size of garnet included in amphibole differ little from those of garnet in the matrix (Fig. 2). In addition, phengite and rarely talc can be found enclosed in amphibole and clinozoisite porphyroblasts. In contrast, phengite porphyroblasts show abundant inclusions of omphacite that are hardly smaller than the matrix omphacites whereas garnet in phengite is less abundant than in the matrix. The quantity of phyllosilicates in the matrix amounts to about 5 vol. % of the rock. Phengite is more frequent than talc in the matrix. Paragonite is only an accessory phase, forming well-shaped flakes (Fig. 3), whereas talc often appears marginally corroded. Further accessory phases are sulphides (pyrite and chalcopyrite) and apatite, which form randomly distributed larger grains. Rutile (1 vol. %) occurs as very fine grains enclosed in the various silicates and some small grains relatively homogeneously distributed in the matrix.
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MINERAL COMPOSITIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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Analytical techniques
The compositions of phases in a single thin section of sample 18333 were analysed by electron microprobe. We used a CAMEBAX (analysis numbers beginning with 1035 and 1036) with three wavelength-dispersive spectrometers (WDS) and a Cameca SX100 electron microprobe with five WDS to determine the concentrations of F, Na, Mg, Al, Si, K, Ca, Ti, Cr, Mn, Fe and Ba. Counting times were 20 s at the peak and on the background. For a few elements (Na, Ti, Cr and Mn) longer counting times up to 60 s and a large PET and LiF analysator crystal (for Ti, Cr, Mn and Fe) were used on the Cameca SX100 (analysis numbers beginning with 1702 and 1802). We used synthetic minerals (e.g. pyrope for Mg, Al, and Si for the Camebax) and natural minerals, glasses (e.g. Ba glass for the BaL
1 peak), and pure oxides as standards. The applied acceleration voltage and electric current were 15 kV and 15 nA, respectively. Beam diameter was about 5 µm but also 2 µm in the case of small minerals (analysed in manual mode). The PaP correction procedure provided by Cameca was applied. Concentration maps for major elements (number of simultaneously prepared maps according to the number of WDS) were prepared by stepwise movement of the thin section under the electron beam of the microprobe and subsequent computer-aided evaluation (e.g. Bernhardt et al., 1995). Counting times per step were 1 s on the Camebax and 100 or 200 ms on the Cameca SX100. The energy-dispersive systems of the microprobes served as means of identification of phases during the selection of spots for full analyses.
Garnet and omphacite
Garnet analyses are given in Table 1 and shown in Fig. 4. Compositions of garnet vary significantly between 13 and 28 mol % pyrope component. Variations in the grossular (24–32 mol %), spessartine (1·2–4·5 mol %), and almandine (45–53 mol %) components are significant but not as strong; andradite is less than 2 mol %. According to Fig. 4 and the element concentration maps shown in Figs 5–7, the highest pyrope and lowest grossular concentrations occur on garnet rims (GtE of Table 1), where spessartine is around 2 mol %. Rims of garnet outside amphibole and clinozoisite porphyroblasts can be somewhat richer in Mg compared with rims of garnets enclosed in these porphyroblasts. This could be interpreted as a proof for further matrix garnet growth after inclusion of garnets in these porphyroblasts. The cores or earliest garnet compositions are more difficult to define than the rims, as the lowest pyrope concentrations (GtB of Table 1) were not observed to occur in a regular concentric growth zone around the geometric centre of the garnets richer in Mg (GtA of Table 1) but in several small centres in garnet. For instance, it can be seen in the Mg image of Fig. 6 that in several garnets dark blue spots appear in light blue to green–blue areas. However, in the Ca images (see Figs 6 and 7) a fairly concentric zonation pattern of garnet is discernible, which would, for instance, exclude oscillatory zoning, often observed in Ca images of garnets from low- to medium-temperature metamorphic rocks (e.g. Schumacher et al., 1998; Willner et al., 2001), to explain the above Mg zonation. Either small garnet grains with different size (see dark blue ‘centres’ in garnet of the Mg image of Fig. 5) were amalgamated and overgrown by garnet of higher Mg content (GtC–GtD of Table 1), or garnet with an intermediate pyrope concentration (GtA) in the geometric centre of the grains formed first. An observation supporting the latter interpretation is that the Mn and Ti contents are highest in GtA and, thus, differ from those of GtC–GtD.
