Metamorphic Evolution of the Middle Seve Nappe in the Snasahögarna Area, Scandinavian Caledonides

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Examensarbete vid Institutionen för geovetenskaper Metamorphic Evolution of the Middle Seve Nappe in the Snasahögarna area, Scandinavian ISSN 1650-6553 Nr 277 Caledonides

Åke Rosén Metamorphic Evolution of the

In the Snasahögarna area, situated in the central Scandinavian Middle Seve Nappe in the Caledonides, the Middle Seve Nappe Complex mainly consist of pelitic Snasahögarna area, Scandinavian gneisses. Within these garnet and kyanite bearing gneisses the first Caledonides discovery of metamorphogenic diamond in Sweden has been confirmed by micro raman spectroscopy. The garnet hosted diamond together with associated carbonate and fluid inclusions were deposited by carbon saturated fluids at a pressures exceeding 3.1 GPa. The peak mineral assemblage, Phg Mgs-Sid Grt Ky Jd Rt Coe (Dia), is constrained by Åke Rosén microscopic observations and thermodynamic phase equilibrium modeling. Modeled garnet cores equilibrated at c. 750oC and 1.3 GPa together with phengite rims. This early exhumation stage was followed by heating reaching a peak temperature of 880-910oC at 1.05-1.18 GPa. Partial melting following the increased temperature is confirmed by anatectic segregations and microscopic melt related textures in the paragneiss. A fluid restricted environment at the late high temperature stage is inferred from mineral textures, mineral chemistry and low bulk rock loss on ignition. This is in agreement with thermodynamic models.

Uppsala universitet, Institutionen för geovetenskaper Examensarbete E1, 30 hp i Berggrundsgeologi ISSN 1650-6553 Nr 277 Tryckt hos Institutionen för geovetenskaper, Geotryckeriet, Uppsala universitet, Uppsala, 2014. Examensarbete vid Institutionen för geovetenskaper ISSN 1650-6553 Nr 277

Metamorphic Evolution of the Middle Seve Nappe in the Snasahögarna area, Scandinavian Caledonides

Åke Rosén

Supervisor: Jaroslaw Majka

Copyright © ¯LF3PTÏO and the Department of Earth Sciences Uppsala University Published at Department of Earth Sciences, Geotryckeriet Uppsala University, Uppsala, 2014 Abstract In the Snasahögarna area, situated in the central Scandinavian Caledonides, the Middle Seve Nappe Complex mainly consist of pelitic gneisses. Within these garnet and kyanite bearing gneisses the first discovery of metamorphogenic diamond in Sweden has been confirmed by micro raman spectroscopy. The garnet hosted diamond together with associated carbonate and fluid inclusions were deposited by carbon saturated fluids at a pressures exceeding 3.1 GPa. The peak mineral assemblage, Phg Mgs-Sid Grt Ky Jd Rt Coe (Dia), is constrained by microscopic observations and thermodynamic phase equilibrium modeling. Modeled garnet cores equilibrated at c. 750oC and 1.3 GPa together with phengite rims. This early exhumation stage was followed by heating reaching a peak temperature of 880-910oC at 1.05-1.18 GPa. Partial melting following the increased temperature is confirmed by anatectic segregations and microscopic melt related textures in the paragneiss. A fluid restricted environment at the late high temperature stage is inferred from mineral textures, mineral chemistry and low bulk rock loss on ignition. This is in agreement with thermodynamic models.

Sammanfattning Den mellersta Seveskållan i området kring Snasahögarna i Jämtland domineras av granat- och kyanitförande paragnejser. I dessa har metamorfogena diamanter identifierats med hjälp av ramanspektroskopi. Diamanterna återfinns som inneslutningar i granat tillsammans med associerade karbonater samt vätskeinneslutningar. Dessa avsattes från kolmättade vätskor vid tryck överstigande 3.1 GPa. Mineralsammansättningen vid maximalt tryck, Phg Mgs-Sid Grt Ky Jd Rt Coe (Dia), har härletts från mikroskopobservationer samt termodynamisk fasmodellering. Jämvikt för anomala granatsammansättningar inneslutna i senare homogeniserad granat uppnåddes nära 750oC och 1.3 GPa. Efterföljande tryckavlastning medförde värmeökning till 880-910oC vid 1.05-1.18 GPa. Anatektisk textur samt smältrelaterade mikrostrukturer tyder på partiell uppsmältning till följd av ökad temperatur och minskat tryck. Mineraltexturer, mineralkemi och en låg viktförlust vid upphettning av prov indikerar låg vätsketillgång vid den sena högtemperatursfasen. Detta bekräftas av termodynamisk fasmodellering. Contents 1. Introduction ...... 1

