Pohorje eclogites revisited: Evidence for ultrahigh-pressure
metamorphic conditions
Ultravisokotlačni metamorfizem pohorskega eklogita
Mirijam VRABEC
University of Ljubljana, Faculty of Natural Sciences and Engineering, Department of Geology, Aškerčeva 12, SI-1000 Ljubljana, Slovenia; e-mail: mirijam.vrabec@ntf.uni-lj.si
Prejeto / Received 4. 3. 2010; Sprejeto / Accepted 20. 5. 2010
Key words: ultrahigh-pressure metamorphism, eclogite, geothermobarometry, collisional orogen, Pohorje, Eastern Alps, Slovenia
Kljucne besede: ultravisokotlačna metamorfoza, eklogit, geotermobarometrija, kolizijski orogeni, Pohorje, Vzhodne Alpe, Slovenija
Abstract
Kyanite eclogites from the Pohorje Mountains, Slovenia, are providing the first evidence of ultrahigh-pressure Eo-Alpine metamorphism in the Eastern Alps. Polycrystalline quartz inclusions in garnet, omphacite and kyanite are surrounded by radial fractures and exhibit microtextures diagnostic for the recovery after coesite breakdown. The non-stoichiometric supersilicic omphacites found in Pohorje eclogites contain up to 5 mol % of Ca-Eskola molecule. Such clinopyroxenes are known to be stable exclusively at high-pressure conditions exceeding 3 GPa. Their breakdown during decompression resulted in exolution of quartz rods and needles that are oriented parallel to omphacite c-axis. The absence of coesite is a consequence of near-isothermal decompression during the first stages of exhumation.
Pressure and temperature conditions for the formation of the peak metamorphic mineral assemblages have been assessed through a consideration of a) Fe2+-Mg partitioning between garnet and omphacite pairs, based on different calibrations; b) the equilibrium between garnet + clinopyroxene + phengite ± kyanite ± quartz/coesite assemblage. Estimated peak pressure and temperature conditions of 3.0-3.1 GPa and 750-783 °C are well within the coesite, i.e. the ultrahigh-pressure stability field.
Izvleček
Pohorski kianitovi eklogiti predstavljajo prvi dokaz za obstoj eo-alpinske ultravisokotlačne metamorfoze v Vzhodnih Alpah. Radialne razpoke okoli polikristalnih kremenovih vključkov v granatu, omfacitu in kianitu ter njihove specifične mikrostrukture pričajo o obstoju coesita, ki je med ekshumacijo zaradi dekompresije prešel v kremen. Popolna odsotnost coesita je posledica izotermne dekompresije v začetni stopnji ekshumacije pohorskih eklogitov. Nestehiometrični omfaciti z visoko vsebnostjo SiO2, ki vsebujejo do 5 mol % Ca-Eskola molekule in so obstojni izključno pri tlakih večjih od 3 GPa, potrjujejo izpostavljenost pohorskih eklogitov ultravisokotlačnim metamorfnim pogojem. Zaradi dekompresije so se v njih izločile kremenove iglice in paličice, ki so orientirane vzporedno z omfacitovo c-osjo.
Tlačni in temperaturni pogoji nastanka pohorskih eklogitov so bili določeni s pomočjo različnih geotermometrov, ki temeljijo na izmenjavi Fe2+ in Mg ionov med granatom in omfacitom, kot tudi na osnovi ravnotežja med mineralnimi fazami: granat + monoklinski piroksen + fengit ± kianit ± kremen/coesit. Izračunane tlačne in temperaturne vrednosti se gibljejo v razponu od 3.0-3.1 GPa in 750-783 °C ter odgovarjajo ultravisokotlačnemu (coesitovemu) stabilnostnemu območju.
Introduction
Ultrahigh-pressure (UHP) metamorphism is an important type of orogenic metamorphism that has been recognized in many Phanerozoic collision belts (e.g. Liou et al., 1998; Chopin, 2003, and references therein). Well investigated intracrato-nic collisional orogens that exhibit scattered effects of subsolidus UHP recrystallization include the Quinling-Dabie-Sulu belt of east-central China, the Kokchetav Complex of northern Kazah-
stan, the Dora Maira massif of the Western Alps, and the Western Gneiss Region (WGR) of Norway (Liou et al., 1994; Coleman & Wang, 1995; Ernst et al., 1995). This four classic and several other UHP terranes (Figure 1) share common structural and lithological characteristics (Liou, 2000). Su-pracrustal rocks of these UHP regions experienced subduction-zone metamorphism at mantle depths, followed by a retrograde amphibolite-granulite facies overprint during exhumation, and finally thermal recrystallization and defor-
Figure 1. Distribution of recognized UHP metamorphic terranes in the world through time and space. So far, evidence of UHP metamorphism was not found only in Australia and North America (modified from Liou et al., 1998; Chopin, 2003).
mation during postcollisional granitic intrusions and orogeny.
With the discovery of UHP metamorphism geologists realized that, contrary to general belief, continental crust in convergent settings may be subducted to enormous depths, and that the most formidable geodynamic problems concerning UHP metamorphic rocks are not the mechanisms of their deep burial, but the mechanisms which facilitated their subsequent exhumation to the Earth's surface without a complete breakdown of the UHP mineral assemblages. From the recurrent occurrences both in time and space, since late Proterozoic, it is clear that UHP metamorphism is a common process, inherent to continental collision.