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Omphacite is variable in composition (see Table 2). In Fig. 8 a clear trend of decreasing Na and Al contents with rising concentrations of Ca and Mg is discernible. This is due to the substitution Na + Al(Fe3+) or jadeite (+ acmite) component = Ca + Mg(Fe2+) or diopside (+ hedenbergite) component. Na per formula unit (p.f.u.) in omphacite ranges between 0·39 and 0·46. However, the temporal succession of these compositions is not as clear as in garnet because in the small omphacite grains a concentric zonation pattern is not detectable. We occasionally observed an increase of diopside (+ decrease of Na p.f.u.) from inner to outer portions of the omphacites. Relatively high concentrations of Na p.f.u. are, however, common at the omphacite rims (see Fig. 8). We concluded from these observations that the earliest omphacite compositions, which are still preserved in the sample, are rare Na-rich compositions in omphacite cores (OmA in Table 2), which are additionally richest in acmite component (highest calculated Fe3+ content). These are chronologically followed by omphacite compositions with the lowest Fetot contents (OmB in Table 2). Further, towards the rim (OmC to OmE in Table 2), Mg and Fetot concentrations tend to decrease and increase, respectively.
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Phyllosilicates
Talc compositions vary slightly (see Fig. 9) containing around 4 and 0·4 wt % of FeO (no Fe3+ considered) and Al2O3, respectively. Low concentrations of Na2O (
0·1 wt %) could be proven. Cr, Mn and Ti (for the last see Fig. 9) contents are above the analytical detection limits for the entire population of talc analyses. The compositional change with time was as difficult to detect as for omphacite, for the same reasons as mentioned above. However, according to the observations made from the element concentration maps (e.g. Fig. 5) we consider that early talc was poorer in iron than late talc, leading to the temporal sequence of TcA to TcC in Table 3.
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Phengites are significantly zoned and occasionally show clear zonation patterns (see Fig. 5). The zonation refers to the K = Na exchange and the Tschermak's substitution. The latter, expressed by the Si content per double formula unit (p.d.f.u.), caused a variation between 6·66 and almost 7·00 Si p.d.f.u. (see Fig. 10 and Table 4). The MgO contents decrease with falling Si contents whereas FeO (we assume almost no Fe2O3) is nearly constant. Consequently, the Fe/(Fe + Mg) ratio increases with falling Si p.d.f.u. as does the Na/(Na + K) ratio (Fig. 10). The latter trend (i.e. increase of paragonite content with increasing muscovite component in phengite) is very strong. In contrast, Ti and Ba contents are more or less constant and low for the entire population of phengite analyses. From core to rim (i.e. early to late stage; PheA to PheE in Table 4) the Si content p.d.f.u. of phengite decreases. As the Si contents of phengite porphyroblasts are generally below 6·80 p.d.f.u. (see Fig. 10) these grains must have formed late in the textural sequence. In addition, the compositional spread is relatively low in the phengite porphyroblasts.
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Several paragonite analyses gave constant results pointing to about 10 mol % of muscovite component in paragonite (see Table 4) and minor Tschermak's component (Mg
0·05 p.d.f.u.). It is possible that Fe is trivalent in these paragonites.