2. Background ...... 1

2.1. Pre-Caledonian Baltica ...... 1

2.2. Caledonian orogeny ...... 3

2.3. Scandinavian Caledonides ...... 4

2.4. The Snasahögarna area and surroundings ...... 9

3. Previous studies ...... 11

3.1. Geothermobarometry ...... 11

3.2. Geochronology ...... 12

4. Methods ...... 13

4.1. Thin sections ...... 13

4.2. Microprobe analysis ...... 13

4.3. Raman spectroscopy ...... 14

4.4. Whole rock chemistry ...... 14

4.5. Geothermometry ...... 14

4.6. PT Pseudosections & Isopleths (thermodynamic modeling) ...... 14

5. Results ...... 15

5.1. Petrography ...... 15

5.2. Raman spectroscopy ...... 35

6. Geothermometry ...... 42

6.1. Garnet-Biotite thermometer ...... 42

7. Thermodynamic phase equilibrium modeling ...... 44

7.1. Model 1 ...... 44

7.2. Model 2 ...... 45

7.3. Exhumation PT path ...... 45

8. Discussion ...... 50

8.1. Ultrahigh pressure metamorphism (UHP) ...... 50

8.2. Metamorphogenic diamond ...... 51

8.3. Metamorphic evolution ...... 52

8.4. Exhumation processes ...... 54

8.5. Timing of subduction ...... 54

9. Conclusions ...... 55

Acknowledgements ...... 56

References ...... 57

1. Introduction The Scandinavian Caledonides is a continent-continent collisional style orogen in many aspects well suited for studying the effects of orogenic processes. It was recognized several decades before the notion of plate tectonics that the nappes of the Scandes must have been vastly laterally displaced from where they had originally been deposited (Törnebohm 1888). It is now generally accepted that allochthonous units of the Scandinavian Caledonides were originally derived from as far away as the Laurentian margin (Gee et al. 2008). Given the age of the Scandes, through uplifting and erosional processes it is possible to study rocks and structures whose analogues are hidden deep below the ground surface in younger orogens. The evolution of the mountain chain and the timing of its different stages along with the proposed mechanisms of nappe emplacement and exhumation of subducted units is still not fully understood. This study focuses on the metamorphic evolution in a small part of the Seve Nappe Complex which is of significance to the interpretation of the tectonic history of the Scandinavian Caledonides. In the light of recent results on Åreskutan migmatites which indicates underestimated metamorphic grades in the area, this work aspires to reevaluate the progression of metamorphic conditions in crustally derived gneisses from Snasahögarna. This is done using state of the art techniques such as microprobe analysis, raman spectroscopy and thermodynamic modeling together with conventional methods such as textural studies and geothermometry. The hypothesis is, that the studied rocks have undergone higher metamorphic conditions than previously described.

2. Background The setting before the Caledonian orogeny followed by a description of the orogenic events and the present Scandinavian Caledonides are outlined below.

2.1. Pre-Caledonian Baltica The geological history of the Baltic Shield is long. Early continental core remnants formed in the Mesoarchean prior to the assemblage of the Fenno-Karelian protocontinent (Slabunov et al. 2006). Rifting followed in the Paleoproterozoic with the emplacement of layered gabbroic plutons in the northeastern Fennoscandian shield at c. 2.5 Ga (Lahtinen et al. 2008). During and following the Svecofennian orogenic events (1.95-1.85 Ga), plutonic and volcanic rocks were added through accretion. These units build up the dominant part of the present Baltoscandian platform which borders Archean crystalline basement in the northeast. Westwards crustal growth by accretion and continental magmatism lasted to the end of the

1 disputed Gothian orogeny (1.55 Ga). During this time rapakivi granites were emplaced in the interior of Fennoscandia. Later in the Mesoproterozoic the tectonic setting changed drastically with orogenic growth giving way to a generally extensional regime with deposition of sediments and magmatic intrusions until 1.2 Ga (Bogdanova et al. 2008).

The following Sveconorwegian orogeny (1.14-0.9 Ga) affected the present southwestern parts of Norway and bordering parts of Sweden. This time period was defined by high grade metamorphism and reworking of Svecofennian crust. A model involving four distinct orogenic phases was suggested by Bingen et al. (2008). At 1.14-1.08 Ga the first phase, named Arendal after the town in southeastern Norway, led to the formation of two orogenic wedges. These consist of various plutonic rocks and associated metasediments found in and north of the present Arendal area. The subsequent main collisional Agder phase (1.05-0.98 Ga) is characterized by nappe deposition in the present Swedish west coast area, exhumation of high pressure units and plutonic intrusions in southern Norway. During the Falkenberg phase (0.98-0.96) eclogites were emplaced and partly exhumed. Remnants of these are now found in the swedish region surrounding the town of Falkenberg verifying deep subduction during this orogenic stage. At the end of the Sveconorwegian event, extension during the Dalane phase (0.96-0.90 Ga) is presumed to have caused successive exhumation of metamorphic units (Bingen et al. 2008). Baltica is thought to have been incorporated into the supercontinent (Rodinia) sometime during the late Proterozoic eras. This still controversial idea places the Laurentia paleocontinent next to Baltica and Amazonia during a substantial part of the Neoproterozoic (e.g. Meert & Torsvik 2003).

Early rifting between the adjacent Baltica and Laurentia is considered to have started in middle Neoproterozoic time. The final stage of the continental break up is marked by vast intrusion of dolerite dykes at c. 600 Ma within the rifted units (Nystuen et al. 2008). Cambrian siliclastic sediments were deposited on the margin of Baltica which was separated from Laurentia by the newly formed Iapetus Ocean. Alum shales accumulated over conglomerates and sandstones on Baltica in the Middle Cambrian, giving way to carbonate deposition first in the Ordovician. On the Laurentian margin, deposition of carbonates started already in the Cambrian (Gee et al. 2008).

2.2. Caledonian orogeny The Caledonian orogeny constitutes the last accretionary event to add lithosphere to the Baltoscandian shield. Convergence of Baltica and Laurentia is considered to have started in the early Ordovician with subduction zones forming along both the Laurentian and the Baltican margin. The main collisional stage was focused sometime between the mid Silurian and the onset of the Devonian. During this time, the margin of Baltica was underthrusting Laurentia. This lead to partial melting of sediments which are now found amongst translated nappes that were stacked, along the Scandinavian Caledonides, over the Baltoscandian platform. A section of the suture zone can be inferred stretching from east of Svalbard and recognized down through the Scandinavian Caledonides to the south of Scandinavia (Gee et al. 2008).

The underthrusting and collision left thrust related structures that are dominantly east-vergent in Scandinavia and west-vergent on Greenland. After the converging stage gravity driven extensional collapse followed in the Devonian (Gee et al. 2008). This left extensional structures that can be observed in Caledonian outcrops in the mountain region of western Scandinavia today (e.g. Greiling et al. 1998).

In the north Atlantic, the Caledonian orogen can be followed from the far north in the arctic archipelago northeast of Svalbard, known as Franz Josef Land. Southwards, it is recognized through Svalbard down to northeastern Greenland and westernmost Scandinavia. In the far northeast of Norway it terminates at the northwesternmost part of the Timanides. Further south the Caledonides branches across the northern British Islands and can be traced in central Europe towards the border between Germany and Poland.

2.3. Scandinavian Caledonides In Scandinavia, the Caledonian orogen largely covers Norway and the most western parts of Sweden, from the far north, all along the border between Norway and Sweden. The stratigraphy of the Scandinavian Caledonides is by convention divided, on a tectonic basis, into the Autochthon and Parautochton, the Lower Allochthon, the Middle Allochthon, the Upper Allochthon, and the Uppermost Allochthon (Gee & Zachrisson 1979), (Fig 1). Generally, the lower translated tectonic units consist of Baltica basement sedimentary cover and detached basement units. Moving up the tectonostratigraphy, higher units consist of Baltica margin sediments and further Iapetus Ocean sediments with related intrusive units to finally, in the uppermost units, rocks of Laurentian affinity (Gee et al. 2008).