Ultrahigh-pressure metamorphism is synonymous with eclogite-facies metamorphism that has occured within the stability field of coesite (Figure 2). Unequivocal identification of UHP conditions depends on the presence of relict coesite or diamond, high-pressure polymorphs of silica and carbon, as direct indicators of metamorphic pressures of at least 3 GPa (coesite) or 4 GPa (diamond). But since the metastable preservation of relict UHP phases during exhumation and decompression is now known to be very rare, a simple microscopic identification of UHP metamorphic rocks is normally not possible, and the evidence for UHP conditions must be deduced from indirect petrographic and microtextural observations. In absence of an actual coesite relict, polycrystal-line quartz aggregates are strongly indicative tex-
Temperature (°C)
Figure 2. Simplified P-T facies diagram. UHP eclogites are defined as metabasic rocks with a dominant plagioclase-free, eclogite-facies mineral paragenesis of garnet and omphacitic clinopyroxene, but with additional petrographic or mineral-chemistry features that indicate equilibration at pressures within the coesite P-T stability field. Subdivision of the eclogi-te facies field is based on Okamoto & Maruyama (1999). Also shown are the stability fields for diamond and coesite according to Holland & Powell (1998).
tural feature of former coesite existence (Smith, 1984; Gillet et al., 1984) but tend to disappear due to recrystallization during prolonged thermal annealing. A further distinctive petrographic feature of partly or completely replaced coesite inclusions is the developement of radial expansion cracks, extending from the inclusion boundary into the enclosing mineral. This reflects the roughly 10 % volume increase on transition from coesite to a-quartz. Several other uncommon petrological features and assemblages that have been reported from various UHP metamorphic rocks provide additional evidence for UHP metamorphism. Quartz rods have been observed in omphacite from eclog-ites of several UHP terranes. In all cases, SiO2 needles and rods in omphacite have been interpreted as exsolution products from a preexisting super-silicic clinopyroxene that contained excess silica at peak metamorphic conditions (e.g. Smith, 1984; Liou et al., 1998; Katayama et al., 2000; Schmädicke and Müller, 2000; Zhang & Liou, 2000; Dobrzhi-netskaya et al., 2002). Potassium-bearing clinopy-roxene with extremely high potassium content (up to 1.5 wt% K2O) is stable at pressures of 410 GPa. During decompression, potassic clinopy-roxene develops characteristic textures with oriented precipitates of K-feldspar (Sobolev & Shat-sky, 1990). Orthopyroxene exsolutions in garnet require formation pressures in excess of 6 GPa and hence over 200 km of depth. Exsolutions suggest the existence of a super-silicic precursor garnet with several mol % of majorite component (van Roermund et al., 2001). a-PbO2-type TiO2 inclusions in garnet indicate achieved pressures in the range from 4.5 to 6.5 GPa at a temperature of 1000 °C (Huang et al., 2000). Although the indirect indicators of UHP metamorphism are very useful since the preservation of metastable coesite and diamond is very rare, they cannot be used alone as a proof that UHP metamorphism was achieved.
When direct mineral indicators are absent, converging indirect pieces of evidence, together with reliable geothermobarometrical calculations, are needed to verify the existence of UHP metamorphism.
Pohorje in north-eastern Slovenia (Figure 3) is the south-easternmost prolongation of the Eastern Alps. It is a part of the extensive Alpine oro-gen where UHP metamorphism was documented in the Western Alps (e.g. Chopin, 1984) and was shown to be related to the Tertiary orogeny (e.g. Tilton et al., 1991). Metamorphic processes related to the older, Cretaceous Alpine orogeny are mainly recognized in the Austroalpine units of the Eastern Alps (e.g. Thoni & Jagoutz, 1992). So far, up to high-pressure (HP) eclogite facies metamor-phic conditions were recognized in the Koralpe and Saualpe areas situated just north of Pohorje (Miller, 1990). The Pohorje Mountains consist of a stack of Cretaceous Austroalpine nappes, predominately composed of micashists, gneisses and amphibolites, but also include several lenses of eclogitic rocks that are of special interest since they have high preservation potential for ultra-high-pressure metamorphic indicators. Eclogites from Pohorje were previously investigated by Hin-terlechner-Ravnik (1982), Hinterlechner-Ravnik et al. (1991) and Koch (1999). Geothermobarome-tric estimations from the first two works are rather broad, with estimated pressure ranging from 1.21.8 GPa at temperature from 460-900 °C. Pressure and temperature estimates by Koch (1999) fall into the same (high-pressure) range, but are more narrowly constrained to 1.5 GPa at 760 °C.
First two samples of eclogites from Pohorje indicating possible UHP conditions were investigated by Janak et al. (2004). This work presents new samples from several new localities bringing undisputed mineralogical, petrological, microtextural and microchemical evidence for ultra-
Figure 3. Simplified geologic map of Pohorje and adjacent areas (modified from Mioc & Znidarcic, 1977) showing locations of the investigated eclogite samples.
high-pressure metamorphism of eclogites in the Austroalpine units of the Eastern Alps, exposed in the Pohorje Mountains of Slovenia. The evidence for ultrahigh-pressure conditions is strongly supported by extensive and precise geothermoba-rometric calculations based on different widely accepted calibrations.
Methods
Electron Probe Micro-Analysis
Representative microchemical analyses of the main constituent minerals were determined by EPMA technique using a CAMECA SX-100 electron microprobe at Dionyz Stur Institute of Geology in Bratislava. Bombarding of micro-volumes of sample with a focused electron beam (5-30 keV) induced emission of X-ray photons. The wavelengths of collected X-rays were identified by recording their WDS spectra (Wavelength Dispersive Spectroscopy). Analytical conditions were 15 keV accelerating voltage and 20 nA beam current, with a peak counting time of 20 s and a beam diameter of 2-10 ^m. Raw counts were corrected using a PAP routine.
Garnet-clinopyroxene Fe2+-Mg exchange geothermometery
Temperature conditions of metamorphism were obtained using the partitioning of Fe2+ and Mg between co-existing garnet and clinopyroxene (omphacite). Due to the common appearance of clinopyroxene and garnet in a mineral assemblage of high-grade metamorphic rocks of basic and ultrabasic composition this is one of the most widely used methods in geothermometry of such rocks. When these two minerals are contiguous phases, they effectively exchange the two elements and the exchange balance is a function of temperature.