Amphibole and clinozoisite
The diagrammatic presentation of amphibole analyses in Fig. 11 shows clear compositional variations. For instance, the Al2O3 and Na2O contents range between 2 and 9 wt % and 1 and 4 wt %, respectively. After calculating the structural formula of amphibole (for representative analyses see Table 5), it is obvious that the chemical variations are mainly related to the contents of Si, Mg and eight-fold coordinated Na (= Na[8]). This compositional spread is mostly seen in small amphibole grains in the matrix whereas the variation in composition is moderate for amphibole porphyroblasts. The clear concentric amphibole zonation described by Massonne (1992) for a similar, although somewhat coarser-grained, eclogite from the Erzgebirge, 3·5 km NW of Marienberg, could not be observed here (see Fig. 6). Nevertheless, the rims of the porphyroblasts are richer in Ca and Mg than the wide cores of these amphiboles. Si contents are around 7·5 p.f.u. for the entire porphyroblast whereas compositions close to tremolite with up to 0·4 Na[8] p.f.u. (AmphA and AmphB in Table 5) were detected for cores of small amphibole grains. The highest Na[8] contents (>0·8 p.f.u.) were, however, observed for rims of matrix grains or for wide inner areas of the amphibole porphyroblasts (AmphD in Table 5).
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Representative analyses of clinozoisite are given in Table 6. These show a spread in epidote component [Ca2Al2Fe3+Si3O12(OH)] between 11 and 43 mol % (all Fe was assumed to be trivalent). However, the matrix grains (CzsA to CzsC in Table 6) are significantly richer in epidote component than the porphyroblasts (Fig. 12), with no compositional overlap as observed for phengite and amphibole porphyroblasts and corresponding matrix grains. The clinozoisite porphyroblasts are only weakly zoned (see Fig. 7) with higher Fe3+ contents in core regions (CzsD in Table 6) than in rims (CzsE in Table 6). On the outermost margins of porphyroblasts (CzsF in Table 6), the Fe3+ contents are similar to those in the cores. Small amounts of Mg and Ti could be systematically detected in the clinozoisites with Mg and Ti decreasing as Al increases (see Fig. 12).
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GEOTHERMOBAROMETRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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General
With the exception of calibrations for the Fe2+–Mg partitioning between garnet and omphacite (see below), we have considered thermodynamic relations to calculate P–T conditions on the basis of various multivariant mineral equilibria. For that purpose, we have applied the internally consistent dataset of Berman (1988)
augmented by compatible data (some are slightly modified in regard to those in Berman's dataset) for almandine (Berman, 1990), pyrope (Massonne, 1992), Ti-pyrope (Mg3Al2Ti3O12; Brandelik & Massonne, 2001), Mg–Al-celadonite (Massonne, 1992), Fe2+–Al-celadonite (Massonne, 1997), Ti-muscovite [KAl3SiTi2O10(OH)2; Massonne et al., 1993], talc (Massonne, 1995a), Fe-talc (Stöckhert et al., 1997), and glaucophane (Evans, 1990). Complex activity formulations were used for components in garnet (Massonne, 1992; for Ti-in-garnet: Brandelik & Massonne, 2001), omphacite (Massonne, 1992) and phengite (Massonne, 1997; for Ti-in-phengite: Massonne et al., 1993). For other mineral components mostly simple activity relations (see below) were applied. The element symbols in the corresponding equations refer to the content of this element per (double) formula unit in the specific mineral phase. We used Ge0-Calc (Brown et al., 1988) to determine stoichiometric coefficients for complex mineral equilibria but most of the geothermobarometric computation was undertaken with PTGIBBS (Brandelik & Massonne, 2004). The onset of blastesis (stage III) and, thus, the inclusion of the previously formed minerals in amphibole, clinozoisite and phengite porphyroblasts was the only time frame that we could apply for relating mineral compositions to equilibrium assemblages. Because of the simple inclusion assemblage in garnet (rutile and a SiO2 phase, probably only quartz) we could only use the zonation of the minerals in the matrix and enclosed in the porphyroblasts for a correlation of coexisting mineral compositions (stages I and II). The early compositions of the zoned minerals, as explained in the previous section, were related to the earliest metamorphic stage I. The later formed compositions, before the onset of blastesis, were assigned to stage II. Stages I and III were further subdivided into an early and a late stage (notation a and b, respectively). Thus, stage IIIb refers to the rim composition of the porphyroblasts. Stage IV is related to fine-grained decomposition products (found exceptionally at the outermost rims of minerals), mainly of omphacite. Because of the uncertain compositional correlation within metamorphic stages I and II, we have combined various compositions of coexisting minerals in the subsequent geothermobarometric evaluation.