Fig 1. Tectonic map of the Scandinavian Caledonides. Modified from Gee et al. 2010

Autochthon and Parautochthon The contact between the gently westward dipping Authochton (Precambrian basement with overlying sediments), the Parautochton (sediments resting on mobilized basement) and overriding units can be traced along the Caledonian front, stretching from the western far north of Sweden to the southwest of Norway and in windows across the central parts of the orogen. The Autochton generally consists of eroded granites and volcanics with a sedimentary cover, usually not reaching more than some tenths of meters in thickness, dominated by various sandstones overlain mainly by Cambrian alum shales, but with local variations in lithology (Gee et al. 1974; Nystuen et al. 2008). In Jämtland, antiformal windows reveal a sedimentary succession of conglomerates, quartzites, alum shales and locally limestones that were deposited on the commonly horizontally and/or vertically displaced Svecofennian basement under the now overriding Caledonian front. Westwards into Norway windows reveal successively increasing deformation affecting the basement rocks (Gee 1980).

Lower Allochthon The Lower Allochthon consists of Baltica margin derived sediments with minor incorporated basement. These Jämtlandian and related nappes typically crop out in and around the Jämtland area of central Sweden, in the far north of Sweden and Norway and in the south through to the southwest coast of Norway (Gee et al. 2010). Lithologies include Neoproterozoic to older Cambrian sandstones, turbidites and shales with overlying later (Mid to Late Cambrian) shales, limestones and greywackes. Neoproterozoic sandstones and tillites are represented further up in the nappe stack. The sediments and incorporated basement of the Lower Allochton is generally metamorphosed at greenschist facies conditions, but reaches amphibolite facies in the westernmost outcrop region in Norway causing difficulties in distinguishing them from the Parautochthon (Roberts & Gee 1985). Alum shales below the base of the Lower Allochton provided a competence contrast aiding the formation of the Caledonian frontal detachment zone (Gee 1980).

Middle Allochthon The Middle Allochthon is dominated by various metasediments and Precambrian crystalline rocks (Roberts & Gee 1985) displaying inverted metamorphism with estimated temperature and pressure increasing from the base upwards in the tectonic stratigraphy (Andréasson & Gorbatschev 1980). In Jämtland, thrusted basement in the Vemån Nappe and the Tännäs augen gneiss Nappe makes up the sole beneath the Offerdal Nappe (Gee et al. 2010). The Offerdal Nappe generally consists of late Precambrian metasediments including sandstones,

6 conglomerates and phyllites which although having undergone quite intense deformation still retain recognizable primary sedimentary structures (Plink-Björklund et al. 2005).

Overlying the Offerdal Nappe in Jämtland are the Särv nappes, together comprising a tectonic unit which is dominated by feldspathic quartzites mostly derived from Meso-Neoproterozoic Baltica. Generally, the Särv nappes are metamorphosed at greenschist facies. Their largest outcropping area is in Härjedalen and Jämtland, but they are also represented at several localities further west in Tröndelag, Norway (e.g. Be’eri-Shlevin et al. 2011). The Särv quartzites are densely intruded by dolorites related to the continental rifting at c. 600 Ma, which led to the formation of the Iapetus Ocean (Svenningsen 2001), before the sediments were displaced laterally towards the Caledonian hinterland during the orogeny. These frequent intrusions clearly separate the Särv nappes from underlaying units of similar lithology, a feature representing evidence of vast nappe translation.

The Seve Nappe Complex makes up the uppermost part of the Middle Allochthon and extends for c. 800 km along the Scandinavian Caledonides (Zachrisson 1973). In Norway, the related Kalak nappes (Andréasson et al. 1998) stretch westwards in the far north. The Seve Nappe was first individually named, together with the overlaying Köli Nappe, by Törnebom (1872) based on lithological distinctions. Later, it was placed within the Upper Allochton, then considered to make up the lower part of what was commonly named the Seve-Köli Nappe Complex (e.g. Zwart 1974).

In most areas the Seve Nappe can be subdivided in three subunits where the lower unit generally consists of metasandstones of amphibolite to locally eclogite facies. The middle unit, including Åreskutan and Snasahögarna, is dominated by pelitic gneisses and migmatites. The upper subunit mostly consist of amphibolites with minor metasediments. The metasediments of the Lower Seve Nappe show some lithological similarities with underlying units but what distinctly sets them apart is the metamorphic grade (Gee et al. 2010). Eclogites are found in two main outcropping areas, northern Jämtland (e.g. Zwart 1974) and further north in southern Norrbotten (e.g. Andréasson et al. 1985). Results from previous metamorphic studies on these are briefly outlined in section 3.1.

Spinel peridotites are present throughout the Seve Nappe Complex (Bucher-Nurminen 1991), whereas garnet-peridotites are only found in northern Jämtland. Both types of peridotites were suggested to be of alpine type derived from subcontinental lithosphere by Brueckner et al. (2004). They suggested a model where peridotites with varying metamorphic imprint were

7 incorporated into subuducting continental crust during the early (c. 450 Ma) collision between Baltica and a previously rifted microcontinent.

Upper Allochthon The Köli Nappe Complex (including the correlated Trondheim nappes) comprises the Upper Allochton overlaying the Seve nappes (Roberts & Gee 1985). Generally, these units consist of Cambrian-Silurian oceanic metasediments and island arc terranes (Greiling & Grimmer 2007) mainly metamorphosed at greenschist facies (Gee et al. 2010). The lower Köli Nappe, best exemplified in the Björkvattnet-Virisen area (Jämtland-Västerbotten), is predominantly of sedimentary origin, including limestones. Here a fossil record is preserved within various units in the upper parts of the stratigraphy. Middle Köli Nappe sedimentary rocks include schists and calcareous turbidites. Volcanics found in association with the turbidites are considered older than the sediments based on structural criteria. Sediments in Upper Köli Nappe include graphitic schists, turbidites and various carbonates. Volcanics here occur as debris, pillow lavas and interlayered units, at various structural levels, ranging in composition from basic to rhyolitic. (Stephens & Gee 1985).