The exchange of iron and magnesium between clinopyroxene and garnet is represented by exchange reaction:
pyrope + 3 hedenbergite —> almandine + 3 diopside
The equilibrium constant for the considered equilibrium is expressed by the following function:
n3
Km =
a
"Fe aMg
acpx UMg
ncpx "Fe
where a{ is the activity of component i in phase j. If the minerals are ideal solid solutions, activities are equivalent to concentrations and
vgrt	vcpx
„ _ A. Fe	A- Mg
e9~ Vgrt	y cpx '
■A- Mg	F*
■Fe
{Fe2+/Mg) {Fe2+/Mg)
grt
cpx
=K,
D
where KD is the distribution coefficient, Xf£ is the mole fraction of Fe2+ in the three equivalent divalent sites in garnet structure, X^ is the mole
fraction of Fe2+ in the clinopyroxene, etc. If the minerals are not ideal (i.e. their compositions differ from those of pure ideal end-members used in experiments) compositional differences are corrected by the introduction of activities, which express the thermodynamically effective concentrations of components. Then
a ■
-X-Y
where y is defined as the activity coefficient and thus
Y Xcpx	Vcpx
X — Fe , MZ ,'Fe > Mg _ £ g eq Y grt y cpx grt cpx D y AMg ^Fe ¡Mg IFe
Temperatures have been estimated using six different quantitative calibrations of the garnet-clinopyroxene system revealing garnet-clinopy-roxene Fe2+-Mg exchange geothermometers of Ellis & Green (1979), Powel (1985), Krogh (1988), Pattison & Newton (1989), Ai (1994) and Krogh Ravna (2000).
Ellis & Green (1979) determined the distribution coefficient as a function of P, T and Ca-con-tent in garnet (Xc£) and derived the empirical relation:
T EG-7Á O — "
3104 -XZ. + 3030 + 10.86 -P(kb)
Ca_
In Kd +1.9034
-273
where XqJ is defined as

Ca
Ca + Mg + Fe
2+
They have shown that KD is apparently independent of the Mg/(Mg + Fe2+) content in clinopyroxene and garnet, but that there is a marked dependence of Kd upon the Ca-content of garnet. This Ca-effect is believed to be caused by a combination of non-ideal Ca-Mg substitutions in garnet and clinopyroxene. Consequently, a rectilinear correction for XcÜ in garnet was proposed.
Powell (1985) made an upgrade to Ellis & Green (1979) geothermometer defining:

2790 +10 • P(kb) + 3140 -Xfa
hvKD +1.735
--273
which gives slightly lower temperatures than calibration of Ellis & Green (1979).
Krogh (1988) suggested a curvilinear relationship between lnKD and Xg! in garnet:
Tg.-88( O -
1879 + 6731 ■ Xfa - 6173 • (X^a f +100 • P{GPd)
_Ca -
\aKD +1.393
— 273
at least for the compositional range = 0.100.50. The Ca-content in garnet was calculated as:
j^grt _	Ca
Ca + Mn + Fe2+ + Mg
Calculated temperatures do not vary with the Mg/(Mg + Fe2+) content in garnet and Na-content in the clinopyroxene. Temperatures below 900°C are a bit lower than those obtained by the method
of Powell (1985), and the difference is larger for lower temperatures and lower values of Xc£.
Pattison & Newton (1989) performed multiple regression of a large set of data on the Fe2+-Mg equilibrium between garnet and clinopyroxene resulting in the following relationship:
1PN-
qOQ =
561+3395 • Jg - 2388 • (Xg)2 +9781 • X%g - 31026 • (X
hKD +0.512 mn-(xfjf+m.i-P{GPa) In Kd +0.512
which includes the curvilinear corrections for XqJ and also for Mg-number (Mg#) in garnet:
Mg
Mg#= 100- XMg =100
Mg + Fe
2+
This thermometer works well with experimental data of Pattison & Newton (1989) but commonly yields unrealistically low temperatures for natural rocks. An important feature discovered by Pattison & Newton (1989) and later supported by other researchers (e.g. Ai, 1994; Berman et al., 1995) is that Kd decreases with decreasing Xf?g at all temperatures.
Ai (1994) investigated about 300 garnet-clinopy-roxene pairs and by multiple regression arrived to the expression:
1 A-94
CC) =
_ 1987 + 3648.55 • Xfa -1629 • (Xfa f - 659 ■ X
In Kd +1.076
| 176.6 P(GPa) 2?3
\tîKd +1.076
with curvilinear correction for Xc£ and rectilinear correction for XfJg in garnet. Ai's thermometer is suitable especially for the systems with low-Ca and high-Mg garnets.
In addition to significant dependence between the distribution coefficient KD and XqJ and Xf?g, Krogh Ravna (2000) incorporated the effect of Mn-content in garnet Xj£,:
vgrt Mn
Mn
Ca + Mn + Fe2+ + Mg
and proposed the following P-T-compositional relationship:
1KR-m
(°C):
1939 + 3270 • Jg -1396 • (Jg)2 + 3319 • Xfc - 3535 • (X^f ln^ +1.223
1105-J^ - 3561 ■ (XS)2 + 2324 • {Xflf +169.4■ P(GPa)
ln^ +1.223
-273
which confirmed his conclusion from 1988 that the Fe2+-Mg equilibrium between co-existing garnet and clinopyroxene is independent of the variations in the Na-content of the clinopyroxene, at least in the Xns range from 0 to 0.51. This means
that the jadeite content in clinopyroxene has a certain influence on calculated temperatures only at higher X^a values. Since the natural jadeite content of eclogites is commonly lower than the recommended value, Na is not a problematic additional component in the garnet-clinopyroxene exchange thermometer for eclogites. Application of this thermometer gives reasonable results for most compositional ranges covered by garnet-clinopyroxene pairs from natural rocks.
The uncertainty of the garnet-clinopyroxene geothermometer is usually estimated to ±30 to 50 °C and in most cases almost all of the above-mentioned geothermometers, give temperatures within this interval. However, there is one important drawback that should always be considered when using Fe2+-Mg exchange geothermometers. The problem is that generally the data is only available for total iron content (Fetot), therefore it is not known how much garnet and clinopyrox-ene iron is present in the exchanging ferrous (Fe2+) and non-exchanging ferric (Fe3+) states. There are several methods to calculate the Fe3+/Fe2+ ratio in garnet and clinopyroxene (charge balance criteria, equalizing the amount of Fe3+ with the Na excess over (Altot+Cr), standard titration method, etc.) but they are unfortunately very sensitive to analytical errors and not always reliable.
Garnets are less of a problem than clinopy-roxenes because of their customary higher Fetot contents and much lower Fe3+/Fe2+ ratios (Car-swell et al., 1997). In Fe-rich garnets all iron can be treated as ferrous without affecting the calculated temperatures significantly. But in Fe-poor garnets the underestimation of Fe3+ will result in higher KD value, and consequently, in underestimated temperatures (Krogh Ravna, 2000). In common eclogites the Fe3+/Fetot ratio is reported to be low, in the range of 0.0-0.06 (Carswell et al., 2000; Schmid et al., 2003) and can be therefore calculated by using stoichiometric charge balance.