Mineral equilibria and calibrations applied
The metamorphic T of the various metamorphic stages was determined on the basis of the Fe2+–Mg distribution between omphacite (Om) and garnet (Gt) as a result of the equilibrium
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and Krogh Ravna (2000) considering the calculated Fe2+ contents for both minerals (Tables 1 and 2). Equilibrium (1) was also considered by computing various independent equilibria for the components almandine (Alm), grossular (Gro) and pyrope (Pyr) in garnet, diopside (Diop) and hedenbergite (Hed) in omphacite, and muscovite (Ms), Mg–Al-celadonite (MgAlCel) and Fe2+–Al-celadonite (FeAlCel) in phengite (Phe). Among these are
 | (2a) |
 | (2b) |
 | (2c) |
 | (3) |
) was used and the activities of quartz (Qz) and rutile were set to unity.
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Because of the coexistence of talc (Tc) and amphibole (Amph) in the eclogitic matrix we could apply two further equilibria for geobarometry that are independent of water activity (aV):
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 | (4b,c) |
. We used PTGIBBS, the previously applied activity formulation for omphacite, and the activity formulation aTc = [Mg/(Mg + Fe + Mn + Na)]3·{[4 – (Al + Cr)/2]/4}4 for the activity of the talc component in talc. For the tremolite (Trem) and glaucophane (Glau) component in amphibole, the activity formulations by Massonne (1992), aTrem = XCa2 · 
1 · 
· 2 · Mg/(5 – Na[8]) and aGlau = (XAl · Na[8]/2)2 · 1 · · 2 · XMg (Massonne, 1995b) with XCa = Ca/(Ca + Na[8]), XAl = Al/(Al + Fe3+), indicating vacancy in regard of the A-site of the amphibole structure, and XMg = Mg/(Mg + Fe2+ + Mn), were applied. Because these formulations, especially the one for aGlau, were mainly designed for low-T amphiboles close to the tremolite–riebeckite–glaucophane series, we have also tested the activity model aGlau = (XAl · Na[8]/2)2 · 1 · · 2 · XMg3 · XSi8, which accounts for Si contents p.f.u. deviating from the ideal value of eight (XSi = Si/8). This model leads to activities closer to those obtained by applying the formulation aGlau = (XAl · Na[8]/2)2 · XMg3 as proposed by Evans (1990). For instance, for amphibole composition AmphE in Table 5, Evans's model yields 0·037 for aGlau, whereas a value of 0·1139 [Massonne's (1992) activity model: 0·2487] is obtained with our new model for 700°C.
The additional presence of clinozoisite (Czs) in the eclogitic matrix allows the calculation of further H2O(V)-independent equilibria; for example, the equilibrium
 | (5) |
 | (6) |
 | (7) |
Two equilibria of geothermometric importance were considered in addition to the above ones [equilibria (1) and (2b)]. These are
 | (8) |
), and
 | (9) |
In an attempt to better define the late stage of metamorphism for the Stümpelfelsen eclogite, we have also considered two equilibria that are dependent on aV. These are
 | (10) |
 | (11) |
Finally, it should be noted that the F contents of the OH-bearing minerals, amphibole, clinozoisite, micas, and talc are so low in the studied sample that we could ignore these. For calculations similar to those performed here but involving F-bearing phases the corresponding activity formulations can be enlarged (multiplied) by the term [OH/(OH + F)]Y with Y being the number of OH per formula unit considered in the activity model. For muscovite (or phengite) Y might only be unity even though there are 2 OH p.f.u. (see Massonne & Böving, 1990).