Uppermost Allochthon The Uppermost Allochthon crops out solely in Norway. Concentrated to the Nordland and Troms regions the main lithologies consist of amphibolite facies volcanic and granitic gneisses, schists and carbonates with local conglomerates and metasedimentary iron ores (Stephens & Gee 1985; Roberts et al. 2007). Structural evidence (Roberts et al. 2007) combined with the stratigraphic record (Roberts et al. 2002) and the fossil record in the form of rare brachiopods and trilobites found in the underlaying Upper Allochthon in Scandinavia (Bruton & Harper 1985) points to the Uppermost Allochthon largely having been originally derived from the Laurentian continental margin.

2.4. The Snasahögarna area and surroundings The Snasahögarna mountains are situated in southwestern Jämtland close to the Norwegian border (Fig 2). To the Northeast is the town of Åre, resting beneath the southwestern slope of the classical Caledonian locality Åreskutan. Westwards, on the border and moving into Norway the Northwest trending, regional scale, Röragen detachment zone separates the Trondheim Nappe Complex of the Upper Allochthon from lower tectonic units to the East. The area nearest surrounding Snasahögarna is commonly referred to as the “Handöl area” after the small village of Handöl on the western shore of lake Ånnsjön.

The Seve and Köli nappes of the Handöl area have previously been a target of structural and metamorphic studies as the complexity of these translated units gives evidence of a multiphase (Sjöström 1983a) and highly interesting tectonic and metamorphic history.

The Köli Nappe in this area is represented by the Bunnerviken lense which encompasses the Handöl formation (Sjöström 1983b) in the Tännforsen synform (Beckholmen 1984). The Handöl formation generally consists of schists with minor calcareous schists. These schists sometimes display a garbenschiefer texture (Sjöström 1986).

In the Handöl-Snasahögarna region, the Seve Nappe Complex is represented by Lower, Middle and Upper subunits. The Middle Seve Nappe has been more strongly metamorphosed than the Lower and Upper units. Both the Upper and Lower units are dominated by amphibolites and mica schists whilst the Middle Seve here mainly consists of low granulite facies paragneisses (Bergman & Sjöström 1997).

The studied area occupies an N-S to NNW-SSE trending synform that is separated from underlying units by basal thrust zones. The Middle Seve Nappe, here locally referred to as the Snasahögarna Nappe, correlates with the stratigraphy of the Åreskutan Nappe. The mainly gneissic polymetamorphic sedimentary rocks are characterized by alkali feldspar-sillimanite assemblages (Sjöström 1983a). Secondary to these, felsic paragneisses, are minor lense shaped pyroxene gneisses. Based on these mineral parageneses Sjöström (1986) considered the main lithological units of Snasahögarna to be of upper amphibolite to granulite facies. This early metamorphic record would then have been overprinted by a later greenschist- amphibolite facies thrust related deformation, which is described as being of the least intense grade at the base of the Lower Seve subunit (Sjöström 1986, Bergman & Sjöström 1997).

Fig 2. Geological map of the Snasahögarna-Handöl area with red ellipse marking the study area (Snasahögarna mountains). Modified from SGU database 2012.

3. Previous studies This section briefly outlines some previous published and unpublished results from metamorphic studies and related geochronology within the Snasahögarna area and the Seve Nappe Complex.

3.1. Geothermobarometry A geothermobarometric study by Sjöström (1983c) on the Seve-Köli Nappe Complex, later subdivided into the Seve Nappe and Köli Nappe complexes respectively, was done using the distribution coefficient (KD) for the Fe and Mg distribution between garnet and biotite yielding temperature estimates and garnet-(Al2SiO5)-silica(quartz)-plagioclase (GASP) barometry using a variety of published calibrations. Estimated pressure-temperature (PT) conditions for samples collected in the Snasahögarna Nappe were in the range of 660-680oC and 600-700 MPa. Notably, three different samples yielded temperatures exceeding 800oC for two or more different calibration methods. These values were considered unreasonably high and suggested to be caused by a homogeneous increase of the Fe/Mg ratio in biotite due to retrogression which in turn would have increased the KD.

PT-paths for the Seve and Köli units in the Handöl area were constructed by Bergman (1992). This author used the Fe-Mg exchange between biotite and garnet to constrain peak temperatures, based on the Perchuk & Lavrenteva (1983) calibration. The garnet-biotite- muscovite-plagioclase (GSMB) assemblage was used to constrain pressures using the calibrations of Powell & Holland (1988) and Hoisch (1990). These methods gave temperatures of 630-770oC and pressures of 650-750 MPa for the Snasahögarna Nappe. Although complete PT-paths were presented for other local units in the Seve nappes the Snasahögarna Nappe was not included. Bergman (1992), however, interpreted calcium enriched garnet rims as evidence for pressure increase related to the superposition of the Köli Nappe on the local Seve nappes.

In more recent times thermodynamic modeling on stratigraphically correlated pelitic gneiss of Åreskutan gave evidence of a peak pressure phase of 2.6-3.2 GPa at 700-720oC followed by temperature increase to 800-820oC during exhumation (Klonowska et al. 2013).

In northern Jämtland, peak metamorphic conditions for eclogite, found at the top of mount Tjeliken near the village of Blomhöjden, was recently constrained to 2.5-2.6 GPa at 650- o 700 C (Majka et al. 2013). An essential question regarding this eclogite has been whether it is correlated with the surrounding Lower Seve Nappe or constitutes an overriding klippe as

11 suggested by Strömberg et al. (1984). The former explanation contradicts earlier pressure- temperature estimates for the Tjeliken eclogite (c. 1.4 GPa and 550oC) by Van Roermund (1985) and for associated gneisses (c. 1.65 GPa and 650-680oC) by Litjens (2002). A few tenths of kilometers north, garnet peridotites at lake Friningen have been estimated to have undergone 2-3 GPa and 700-800oC (Brueckner et al. 2004) and in the same area peak conditions for peridotite hosted eclogites was later constrained to c. 3 GPa and 800oC (Janák et al. 2012). Further north in Norrbotten pressure temperature conditions of 2.0-2.7 GPa at 650-750oC has been reported for eclogites within the Vaimok and Tsäkkok lenses (Albrecht 2000).