Clinopyroxenes are more problematic, especially because they tend to be non-stoichiometric under HP/UHP conditions due to the presence of the Ca-Eskola molecule. The most problematic are Fe-poor clinopyroxenes which show a large spread in calculated Fe3+/Fetot ratios. The charge balance calculations are unsuitable in this case. The published Fe3+/Fetot ratios from omphacites vary between 0.0 and 0.5 (Carswell et al., 1997) and were proven by Mossbauer studies (Carswell et al., 2000) and micro-XANES (Schmid et al., 2003).
For eclogitic rocks from Pohorje, the Fe3+/ /Fetot ratio directly determined by Mossbauer spectroscopy is reported to vary between 0.15 to 0.41 in omphacite, with a mean of 0.30, and tends to be low and constant in garnets, with the value of 0.02-0.03 (Koch, 1999).
Geobarometry based on garnet-clinopyroxene-phengite-kyanite-quartz/coesite assemblage
Reliable geobarometers applicable to HP and UHP metamorphic rocks are rather scarce. In eclogites containing an assemblage of garnet +
clinopyroxene + phengite ± kyanite ± quartz/coe-site, an equilibrium between these phases may be used for thermobarometric estimations. Such geobarometer is based on net-transfer reactions representing a balanced reactions among phases (or components of phases) in which progress of the reactions result in a change in the modal amounts of the phases. This means that net-transfer reactions cause the production and consumption of phases and therefore result in large volume changes making the equilibrium constants pressure sensitive.
Possible net-transfer reactions defining this equilibrium are given as follows (Krogh Ravna & Terry, 2001; Krogh Ravna & Paquin, 2003):
[phe]: 3 diopside + 2 kyanite —>
1	grossular + 1 pyrope + 2 coesite/quartz [di, grs]: 1 pyrope + 3 muscovite + 4 coesite/quartz
3 Al-celadonite + 4 kyanite [prp]: 3 muscmite + 3 diopside + 2 coesite/quartz —>
3 Al-celadonite + 1 grossular + 2 kyanite [Si02,ky]: 6 diopside + 3 muscovite —>
2	grossular + 1 pyrope + 3 celadonite
These reactions define an invariant point in both the coesite and quartz stability field, depending on which SiO2 polymorph is stable. Phases in square brackets are absent in the reactions.
For pressure calculations the calibrations of
Waters & Martin (1996), Krogh Ravna & Terry (2001) and Krogh Ravna & Terry (2004) have been applied.
Waters & Martin (1996) calibrated geobarometer aplicable to HP and UHP eclogites with the garnet + clinopyroxene + phengite assemblage, using the reaction [SiO2, ky] and thermodynamic data set of Holland & Powell (1990):
PWM-96 War) = 285 + 0.02044• T - 0.003539-T-\nK
where lnK term is calculated as follows: = 61nad. -Inaprp ~2\nagrs +31nainvphe The phengite activity may be calculated from:
_ aidealms _ ^AIM, ' ^AIT, ainvphe ~	~ y	y
"ideal eel MgM, ' ^ SiTt
with XA1Ti=4-Si, XSiTi = Si-2 and Mi the octahe
dral cation sites. Activity model for diopside was taken from Holland (1990) and for garnet from
Newton & Haselton (1981).
Krogh Ravna & Terry (2001, 2004) used all four above reactions and constructed geobarometric expressions for UHP (with coesite) and HP (with quartz) assemblages. The corresponding equilibrium constants are:
ngrt _ grl I coe/qkf	I phe ^ ( h V
g\phe,coe/qtz] _ UFP Srs rfl02 > ^\di,gn,coe/qtz]_ \"cd I \uAl2SiOsl
glmcoe/qtz] _ "pi ' K^f ' (aAl2SioJ	=a%'\a&J' \acdJ
" krfki
For the above reactions they formulated linearized barometric expressions, which are:
P^T](GPa) = 7.235 - 0.000659 T+0.001162-\nK^q,z] P^^GPa) = 11.422 - 0.001676 • T + 0.002157 ■ T - In K^"^ P[ÎZT](GPa) = -2.624 + 0.005741 • T + 0.0004549 • T ■ In K{di'ps'qtz] P[ÎrT'](GPa) = "°-899+0.003929 • T + 0.0002962 T-ln K^"^ P^f](GPa) = 0.355 + 0.003695-T+0.0003059 -T ■ {nK^"^ P^e\GPa) = 0.568 + 0.003345 • T + 0.0002705 • T ■\RKbrpM Pl°^\GPa) = 1.801 + 0.002781 • T + 0.0002425 • T ■ In ^
The intersection of "quartz absent" lines defines a single point within the coesite (UHP) stability field, and analogously, the intersection of "coesite absent" lines defines another single point within the quartz (HP) stability field (Figure 4). Therefore the intersection of any two of these sets of lines uniquely defines P and T for a single sample.
5,0-1-
4,5
1,5
1,0-1-1-i->-,-
500 600 700 800 900 1000
Temperature fC)
Figure 4. The intersection of reaction lines in the quartz and coesite stability fields (after Krogh Ravna & Terry, 2004). Phases in square brackets are absent in the reactions.
The content of ferric iron was calculated as:
Fe3+ =2-(Al + Cr + Ti)
The uncertainty limits for this thermobarome-ter are ±65°C and ±0.32 GPa. These geothermoba-rometric methods are supposedly less affected by subsequent thermal re-equilibration than common cation exchange thermometers, and the methods
also diminish the problems related to estimation of Fe3+ in omphacite (Krogh Ravna & Paquin, 2003). Activity models for phengite, clinopyroxene and garnet were taken from Holland & Powell (1998), Holland (1990) and Ganguly et al. (1996), respectively.
For calculating peak metamorphic conditions for a specific eclogite sample, garnet with maximum	omphacite with minimum a^x (and thus maximum jadeite content) and phengite
with maximum a^f (maximum Si content) are required. Analyses of phases used for geothermo-barometryare given in Table 1.
matrix often accompanied with elongated grains of blue kyanite.