Results
Selected results of the application of the above geothermobarometers are given in Table 7. These results refer exclusively to the mineral compositions given in Tables 1–6. The geothermometric results using equilibrium (1) lie in the range 450–750°C. The calibration of Krogh Ravna (2000) and the thermodynamic calculation with Massonne's activity formulations [method (2) in Table 7; see IP1) gave nearly identical T results. The temperatures obtained with the calibration of Ellis & Green (1979) are about 60°C higher. The Fe2+–Mg exchange thermometer for the phengite–garnet pair [equilibrium (2b)] yielded also somewhat higher temperatures (see T quotations for IP2 in Table 7) compared with the calibration of equilibrium (1) by Krogh Ravna (2000). Assuming a 1
error as low as ±20°C for the above applied geothermometers [three calibrations of equilibrium (1), one calibration of equilibrium (2b)] and compositions of minerals (nearly) in equilibrium (e.g. GtC–OmB–PheB; see data in bold type in Table 7), the temperatures obtained for a single metamorphic stage are compatible with each other. Higher T also resulted from the Ti-in-garnet thermometry [equilibrium (8) applied to garnet A and D], which is, however, somewhat uncertain (1 error assumed to be ±60°C) because of the low Ti contents in our garnets. Temperatures about 100°C higher compared with Krogh Ravna's calibration were calculated on the basis of the Fe2+–Mg partitioning between garnet and talc [equilibrium (9)]. However, to obtain reliable results, this geothermometer has still to be calibrated. Our proposed activity model for Fe-talc is probably not sufficient to obtain better results than those obtained by applying the Ti-in-garnet thermometry. For reasons discussed in the subsequent section, we prefer the lower T estimates, which have been combined with the geobarometric calculations [methods (3)–(5), (7), and (11b) (aV for stage IIIa = 1) in Table 7]. The resulting P–T points, after selection of coexisting minerals with specific compositions, are plotted in Fig. 14. This selection was based on compositional combinations for method (2) [Fig. 13, equilibria (1)–(2c), etc.]. The results demonstrate that, for instance, the combination of GtA is less likely than GtB with OmA and PheA (see Table 7) because the IPs are closer together for the latter (the ideal case would be a coincidence of the IPs). A likely combination is also GtD, OmE and PheC (Stage IIb) with only a relatively small difference of the P–T positions of IP1 to IP3. In spite of the compositional selection, there is a significant scatter of P estimates but, nevertheless, systematic changes in P–T are discernible by relating the P–T data to a specific method. For instance, geobarometric methods (2) and (3) [plus method (1) for the T estimate] clearly indicate a P–T history involving slightly rising pressures with rising temperatures as the eclogite progressed from stage Ib to IIIa.
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The absolute errors of the various geobarometers are certainly significantly larger than suggested by the scatter of the calculated pressures with a specific geobarometer and somewhat variable mineral compositions. For the phengite barometers based on equilibria (2a) and (2c), the 1
accuracy is assumed to be in the range 1–2 kbar for the T interval 500–700°C despite the many experimental data on phengites in the systems K2O–MgO–Al2O3–SiO2–H2O and K2O–FeO–Al2O3–SiO2–H2O, as well as the evaluation of phengite–garnet parageneses of natural rocks (Massonne, 1997). However, a significant overestimation of the pressure using equilibrium (2a), as criticized by some petrologists, can be excluded. Evidence for this is, for example, the calculated pressure of 31 kbar at 620°C by Zhang & Liou (1994) using the above equilibrium and the composition of phengite inclusions in a garnet zone that also contained coesite relics. Thus, because of the coesite–quartz transition, the pressure can be overestimated by 4 kbar at most. For the Ti-in-phengite barometer, the 1 accuracy might be above 2 kbar (T > 500°C) because the underlying dataset (Massonne et al., 1993) considered no experiments but about 50 natural rocks (P, T, phengite composition). For the various barometers considering amphibole components, the accuracy is certainly not better than for the Ti-in-phengite barometry. |
DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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The above P–T estimates point to metamorphic pressures in excess of 22 kbar for the formation of the eclogitic matrix and thus broadly confirm the previous estimates by Massonne (1992)