3.2. Geochronology Early ion microprobe U-Pb dating of zircon placed the high grade metamorphic events of Åreskutan at c. 441 Ma (Williams & Claesson 1987). This was later supported by conventional U-Pb monazite dating of the Åreskutan Nappe, 440-435 Ma and Lower Seve Nappe, 437-427 Ma (Gromet et al. 1996) in the same area. Recent U-Th-total Pb monazite dating of the Åreskutan migmatites places a progressive high pressure metamorphic event at c. 455 Ma (Majka et al. 2012) which is slightly older than zircon rim ages dated to 441-442 Ma by Ladenberger et al. (2010). In the Snasahögarna area, Be’eri-Shlevin (unpublished) dated calc-silicate rock from Tväråklumparna to 441 ± 4 Ma using U-Pb zircon geochronometry, adding further support to a late Ordovician high pressure episode in central Jämtland.

Dating results for the high pressure event recorded within the Seve Nappe Complex in northern Jämtland are inconclusive, but generally give older dates than those of central Jämtland. Brueckner & Van Roermund (2004; 2007) used garnet-pyroxene–amphibole-whole rock to construct isochrones in the Sm-Nd system in eclogite and garnet pyroxenites from the Middle and Lower Seve Nappe Complex in Jämtland. These authors obtained ages with a weighted average of 458 ± 4 Ma, concluded to all be within the error range, and proposed to name this high pressure phase the Jämtlandian Orogeny. The age is disputed by later metamorphic TIMS U-Pb zircon dating on the Tjeliken eclogite of 446 ± 1 Ma by Root & Corfu (2012).

In Norrbotten a high pressure stage within the Seve Nappe Complex clearly seems to be older than further south. Ar-Ar hornblende dating of the Grapesvare eclogite by Dallmeyer & Gee (1986) yielded an age of 491 ± 8 Ma. Similar (mineral and whole rock) Sm-Nd isochron ages

12 of 505 ± 18 Ma and 503 ± 14 Ma for the Grapesvare eclogite were obtained by Mørk et al. (1988). Mørk et al. (1988) suggested an eclogite forming event in present Norrbotten, predating the main Caledonian collision. Supporting such an event are conventional U-Pb dating results of titanite from the Grapesvare Nappe by Essex et al. (2007) which range from 500-475 Ma. The timing of this eclogite facies metamorphism is, however, not completely clear. Recent TIMS U-Pb zircon dating yielded 482 ± 1 Ma on the Tsäkkok and Vaimok eclogites (Root & Corfu 2012), falling in the lowest range of previous geochronological studies.

4. Methods The methods used in this metamorphic study on rock samples from the Snasahögarna area are outlined below.

4.1. Thin sections Thin sections were made and polished at the Geological Institute of the Slovak Academy of Sciences in Bratislava, Slovakia. The sections were all made strictly using two steps of diamond polishing agent which were of a consistent 1 µm and 3 µm size. Studies using light microscopy were carried out on the sections recording mineral assemblages and mineral textures.

4.2. Microprobe analysis Wavelength-dispersive X-ray spectroscopy (WDS) and energy-dispersive X-ray spectroscopy (EDS) analyses were performed on selected thin sections using a JXA-8530F JEOL HYPERPROBE, Field Emission Electron Probe Microanalyzer at the Department of Earth Sciences, Uppsala University. Operating conditions during quantitative analyses were 20kV acceleration voltage with a 20nA beam. Backscattered electron (BSE) and secondary electron (SE) imaging was carried out using the same facility during which acceleration voltage and beam current were manually adjusted for each image to obtain maximal resolution.

WDS was used for quantitative mineral chemistry analyses including garnet profiling. EDS was used to identify phases being too small or otherwise unsuitable for reliable optical identification. BSE and SE imaging was employed in the investigation of micron size mineral phases and mineral textures.

4.3. Raman spectroscopy A preliminary Raman study was carried out at the Department of Earth Sciences, Uppsala University. In total 14 spectra were obtained using a green argon laser with an applied power of 10mW and the acquisition time were 30 or 60 seconds depending on noise levels.

Micro Raman analyses mineral phase identification were carried out at AGH University of Science and Technology, Krakow, Poland, using a Thermo Scientific DXR Raman microscope. A red 633 nm, 12 mW maximum power laser and a green 532 nm, 10 mW maximum power laser were utilized. Settings for applied laser power varied between 1-12 mW and a 25 µm pinhole aperture was selected. Estimated laser spot radius on the sample surface was c. 0.7-1.3 µm. Final spectra are averages of 10-15 single spectra.

4.4. Whole rock chemistry A paragneiss rock piece was cut from sample JMY-19a/11, crushed and milled into a fine powder. 15 g of powdered rock was then sent to the ALS laboratory in Luleå, Sweden. Bulk chemistry analysis was carried out on major elements and additional minor constituents according to their ALS G-2 analyses package. 0.1 g of dried sample was melted with 0.4 grams of LiBO2 and dissolved in HNO3. The dissolved sample was then analyzed using ICP- SFMS. The loss on ignition (LOI) was determined at 1000oC with a reported uncertainty of 5 %.

4.5. Geothermometry Equilibration temperatures for garnet and biotite, in paragneiss samples, at 1 GPa based on Fe and Mg distribution coefficients were calculated using the calibrations of Thompson (1976), Holdaway & Lee (1977), Ferry & Spear (1978), Hodges & Spear (1982), Perchuk & Lavrent’eva (1983), Dasgupta et al. (1991) and Bhattacharya et al. (1992).

4.6. PT Pseudosections & Isopleths (thermodynamic modeling) Whole rock chemistry from sample JMY-19a-11 was used to calculate PT pseudosections and compositional isopleths in Perple_X_07 thermodynamical software (Conolly version 2005). Perple_X software allows employment of gridded Gibbs free energy minimization calculating thermodynamically stable mineral assemblages at specified conditions. Using Perple_X it is possible to construct pressure-temperature pseudosections and, with the included werami.exe application, isopleths of phase compositions and pseudo-compounds derived from selected solution models.