The eclogites consist of garnet, omphacite, kyanite, and zoisite as major primary mineral phases. In some samples crystals of phengitic mica, quartz (after coesite?), rutile and rarely zircon are also present. Among the secondary mineral phases are mainly amphibole, diopside, plagioclase, biotite, sapphirine, corundum and spinel. They occur in the coronas, symplectites and fractures of the primary minerals. Secondary minerals developed after peak metamorphic conditions and are related to the exhumation of these rocks.
Table 1. Representative microprobe analyses of mineral compositions used for thermobarometry.
Sample	JV103a	NO1/04A	PO6h	SP1/08	JV103a	NO1/04A	PO6h	SP1/08	JV103a	NO1/04A	PO6h	SP1/08
Mineral	Grt	Grt	Grt	Grt	Cpx	Cpx	Cpx	Cpx	Phe	Phe	Phe	Phe
SiO2	40.99	40.80	40.83	40.35	55.02	5.5	55.28	54.92	53.36	51.53	51.00	52.66
TiO2	0.01	0.02	0.02	0.00	0.11	0.12	0.14	0.10	0.02	0.00	0.74	0.02
Al2O3	23.16	22.99	23.06	22.58	8.34	8.03	8.04	7.95	26.87	26.84	27.59	27.05
&2O3	0.00	0.14	0.04	0.27	0.02	0.19	0.18	0.06	0.30	0.37	0.08	0.05
FeO	15.77	13.92	13.91	14.88	2.55	2.45	2.48	2.28	1.48	1.71	1.03	1.20
MnO	0.38	0.26	0.28	0.34	0.02	0.02	0.00	0.06	0.00	0.01	0.00	0.00
MgO	13.37	15.02	14.22	12.91	12.30	12.43	12.71	12.08	3.62	4.00	4.22	3.90
CaO	7.81	7.42	7.76	7.60	18.19	18.83	18.80	18.22	0.11	0.08	0.00	0.16
Na2O	0.01	0.02	0.00	0.03	3.73	3.69	3.75	3.75	0.06	0.08	0.32	0.06
K2O	nd	nd	nd	nd	bd	bd	bd	bd	9.11	10.06	10.46	9.50
Total	101.50	100.59	100.12	98.97	100.28	100.81	101.38	99.42	94.93	94.68	95.44	94.60
Si	2.978	2.960	2.984	3.005	1.958	1.954	1.949	1.970	3.510	3.431	3.377	3.482
Ti	0.001	0.001	0.001	0.000	0.003	0.003	0.004	0.003	0.001	0.000	0.037	0.001
Al	1.983	1.966	1.987	1.983	0.350	0.336	0.334	0.336	2.083	2.107	2.154	2.108
Cr	0.000	0.008	0.002	0.016	0.001	0.005	0.005	0.002	0.016	0.019	0.004	0.003
Fe3+	0.061	0.106	0.039	0.000	0.000	0.000	0.011	0.000	0.000	0.035	0.000	0.000
Fe2+	0.897	0.739	0.811	0.931	0.076	0.073	0.062	0.068	0.081	0.060	0.057	0.066
Mn	0.023	0.016	0.017	0.021	0.001	0.001	0.000	0.002	0.000	0.001	0.000	0.000
Mg	1.447	1.624	1.549	1.433	0.652	0.657	0.668	0.646	0.355	0.397	0.416	0.385
Ca	0.608	0.577	0.608	0.607	0.694	0.716	0.710	0.700	0.008	0.006	0.000	0.011
Na	0.001	0.003	0.000	0.005	0.258	0.254	0.257	0.261	0.008	0.010	0.041	0.008
K	0.000	0.000	0.000	0.000	0.000	0.000	0.000	0.000	0.765	0.855	0.884	0.801
Total	7.999	8.000	7.998	8.001	3.993	3.999	4.000	3.988	6.827	6.921	6.970	6.866
Analyses oxygens.
(in wt%) of garnet (Grt), clinopyroxene (Cpx) and Abbreviations are as follows: bd - below detection:
phengite (Phe). Garnet is normalized to 12, , nd - not determined.
clinopyroxene to 6 and phengite to 11
Results
Petrography and mineral chemistry
Eclogites from four localities in the Pohorje Mountains have been investigated (Figure 3). The dominant rock type is weakly retrograded eclogite, which occurrs in bands, lenses and boudins within the surrounding continental crustal rocks (orthogneisses, paragneisses and micaschists). In macroscopic scale, eclogites contain big distinctive grains of garnet surrounded by omphacite
Garnets form euhedral to subhedral crystals with size ranging from large porphyroblasts up to several mm in diameter, to very tiny inclusions of few microns, which are found within kyanite and omphacite crystals. Garnets are texturally uniform, unzoned (Figure 5a), and nearly homogeneous in composition, as major element zoning is absent in all observed grains. Garnets with 27-55 mol% of pyrope, 25-48 mol% of almandine and spessartine, and 18-27 mol% of grossular and andradite content belong to the almandine-py-rope-grossular series with high pyrope content, as
Figure 5. Primary and secondary mineral phases in eclogites.
(a) Photomicrograph - backscattered electron image (BSE) of unzoned homogenous garnet grains lacking any inclusions. Symplec-
titic rims of amphibole + plagioclase are replacing garnet rims in contact with omphacite.
(b, c) Typical inclusions in garnet belong to phengite, kyanite, omphacite, quartz and rutile (BSE).
(d)	Homogenous omphacites are replaced by symplectites of diopside + plagioclase + amphibole.
(e)	Quartz rods and needles in matrix omphacite shown in plane-polarized light. Quartz rods are distinctly oriented and parallel to omphacite c-axis.
(f)	Omphacite with tiny quartz exolutions is hosting quartz inclusion surrounded by radial fractures (BSE). Abbreviations after Kretz (1983).
it is common for UHP metamorphic rocks (Chopin, 1984). Inclusions in garnet are relatively rare; these are typically omphacite, rutile, kyanite, quartz and phengite (Figure 5b, c). Garnet rims are commonly resorbed by amphibole + plagioclase symplectite (Figure 5a).