, Schmädicke et al. (1992), and Klápová et al. (1998) for the low-T eclogites from the Saxonian Erzgebirge. However, more importantly, we were able to reconstruct a significant portion of the P–T path of the rock. This path comprises a T rise from stage Ib to IIIa of more than 200°C, independent of which geothermometric method was used (see previous section). As mentioned above, we prefer the lower T estimates yielding peak temperatures around 720°C (see P–T path in Fig. 14). We justify our selection by the argument that at temperatures somewhat above 700°C (and 25 kbar) melting of an H2O-saturated MORB protolith would occur [see summary by Poli & Fumagalli (2003); Lambert & Wyllie (1972) gave 735°C for a quartz eclogite + H2O in excess]. The infiltration of hydrous fluids with aV close to unity [see also below and Table 7: stage IIIa, equilibrium (11b)] at stage IIIa resulted in the formation of porphyroblasts, consisting of hydrous minerals such as amphibole, but not in melting. Another argument for the lower T estimates comes from the garnet–talc thermometry although yielding here higher T [method (9)]. If we compare KD [= (Fe2+/Mg)Gt/(Fe2+/Mg)Tc] values calculated from literature data on talc-bearing eclogites (Menot & Seddoh, 1985; Miller, 1986; Pognante, 1989; Zhang & Liou, 1994; Miller & Thöni, 1995; Zhang et al., 1995; Sisson et al., 1997; Liu et al., 1998; Hoschek, 2001) with our KD values [stage Ib (GtB, TcA) = 62·6; stage IIIa (GtE, TcC) = 19·6] temperatures below 700°C for our peak would result. This is because KD values close to 20, as calculated from the literature data, are related to metamorphic temperatures between 570°C (Miller, 1986) and 660°C (Zhang et al., 1995). However, a KD value of 99 (Miller & Thöni, 1995) cannot be assigned to the given peak T of 750°C as KD values related to Fe2+–Mg partitioning among mineral phases, generally, decrease with rising T. Probably, this example either refers to equilibration at T much lower than 750°C or represents a non-equilibrium assemblage. The KD values of 27 (Menot & Seddoh, 1985) and especially 55 (Pognante, 1989) confirm the early metamorphic T (stage Ib) of our eclogite, because the corresponding T estimates of those workers are 565°C and 500°C, respectively. The discrepancy between the result of geothermobarometric method (9) and the evaluation of literature data is probably due to the poorly known thermodynamic properties for the talc–Fe-talc join.
Phengite geobarometry with the components Mg–Al-celadonite [IP1 of method (2)], Fe2+–Al-celadonite [IP2 of method (2)], and Ti-muscovite [method (3)] yielded similar P results for each stage between Ib and IIIa. We noted that the P–T data of IP2 and method (3) [combined with method (1)] are systematically about 1 kbar and 2 kbar, respectively, above those of IP1. A slight P overestimation by method (3) has already been reported by Hoschek (2001). Barometry applying the glaucophane breakdown [equilibrium (4b,c)] yields almost the same results as that related to IP1 but only when the modified glaucophane activity formulation is used. Because the old model gives significantly lower P compared with phengite barometry, and because there are good thermodynamic reasons to modify the old model (see above), we suggest the application of the new model for future geobarometric evaluations. The corresponding barometry with tremolite [equilibrium (4a)] instead of glaucophane leads to calculated pressures that are lower, on average, by about 3 kbar in comparison with the phengite barometry with the Mg–Al-celadonite component applied here. This may indicate that the tremolite activity formulation has to be revised in future work. The geobarometric application of the equilibrium involving tremolite with clinozoisite, muscovite, and Mg–Al-celadonite [equilibrium (5)] also yields somewhat lower pressures (2·5 kbar) compared with method (2) with the Mg–Al-celadonite component, and a clear tendency of increasing difference with rising T is discernible. The lowest P estimates were obtained with equilibrium (7) involving tremolite and clinozoisite. As the activity formulations for these components in amphibole and clinozoisite are only estimates, we prefer, for the construction of the P–T path for the Stümpelfelsen eclogite, phengite barometry involving the Mg–Al-celadonite component. An additional argument in support of this preference is that there are numerous underlying experimental data on phengites (see Massonne & Szpurka, 1997) that provide a good basis for the activity model for components in phengite. As a consequence of this, the P–T path from stage Ib to IIIa is displayed in Fig. 14 located almost along the quartz–coesite transition curve, which is slightly below the P determinations using phengite barometry with the Mg–Al-celadonite component. This is the result of the consideration of the barometry applied here using amphibole components yielding lower pressures. Moreover, no coesite was detected in the Stümpelfelsen eclogite. Thus, we suggest that the formation of the matrix minerals in this eclogite (without stage Ia) occurred during slight P rise (at P > 24 kbar), and a clear T rise of more than 200°C, reaching peak T conditions of about 720°C. We estimate the error of this early portion of the P–T path to be below 2 kbar and 30°C (see also previous section).