Two PT pseudosections for different metamorphic stages were constructed following the approach of Massonne (2011; 2013). This was done due to mineral assemblages and mineral chemistry indicating variations in bulk composition (H2O and CO2), (see sections 5.1 and 5.2). In model 1 the following solution models were used: Garnet (White et al. 2007), phengitic white mica (Coggon & Holland 2002), biotite (Tajčmanová et al. 2009), plagioclase (Newton et al. 1980) and carbonate (Holland & Powell 1998). In model 2 the same solution models where used but adding the melt model by Holland & Powell (2001), White et al. (2001) and discarding the carbonate solution model.

Calculations were performed using thermodynamical data of Holland & Powell (1998). Isopleths of melt and kyanite volume, Si (a.p.f.u) in phengite, XFe in Biotite, XAn in plagioclase, XGrs and XPrp content in garnet were calculated using werami.exe. The selection of calculated and presented isopleths is based on well-known pressure and temperature dependences of phase modal volumes and mineral compositions in metapelitic systems (e.g. Holland & Powell 1998).

5. Results The results from the previously outlined analytical methods are presented below including general petrography, mineral chemistry and phase identification followed by geothermometrical application and thermodynamical modeling.

5.1. Petrography In this section the studied rock samples are described based on observations from optical microscopy studies and results from microprobe electron imaging and mineral analyses. Samples collected during the summer of 2011, in the Snasahögarna area, were investigated. These samples are listed and generally defined in Table 1. Alongside are abbreviations that are used in the following sections of this thesis. The samples are all ascribed to a pelitic origin but have undergone varying degrees of deformation and partial melting displaying textures ranging from schistose-gneissic to quartz ribboned mylonitic. Samples (JMY-18/11, JMY- 19/11 and JMY-19a/11) display anatectic textures with segregated melanocratic and leucocratic sections although these are most clearly developed in sample JMY19a/11 (Fig 3).

Table 1. Sample labels and general rock type.

Sample Rock type Text abbreviation JMY-16/11 Quartzofeldspatic mylonite 16 JMY-18/11 Anatectic paragneiss 18 JMY-19/11 Anatectic paragneiss 19 JMY-19a/11 Anatectic paragneiss 19a JMY-19b/11 Calcic schist 19b JMY-21/11 Finegrained paragneiss 21 Unmarked Calcic schist X

Fig 3. Photograph showing melanocratic and leucocratic parts in a section cut across the layering in paragneiss (sample JMY-19a/11).

Paragneiss and mylonite The main mineral assemblage in the four studied paragneisses consists of quartz, feldspar, garnet biotite and usually sillimanite. Quartz, orthoclase and plagioclase, which commonly exhibits polysynthetic twinning, dominate the leucocratic matrix. One sample is set apart by the presence of a fine grained matrix. Garnet in this sample is small in size (commonly below 250 µm) compared to other samples and well preserved as it occasionally retains an almost euhedral shape (Fig 4).

Fig 4. Photomicrographs of matrix with small almost euhedral garnets in sample 21 (left) and quartz- feldspar dominated matrix in sample 19a, crossed polarizer (right).

Segregated melanocratic areas in three samples are generally defined by biotite and sillimanite, which often grow together in thick bundles adjacent to garnet. Kyanite occurs in these areas but in very restricted amounts. Larger, mostly anhedral kyanite crystals, found in association with sillimanite, and small grains enclosed in biotite could represent two different generations of growth. Garnet is subhedral to anhedral with one dominating set of parallel, healed fractures running straight through the grains. Abundant inclusions of quartz, biotite, sillimanite, feldspars and accessory minerals are common in the anhedral garnets (Fig 5). Such garnet often consists of several coalesced garnet crystals. White mica is rare. Tiny, highly altered grains of paragonite and phengite (muscovite-celadonite) occur in two samples. These micas are found in the matrix or enclosed in other minerals (commonly garnet). Biotite and occasionally paragonite is partly replacing phengite.

Fig 5. Photomicrographs of (A) Biotite growing in a bundle with sillimanite close to garnet in sample 19a, crossed polarizers (B) Fibrolite and biotite surrounding kyanite in sample 18 (C) Sillimanite and biotite next to kyanite and garnet in sample 19a, crossed polarizers (D) Inclusion rich garnet in sample 19a with spinel below.

The most abundant accessory phases in the paragneisses are rutile and ilmenite which occur in the matrix and as inclusions in garnet. Zircon, monazite, titanite and spinel are other common phases. Less common are carbonates and apatite. Minor amounts of calcite were found in the matrix of one sample. Flaky elongated grains of graphite are often found in association with areas of denser growing garnet and as inclusions in garnet in two samples (19a and 21). Cordierite is not common but was observed in one sample (18), where it usually displays simple twinning and some degree of pinitization.

One mylonite sample (16) is included in the study. Quartz ribbons in this rock define the mylonitic texture. These are coherent, 1-3 mm thick and border large perthitic microcline megacrysts. The megacrysts are mantled by finely recrystallized feldspar and occasionally

18 exhibit crosshatched twinning. Fine grained muscovite regularly occurs along the foliation, parallel to quartz ribbons in this sample.

Calcic schist Studied calcic schist samples (19b and X) show distinct differences to the paragneisses. Disseminated garnet, orthopyroxene and less common clinopyroxene, clearly represent earlier mineral assemblages than the quartz-feldspar dominated matrix in these rocks. Garnet is anhedral, showing no preferred orientation and inclusions of quartz, feldspar or minerals of older assemblages occur in almost every grain. Sillimanite in these rocks mostly occurs as fibres in the matrix. Minor kyanite has also been found. Biotite is not very common in the samples. Instead, most mica is represented by phengite. In one rock, a few grains of rare intensely dark-brown pleochroic barium mica were identified, its chemistry and optical features closest resembling that of oxykinoshitalite (Kogarko et al. 2005).

Signs of fluctuating levels of available H2O and gas components throughout different stages of mineral growth were observed. Phengite is commonly replaced by highly anorthitic plagioclase. Such plagioclase, and a few grains of hornblende found in sample 19b are chloritized around the rims. This appears to have happened at a very late stage, likely during shearing as the chlorite show a preferred orientation of growth. Manganese rich ilmenite was found replacing remnant pieces of rutile which are mantled by titanite, rich in aluminum.