Omphacite occurs in large anhedral grains in the matrix (Figure 5d) or as inclusions in garnet (Figure 5c) and kyanite. Similarly to garnets, omphacites also show uniform grains and almost homogenous composition for all major oxides. The jadeite component of omphacites, calculated from Katayama et al. (2000):
Jd = Na - Fe3+ - 2-Ti
varies between 18-37 mol%. The total cation deficiency (omphacite cation totals are less than 4.00 per six oxygen atoms) and excess Al on the octahedral site suggest the existence of non-stoichio-metric pyroxenes or pyroxenes with octahedral vacancies. Ca and Al combine and form the Ca-Eskola molecule (CaEs - Ca0.5n0.5AlSi2O6; empty rectangle □ represents a vacancy on the M2 site). The cation sum of natural clinopyroxenes, calculated on the basis of six oxygens, decreases progressively from the theoretical value of 4.00 as the
VIAl content increases (Cawthorne & Collerson, 1974). The Ca-Eskola clinopyroxene, which is stable at peak metamorphic conditions, is highly unstable at lower pressure (Smyth, 1980). It rapidly breaks down to Ca-Tschermak component (CaTs - CaAl2SiO6) and quartz, following the retrograde reaction:
CaAl2SiOe + 3 • Si02
2-Ca0:5n05AlSi2O6 which can be simplified to: 2 • CaEs ^ CaTs + 3 • Qtz
The Ca-Tschermak component is assigned to be equivalent to the IVAl content (Katayama et al., 2000), and may be calculated from:
CaTs = 2 -Si=IVAl
The decompressional breakdown of Ca-Eskola molecule results in the exsolution of tiny needles and rods of quartz that are the most striking feature of matrix omphacites from Pohorje eclogites (Figure 5e, f). They display an orientation parallel to the c-axis of omphacite. Microprobe analysis of quartz needles show essentially pure SiO2. The amount of the Ca-Eskola component in omphaci-
tes from Pohorje samples, calculated from Katayama et al. (2000):
CaEs = Al -2 -,vAl-K-(Na- Fei+ -2-Ti)
is reaching 5 mol%. Integral analysis of omphacite, together with SiO2 precipitates, under defocused electron beam (25-30 |im) yields even higher values, up to 9 mol%. Omphacite is very sensitive to retrogression. During decompression, omphaci-tic clinopyroxene must reduce its jadeite content, which results in the production of plagioclase. The typical product of omphacite retrogression is the formation of symplectites of fine-grained diopside + plagioclase + amphibole (Figure 5d).
Phengite occurs sporadically in the matrix (Figure 6 a) and as minor inclusions in garnet (Figure 5b) and omphacite. Measured phengite grains were nearly homogenous in composition since the content of major cations is close to constant. Phengite belongs to white micas which result from solid solution between muscovite and celadonite and is formed by an inverse Tschermak substitution starting from muscovite. The amount of cela-donite component present is evident from the Si content of phengite, which rapidly increases with
Figure 6. Primary and secondary mineral phases in eclogites (photomicrographs - backscattered electron images).
(a)	Preserved remnant of phengite with clearly visible cleavage is surrounded by symplectitic intergrowth of biotite and plagio-clase.
(b)	Dark-gray kyanite grains in comparison with garnet and omphacite. All minerals shown are surrounded by symplectites of secondary mineral phases. The most typical are amphibole + plagioclase after garnet, diopside + plagioclase + amphibole after omphacite and sapphirine + corundum + spinel + anorthite after kyanite.
(c)	Kyanite with omphacite inclusion is surrounded by coronas of secondary minerals. Two distinct symplectitic belts are present. Inner corona, next to kyanite, consists of sapphirine + corundum + spinel + anorthite. Outer corona, next to omphacite, is made of plagioclase + amphibole ± diopside.
(d)	Detail of kyanite corona in which separate symplectitic constituents are clearly visible. Abbreviations after Kretz (1983).
Figure 7. Quartz (former coesite?) inclusions in eclogites.
(a)	Radial fractures surrounding quartz inclusion in omphacite under plane-polarized light.
(b)	Polycrystalline polygonal quartz inclusion (PPQ) surrounded by tiny radial fractures within garnet host under crossed polars. Abbreviations after Kretz (1983).
increasing pressure (Hermann, 2002). The strong variation of phengite composition as a function of pressure makes phengite a crucial mineral component for determination of UHP metamorphism (Massone & Szpurzka, 1997). Investigated phengite grains from Pohorje eclogites contain up to 3.5 Si per formula unit (pfu). The biotite + pla-gioclase intergrowths are typical replacements of phengite (Figure 6a).
Kyanite forms frequently twinned subhedral grains (Figure 6b) and small rod-like inclusions within garnet (Figure 5c) and omphacite minerals. Rare inclusions found in kyanite belong to garnet, omphacite (Figure 6c) and quartz. Microprobe analysis of all measured kyanite grains revealed almost pure Al2SiO5. Retrogression of kyanite is expressed by development of complex coronas consisting of sapphirine + corundum + spinel + anorthite (Figure 6b, c, d). The width of the corona is progressively increasing with the increasing degree of retrogression. Spinel belongs to Fe-Mg
Table 2. Calculated temperatures and pressures.
spinels. Sapphirines contain up to 1.4 Si pfu and are clearly peraluminous. Growth of corundum and tiny lamellar grains of sapphirine is restricted only to domains formerly occupied by kyanite.
Zoisite mainly forms individual elongated grains but may also be found as minor inclusions in garnet and omphacite. It contains rare inclusions of rutile.
Quartz inclusions are present in garnet (Figure 5c), omphacite (Figure 5f, 7a) and kyanite. They are frequently surrounded by radial fractures (Figure 7a) which may imply the possibility for the existence of former coesite. Some of the quartz inclusions surrounded by radial fractures, are aggregates of several polycrystalline quartz grains (Figure 7b). They strongly resemble the PPQ (polycrystalline polygonal quartz) and MPQ (multicrystalline polygonal quartz) textures described by Wain et al. (2000). Those quartz inclusions are interpreted as possible pseudomorphs after coesite.
Temperature (°C)
Sample	EG-79	P-85	K-88	PN-89	A-94	KR-00	Minimum	Maximum	Average
JV103a	843	822	802	573	741	748	741	843	791
NO1/04A	921	903	888	777	817	820	817	921	870
PO6h	824	803	781	588	710	717	710	824	767
SP1/08	801	779	755	516	694	702	694	801	746
	Temperatures calculated at 3 GPa.								