The P–T path from stage Ib to IIIa could roughly mark the lower P stability for talc in eclogites of normal MORB composition, as the breakdown product towards lower pressures, amphibole, is already present in the Stümpelfelsen eclogite. Thermodynamic calculations for simple model systems related to metabasic rocks, such as the system Na2O–FeO–MgO–Al2O3–SiO2–H2O (NFMASH), demonstrate the existence of a stability field for talc + garnet + omphacite (Guiraud et al., 1990). In the above system, this field is limited towards lower pressures at about 32 kbar (T limit 600°C) according to the calculations of Guiraud et al. (1990). However, the oxide component CaO in addition to NFMASH (jadeite 
omphacite) will enlarge the P–T range of this assemblage. Unfortunately, the stability fields of talc + phengite (Massonne & Schreyer, 1989) and talc + zoisite (Hoschek, 1995) do not provide further limits on P here. However, the latter mineral assemblage, found in the Stümpelfelsen eclogite, is an additional argument for the preferred peak T as zoisite + talc breaks down at about 730°C, 25 kbar.
The natural occurrences of talc in eclogite seem to be broadly in accordance with the suggested lower P stability limit. The P estimates of Zhang & Liou (1994), Miller & Thöni (1995), Zhang et al. (1995), Liu et al. (1998) and Zack et al. (2001) are above 26 kbar and consistent with the presence of coesite relics. The estimates of Sisson et al. (1997) and Hoschek (2001) are about 22 kbar. The pressures (<20 kbar) given in the earlier studies by Menot & Seddoh (1985), Miller (1986) and Pognante (1989) tend to be minimum pressures, and the real metamorphic pressures could have been higher. The talc compositions reported in the above studies are broadly similar to those we have observed (see Fig. 9). In the study by Sisson et al. (1997) the FeO and Al2O3 contents of the talc are virtually the same as those reported here. In the other studies, the FeO contents of talc are lower but often above 2 wt %. Al2O3 contents are somewhat higher (mostly above 0·5 wt %); however, kyanite was reported coexisting with talc, except in the studies of Menot & Seddoh (1985), Pognante (1989) and Sisson et al. (1997), with Al2O3 contents in talc less than 0·5 wt %. The absence of kyanite in our sample is also corroborated by the thermodynamic calculation of the equilibrium
 | (12) |
The determination of the P–T conditions of the early metamorphic stage Ia (garnet core composition GtA) was not successful. However, we propose three possibilities for an early path to stage Ib in Fig. 14. Although thermometry with garnet core composition GtA yielded relatively high temperatures (see Table 7), the path along a (subduction zone) geotherm of 6°C/km (see Fig. 14) might be most likely to reach P–T conditions close to 25 kbar and 480°C.