These features can indicate fluctuation of fO2 and fH2O (Harlov et al. 2006), (Fig 6).

Fig 6. BSE images of Ba-mica grain (Left) and Mn-rich ilmenite replacing rutile mantled by Al-rich titanite (Right).

Mineral textures and chemistry Textural investigation of metamorphic minerals can be a powerful tool when trying to decipher the timing of their growth and can give important clues to the conditions that were prevailing during different growth stages. Combining microprobe analysis with textural observations allows for thermodynamical testing of hypotheses regarding metamorphic evolution against empirically derived chemical data. The text below highlights some important textural and analytical observations in the studied paragneisses (with emphasis on samples 19 and 19a).

Phengite is scarce but tiny grains are found preserved in the matrix and as inclusions in garnet. Commonly replaced by biotite or less commonly paragonite the remaining phengite has largely been consumed and later growing minerals including plagioclase and melt related quartz-alkali-feldspar symplectites are growing across the outline of previously coherent phengite grains (Fig 7). High Si atoms per formula unit (a.p.f.u.) in the range of 3.30-3.34 are found both in matrix grains and garnet inclusions (Fig 8). Analyzed phengite display compositional zonation with Si a.p.f.u. decreasing towards the rims (Fig 9; Table 2).

Biotite is commonly found as well preserved grains growing in the matrix and often in close association with garnet and sillimanite (melanocratic areas). These generally show intermediate to high values for Fe/(Fe+Mg) (XFe= c. 0.33-0.44). Anhedral-subhedral garnet hosts abundant inclusions including biotite with substantially lower iron content (XFe < 0.30). Biotite occurring in a biotite-quartz symplectite (Fig 9) in sample 19a gave an average XFe value of 0.38. The highest value (XFe=0.55) was obtained from biotite replacing phengite (Fig 9) which also displays significant Ti enrichment (Ti a.p.f.u=0.39), roughly two times higher than Ti levels common for biotites occurring in the melanocratic matrix. Analyzed biotite enclosed within garnet shows low to no Ti substitution (Fig 10; Table 3).

Late plagioclase from paragneiss matrix gave coherent results of XAn in the range of 0.38- 0.40. In one sample plagioclase present as inclusions in garnet (Fig 11) ranges from almost pure albite to albite dominated compositions. Analyses of plagioclase from quartz-plagioclase symplectites (Fig 11) and grains overgrowing remnant phengite yielded XAn of c. 0.56. (Table 4).

Fig 7. BSE image showing remnant phengite (outlined in white) replaced by biotite and overgrown by plagioclase. Tiny kyanite grains are enclosed within the biotite. Intergrowths of sillimanite with quartz, and quartz-feldspar symplectites are found in conjunction with garnet and phengite.

Fig 8. BSE image showing phengite with high Si contents (highest Si a.p.f.u.>3.30) in a polyphased garnet inclusion.

Fig 9. BSE images showing biotite quartz symplectite (left) and titanium rich, high XFe biotite replacing zoned phengite (right) in sample 19a. Spot analysis are marked with colors and corresponding values for Si a.p.f.u. in phengite are shown in the legend.

Fig 10. Ti contents plotted against XFe in analyzed biotite. Red ellipse marks biotite enclosed in garnet. Blue marks matrix biotite and biotite intergrown with quartz

Fig 11. BSE images showing plagioclase with high anorthite component intergrown with quartz (left) and high albite plagioclase enclosed in garnet together with quartz (right).

Table 2. Representative phengite compositions in sample 19 and 19a.

Phengite Sample 19a 19a 19a 19a 19 19a 19 19 19 19 19 19 19 19a 19 19 19a in in in Position matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix garnet garnet garnet SiO2 46,31 47,44 46,30 49,35 45,68 49,84 45,91 48,10 47,29 46,12 46,37 46,33 51,32 46,91 47,39 46,39 46,92 TiO2 0,01 0,20 0,34 0,06 1,13 0,03 1,66 0,11 0,07 1,46 1,32 1,86 0,06 0,18 1,69 1,03 0,16 Al2O3 29,66 29,68 29,71 31,30 31,40 31,41 31,62 31,84 32,43 32,45 32,62 32,71 32,84 32,87 32,90 32,91 33,00 FeO 5,07 4,03 3,37 4,07 1,24 1,60 1,25 4,01 3,17 1,25 1,21 1,15 0,78 2,65 1,19 1,13 1,98 MnO 0,02 0,00 0,00 0,09 0,03 0,00 0,00 0,03 0,00 0,01 0,00 0,03 0,00 0,00 0,02 0,00 0,00 MgO 1,41 2,78 2,55 1,61 1,17 1,29 1,25 1,76 1,71 1,28 1,29 1,11 0,78 2,00 1,14 1,05 1,67 CaO 0,08 0,14 0,02 0,12 1,90 0,17 0,07 0,13 0,10 0,08 0,06 0,07 1,23 0,05 0,04 0,08 0,04 Na2O 0,09 0,12 0,15 0,08 0,16 0,29 0,19 0,19 0,21 0,22 0,19 0,24 1,17 0,15 0,16 0,16 0,19 K2O 7,75 7,01 10,48 6,48 9,08 8,33 10,16 9,59 10,20 10,63 10,77 10,17 8,45 10,80 10,42 10,46 10,85 Total 90,41 91,41 92,94 93,16 91,81 92,96 92,12 95,76 95,19 93,50 93,83 93,68 96,63 95,61 94,96 93,22 94,81 Si 3,25 3,26 3,20 3,30 3,15 3,34 3,16 3,20 3,17 3,13 3,14 3,13 3,31 3,14 3,16 3,15 3,15 Al 2,46 2,41 2,42 2,47 2,55 2,48 2,56 2,50 2,56 2,60 2,60 2,61 2,50 2,59 2,58 2,63 2,61 Ti 0,00 0,01 0,02 0,00 0,06 0,00 0,09 0,01 0,00 0,07 0,07 0,09 0,00 0,01 0,08 0,05 0,01 Fe 0,30 0,23 0,19 0,23 0,07 0,09 0,07 0,22 0,18 0,07 0,07 0,07 0,04 0,15 0,07 0,06 0,11 Mn 0,00 0,00 0,00 0,01 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Mg 0,15 0,29 0,26 0,16 0,12 0,13 0,13 0,17 0,17 0,13 0,13 0,11 0,07 0,20 0,11 0,11 0,17 Ca 0,01 0,01 0,00 0,01 0,14 0,01 0,01 0,01 0,01 0,01 0,00 0,01 0,09 0,00 0,00 0,01 0,00 Na 0,01 0,02 0,02 0,01 0,02 0,04 0,02 0,03 0,03 0,03 0,03 0,03 0,15 0,02 0,02 0,02 0,03 K 0,69 0,62 0,92 0,55 0,80 0,71 0,89 0,82 0,87 0,92 0,93 0,88 0,70 0,92 0,89 0,91 0,93 Total 6,87 6,84 7,04 6,74 6,92 6,80 6,93 6,96 6,99 6,97 6,97 6,92 6,86 7,03 6,92 6,94 7,01 Mineral formula calculated on the basis of 11 oxygens