Geothermometers: EG-79: Ellis & Green (1979), P-85: Powel (1985), K-88: Krogh (1988), PN-89: Pattison &Newton (1989), A-94: Al (1994), KR-00: Krogh Ravna (2000). PN-89 calibration is excluded from minimum and average temperature calculations.
Pressure (GPa)
Sample	KRT-04[phe- qtz]	KRT-04[di> grs> qtz]	KRT-04[prp- qtz]	KRT-04[SiO2- ky]	WM-96	Minimum	Maximum	Average
JV103a	2.9	3.2	3.1	3.1	3.0	2.9	3.2	3.1
NO1/04A	2.9	3.2	3.1	3.1	2.9	2.9	3.2	3.1
PO6h	2.9	3.2	3.1	3.1	2.8	2.8	3.2	3.0
SP1/08	2.9	3.2	3.1	3.1	3.0	2.9	3.2	3.1
_Pressures calculated at 800 °C._
Geobarometers: WM-96: Waters & Martin (1996), KRT-04: Krogh Ravna & Terry (2001, 2004).
Geothermobarometry
Calculated temperatures from different calibrations of the garnet-clinopyroxene system (Ellis & Green, 1979; Powell, 1985; Krogh, 1988; Patti-son & Newton, 1989; Ai, 1994; Krogh Ravna, 2000)
differ substantially (Table 2, Figure 8). The unrea-listically low temperatures were obtained by Pat-tison & Newton's (1989) geothermometer, while the highest temperatures were calculated from the calibration of Ellis & Green (1979). Powel (1985) and Krogh (1988) calibrations give reasonable
Figure 8. Comparison of geothermobarometric results calculated with different geothermometers and geobarometers. Geother-mometers: EG-79: Ellis & Green (1979), P-85: Powel (1985), K-88: Krogh (1988), A-94: Ai (1994), KR-00: Krogh Ravna (2000). Geobarometers: WM-96: Waters & Martin (1996), KRT-04: Krogh Ravna & Terry (2001, 2004). Abbreviations after Kretz (1983).
results but obtained temperatures are still very high. The most reliable are temperatures obtained by geothermometers of Ai (1994) and Krogh Ravna (2000). With the exception of Pattison & Newton's (1989) calibration which obviously underestimates peak temperature conditions in mafic and ultramafic lithologies, the temperature intervals calculated at 3.0 GPa pressure range from 741 to 843 °C for JV103a sample, from 817 to 921 °C for N01/04A sample, from 710 to 824 °C for PO6h sample, and from 694 to 801 °C for SP1/08 sample. Average peak temperatures obtained are 791 °C (JV103a sample), 870 °C (N01/04A sample), 767 °C (PO6h sample), and 746 °C (SP1/08 sample).
Peak pressure estimations of Waters & Martin (1996) and Krogh Ravna & Terry (2001, 2004) calibrations yielded consistent results (Table 2, Figure 8). The intersections between these two geo-barometers and the geothermometers of Ai (1994) and Krogh Ravna (2000) define average peak pressures of 3.0 GPa for PO6h sample and 3.1 GPa for JV103a, N01/04A and SP1/08 samples, calculated at 800 °C. Excellent fitting is obtained mostly between Waters & Martin (1996) geothermobaro-meter, the pyrope absent reaction with coesite and the SiO2-kyanite absent reaction from Krogh Ravna & Terry (2001, 2004) calibration system.
The combination of garnet-clinopyroxene Fe2+-Mg exchange geothermometer (Krogh Ravna, 2000) with the geobarometric calibrations based on the net-transfer reactions in the garnet-cli-nopyroxene-phengite-kyanite-quartz/coesite system (Krogh Ravna & Terry, 2004), resulted in similar but more precise estimations of peak me-tamorphic conditions. The intersections between the used geothermometer and geobarometers define optimized maximum pressure of 3.0 GPa for samples JV103a, PO6h and SP1/08 at temperature range from 750 to 782 °C; and pressure of 3.1 GPa at temperature 783 °C for N01/04A sample (Figure 9).
All estimated peak pressure and temperature values consistently plot above the quartz-coesite transformation curve and thus correspond well to the ultrahigh-pressure stability field of coesite (Figure 8, 9). The quartz-coesite and graphite-diamond transformation boundaries were calculated from thermodynamic data of Holland & Powell (1998).
Discussion
UHP metamorphism in eclogites from Pohorje is evident both from microtextural observations and from the results of geothermobarometric calculations of peak metamorphic conditions, which revealed very high pressures and temperatures of 3.0-3.1 GPa and 750-783 °C. These values correspond well to the coesite, i.e. ultrahigh-pressure, stability field. Pressures calculated in this work are much higher than the former estimates of 1.8 GPa by Hinterlechner-Ravnik et al. (1991) or 1.5 GPa by Koch (1999).
Remnants of coesite, a direct mineral indicator of UHP conditions, were not found but its exi-
Temperature (°C)
Figure 9. Results from geothermobarometry. Combination of garnet + omphacite + phengite + kyanite + quartz/coesite assemblage (Krogh Ravna & Terry, 2004) with garnet-clinopy-roxene Fe2+-Mg exchange thermometer (Krogh Ravna, 2000). The quartz-coesite and graphite-diamond curves are calculated from thermodynamic data of Holland & Powell (1998). Abbreviations after Kretz (1983).
stence is clearly revealed from: (1) radial fractures around quartz inclusions within robust host minerals (garnet, kyanite and omphacite) that were caused by expansion ensuing the transformation of high-pressure coesite to its low-pressure polymorph; and (2) polycrystalline appearance of these inclusions interpreted as a pseudomorphs after former coesite due to their distinctive PPQ and MPQ microtextures (Wain et al., 2000).
Radial fractures around quartz inclusions may imply the possibility for the existence of former coesite and are therefore an indicator of possible UHP conditions. The roughly 10% of volume increase in coesite-quartz transformation process produces considerable overpressure (more than three times the lithostatic pressure), which buffers the inclusion at the coesite-quartz transformation boundary (Gillet et al., 1984).