For stage IIIb geothermobarometric methods (3), (5) and (6) alone yield a P–T estimate in the range around 450°C and 21 kbar. However, because paragonite is present at this stage the following equilibria can be additionally considered:
 | (11b) |
 | (13) |
 | (14) |
). Geohygrometry applied to this eclogite by Massonne (1992) yielded a water activity close to 0·7 for the peak conditions. The same method gave values for aV above unity for GtE, OmE, and PheD coexisting at IP1 (27·6 kbar, 703°C) and a tremolite activity of 0·372 (AmphD, 700°C). A calculated water activity close to unity is reached at significantly lower tremolite activities. A lower tremolite activity would also lead to higher pressures, when applying method (4a). Thus, we take both as a hint of lower tremolite activities. However, more importantly, we can now conclude that amphibole blastesis was initiated at both eclogite occurrences at the peak T condition through the infiltration of hydrous fluids (aV close to unity) accompanied by decompression of the rock and a significant fall in T. These fluids were then consumed by the formation of the amphibole, clinozoisite, and phengite porphyroblasts. As a consequence of this, the water activity decreased (see also Massonne, 1992). When relatively low temperatures were reached and water was consumed, no further (significant) change of the rock could probably occur, thus explaining the extraordinary freshness of the eclogites from the Stümpelfelsen. The significant (and fast?) T fall just after the T climax could perhaps also explain why the chemical zonations of the early formed matrix minerals are relatively well preserved on the micrometre scale. |
CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLE...MINERAL COMPOSITIONSGEOTHERMOBAROMETRYDISCUSSIONCONCLUSIONSREFERENCES |
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The P–T path derived here for the Stümpelfelsen eclogite, and the previously reported path of the eclogite near Marienberg (Massonne, 1992
), demonstrate that a simple P–T loop, as suggested by Schmädicke et al. (1992) for the eclogites from the southwestern Erzgebirge, is only a very rough approach to the true P–T evolution of eclogites from the MEU (Fig. 15). This may even be a general problem because many P–T paths for eclogites in the literature are based only on peak P–T conditions derived by geothermobarometry. For a better understanding of the mass flow in subduction zones, and the corresponding geodynamic processes, detailed P–T paths for subducted material as derived here or, for example, by Zhang & Liou (1994), Massonne (1995a) and Hoschek (2001), are required. Numerical modelling of mass flows in subduction channel environments (e.g. Gerya et al., 2002) can result in two types of paths. One is a clockwise P–T loop. The other type is characterized by an isobaric segment at the highest P conditions. Such results must be compared with P–T paths derived for actual rocks. For instance, cooling is related to the previously mentioned isobaric segment in the modelling experiments of Gerya et al. (2002), whereas our finding is significant heating at the highest pressures. Heating is explained by the attachment of subducted eclogitic material to the base of the hot overriding plate bounding the subduction channel. At this stage, hydrous fluids (either from below or released by heating the rock unit itself) interact with the eclogitic material to form porphyroblasts of hydrous minerals (amphibole, clinozoisite, phengite). Whether the formation of the porphyroblasts can cause the exhumation of the rock is not clear. Clearly, the (partial) transformation of the dense phases garnet and omphacite to less dense hydrous phases, mainly amphibole, results in a reduction in the density of quartz eclogite. By this process, densities of eclogite even lower than that of garnet peridotite could be reached, creating buoyancy forces that could, at least, facilitate the exhumation of portions of deeply subducted oceanic crust (Massonne, 1996). The exhumation of eclogitic material studied here could be caused by an upward directed mass flow in a subduction channel. In the corresponding P–T environment, it is possible to cool rock units significantly (in our case: 
T > 200°C) at high pressures (P > 15 kbar) by continuing subduction of a crustal plate.
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ACKNOWLEDGEMENTS |
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Our thanks are due to Deutsche Forschungsgemeinschaft for financial support of our study ‘Eklogite im Erzgebirge’ (MA 1160/3, KO 1187/2). Heinz-Jürgen Bernhardt provided the computer program XMAP for visualization of area studies with the wavelength-dispersive system of an electron microprobe. Simon Harley, Gert Hoschek, Grahame Oliver, and Arne Willner reviewed the manuscript. It benefited considerably from their comments and suggestions.
* Corresponding author. Telephone: ++49-711-121-1225. Fax: ++49-711-121-1222. E-mail: h-j.massonne{at}mineralogie.uni-stuttgart.de
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