The multistage transformation of coesite to strain-free quartz was described from experimental studies (Mosenfelder & Bohlen, 1997) and also observed in UHP rocks from many well-established UHP localities (e.g. Wain et al., 2000). Transformation begins with the growth of quartz on the coesite-host contact, which is followed by the development of shear cracks with no open volume (Hirth & Tullis, 1994), providing new sites for quartz growth. With progressing transformation, quartz rim around coesite inclusion succes-
sively changes from textureless to fibrous radial texture and to more irregular, feather-like texture. In the final stage of transformation, quartz inclusions exhibit typical radial palisade texture; if any relic coesite is present, palisades of quartz radiate from it. When absolutely no coesite is preserved, its former existence may be deduced from the three different sequential textural types of quartz inclusions: PRQ, PPQ and MPQ (Wain et al., 2000). Polycrystalline radial-texture quartz (PRQ) inclusions develop prior to radial fracturing in the host mineral and result from stress release. They consist of individual strain-free palisades, which often radiate from a core zone. The core zone is made of quartz which has irregular mosaic texture, sutured boundaries and undulatory extinction. Polycry-stalline quartz inclusions with mosaic texture and sutured contacts containing 30-100 grains also belong to the PRQ group. With proceeding recovery, the number of grains per inclusion is continuously reduced. Polycrystalline polygonal quartz (PPQ) inclusions consist of 10-50 polygonal strain-free quartz grains. Their formation post-dates radial fracturing. The next step of recovery is represented by multicrystalline polygonal quartz (MPQ) inclusions with 2-10 monocrystalline quartz grains per inclusion. Eventually, the recovery is completed when the whole inclusion recrystallizes to a single unstrained monocrystalline quartz surrounded by radial fractures.
Additional evidence for UHP metamorphism is coming from the presence of highly pressure-sensitive non-stoichiometric supersilicic clinopy-roxenes that directly indicate high pressure and temperature conditions (Gasparik & Lindsley, 1980; Fockenberg & Schreyer, 1997). They were reported as a typical constituent of eclogites from various UHP terranes (summarized by Bruno et al., 2002). Such clinopyroxenes containing Ca-Escola molecule in solid solution favourably form at high pressures exceeding 3 GPa (SchmAdicke & Muller, 2000). Due to the pressure decrease quartz needles and rods form as an exsolution product from su-persilicic clinopyroxene. They are oriented along the c-axis of omphacite and are resulting from the omphacitic clinopyroxene that contained excess silica at peak metamorphic conditions (Zhang & Liou, 2000). The occurrence of oriented quartz needles and rods in matrix omphacites is typical feature of many UHP eclogites (e.g. Liou et al., 1998). The eclogites from Pohorje show both, direct and indirect, evidence for pyroxene containing the Ca-Eskola component: the presence of clinopyroxene with Ca-Eskola molecule in solid solution, and quartz rods and needles as an exolu-tion product.
The survival of coesite relics depends on many factors, like the presence of sufficiently rigid host mineral, the extent of metamorphic re-equilibration of a mineral assemblage during exhumation, the trend of exhumation P-T path, the rate of exhumation, the presence of fluid phase on the reactant side of retrograde reactions (e.g. Mosenfelder & Bohlen, 1997; Hermann, 2002). Continental crustal rocks entering the UHP realm are not completely
dry, as one would expect. The majority of UHP assemblages contains hydrous minerals, like phen-gitic mica and epidote, that are stable under UHP conditions and may evolve small amounts of fluids by dehydration reactions (Chopin, 2003). Therefore, UHP parageneses may only be preserved if the entire exhumation path is situated within the stability field of phengite. This can happen either when continuous cooling is accompanying exhumation or when the entire exhumation process occurred at T<700°C (Hermann, 2002). In contrast, in terranes that show isothermal decompression, breakdown of phengite and concomitant partial melting may cause a complete recrystallization of the subducted crustal rocks during exhumation, and thus destruction of all UHP mineral assemblages.
Granulite facies metamorphism is usually required for sapphirine formation (Ackermand et al., 1975). Composition of sapphirines from Pohorje eclogites is clearly peraluminous and is significantly different from the composition of sap-phirines that are stable within granulite facies conditions (Higgins et al., 1979). Furthermore, the formation of sapphirine is very dependent on H2O activity, and for this reason is by itself not an indication that typical granulite facies conditions were reached (Simon & Chopin, 2001). Therefore it is proposed that sapphirine in Pohorje eclogites formed during decompression at lower-temperature conditions than those common for granulite-facies. Presence of sapphirine in coronas around kyanite in eclogites from Pohorje clearly indicates that during the early stages of decompression high temperature conditions persisted. Due to this near-isothermal decompression, fracturing and introduction of fluids was possible already at high temperatures. Because the Pohorje eclogites crossed the transformation border between coe-site and quartz at such high temperatures, there was no possibility for the survival of coesite.
The only known UHP metamorphism of Alpine eclogites so far was documented in the Western Alps, e.g. Dora Maira Massif (Chopin, 1984), Zer-matt-Saas Zone (Reinecke, 1991), and is of Tertiary age (Tilton et al., 1991; Rubatto et al., 1998). The determined P-T conditions of Eo-Alpine me-tamorphism in Pohorje are the highest reported within the Austroalpine units of the Eastern Alps. Thus, the Pohorje eclogites are the first evidence of UHP metamorphism in Eastern Alps, and also the first evidence that UHP metamorphic conditions were also reached during the Cretaceous orogeny.
Conclusions
Kyanite eclogites from Pohorje experienced UHP metamorphic conditions of 3.0-3.1 GPa at temperature range from 750 to 783 °C.
Radial fractures around polycrystalline quartz inclusions with microtextures diagnostic for recovery after coesite, together with oriented quartz exolutions in clinopyroxene, and omphacites with Ca-Eskola molecule are the microtextural evidence for UHP metamorphism.
During the first stages of exhumation, eclogites were exposed to near-isothermal decompression that lead to complete breakdown of coesite.
Pohorje eclogites record the highest pressure conditions of Eo-Alpine metamorphism in the Alps and thus provide the first evidence that UHP metamorphism in the Alps occurred already during the Cretaceous orogeny.
UHP eclogites from Pohorje represent the first occurrence of UHP metamorphic rocks within the Eastern Alps.
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