Table 3. Representative biotite compositions in sample 19 and 19a.

Biotite Sample 19a 19 19a 19a 19 19a 19 19a 19a 19a 19 19a 19 19 19 19a 19a bt-qz bt-qz in in in Position matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix matrix sympl. sympl. garnet garnet garnet SiO2 37,56 36,79 36,32 36,31 37,55 37,27 37,66 35,93 36,42 36,39 38,25 36,88 38,48 37,81 38,29 37,44 36,78 TiO2 7,17 4,50 3,64 3,66 3,95 4,12 3,87 3,68 3,85 4,06 2,93 3,68 2,83 3,22 3,29 0,03 2,08 Al2O3 17,32 17,51 18,02 17,38 17,56 17,52 17,29 17,21 17,13 16,95 17,76 17,07 17,48 17,73 17,51 18,89 18,17 FeO 16,77 15,77 16,10 14,82 14,47 14,68 13,82 14,37 14,28 14,22 13,18 13,18 12,34 10,64 10,07 12,48 11,40 MnO 0,00 0,02 0,04 0,11 0,05 0,02 0,01 0,02 0,00 0,07 0,00 0,03 0,03 0,03 0,03 0,04 0,09 MgO 7,58 10,22 11,64 11,78 12,26 12,33 12,49 13,10 13,26 13,40 13,87 13,98 14,31 15,26 15,60 15,93 16,55 CaO 0,02 0,04 0,09 0,08 0,03 0,05 0,03 0,03 0,02 0,04 0,03 0,01 0,04 0,04 0,04 0,02 0,23 Na2O 0,06 0,10 0,16 0,13 0,05 0,00 0,11 0,12 0,05 0,07 0,05 0,03 0,09 0,10 0,08 0,05 0,10 K2O 9,88 9,35 9,30 9,70 9,68 9,66 9,63 9,76 9,66 9,54 9,53 9,73 9,24 9,57 9,16 8,59 9,30 total 96,36 94,31 95,30 93,98 95,60 95,65 94,91 94,22 94,67 94,74 95,60 94,59 94,84 94,40 94,07 93,47 94,70 Si 2,74 2,74 2,68 2,72 2,73 2,72 2,75 2,68 2,69 2,69 2,75 2,71 2,77 2,73 2,75 2,74 2,67 Ti 0,39 0,25 0,20 0,21 0,22 0,23 0,21 0,21 0,21 0,23 0,16 0,20 0,15 0,17 0,18 0,00 0,11 Al 1,49 1,53 1,57 1,53 1,51 1,51 1,49 1,51 1,49 1,48 1,50 1,48 1,48 1,51 1,48 1,63 1,56 Fe 1,02 0,98 1,00 0,93 0,88 0,89 0,84 0,90 0,88 0,88 0,79 0,81 0,74 0,64 0,61 0,76 0,69 Mn 0,00 0,00 0,00 0,01 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,01 Mg 0,82 1,13 1,28 1,31 1,33 1,34 1,36 1,46 1,46 1,48 1,49 1,53 1,54 1,64 1,67 1,74 1,79 Ca 0,00 0,00 0,01 0,01 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,02 Na 0,01 0,01 0,02 0,02 0,01 0,00 0,02 0,02 0,01 0,01 0,01 0,00 0,01 0,01 0,01 0,01 0,01 K 0,92 0,89 0,88 0,93 0,90 0,90 0,90 0,93 0,91 0,90 0,87 0,91 0,85 0,88 0,84 0,80 0,86 Total 7,40 7,54 7,64 7,65 7,57 7,58 7,57 7,71 7,67 7,66 7,57 7,65 7,55 7,60 7,55 7,68 7,73 XFe 0,55 0,46 0,44 0,41 0,40 0,40 0,38 0,38 0,38 0,37 0,35 0,35 0,33 0,28 0,27 0,31 0,28 Mineral formula calculated on the basis of 11 oxygens

Table 4. Representative plagioclase compositions in sample 19 and 19a

Plagioclase Sample 19 19 19 19 19a 19a 19a 19a 19a 19a 19a 19a 19a 19 19 19 in in in in qz-pl across across Position matrix matrix matrix matrix matrix matrix matrix matrix matrix garnet garnet garnet garnet sympl. phg phg SiO2 67,33 65,74 61,37 59,02 57,94 57,35 57,32 57,09 56,95 56,93 56,93 56,76 56,55 55,12 53,91 53,80 TiO2 0,00 0,01 0,00 0,01 0,02 0,01 0,02 0,04 0,01 0,01 0,19 0,01 0,00 0,00 0,00 0,01 Al2O3 19,73 21,17 23,70 26,32 26,27 26,31 26,26 26,40 26,12 26,22 26,15 26,13 26,40 27,52 28,92 28,38 FeO 0,59 0,40 0,46 0,23 0,04 0,08 0,05 0,06 0,03 0,03 0,79 0,05 0,02 0,02 0,41 0,25 MnO 0,02 0,00 0,06 0,00 0,00 0,00 0,00 0,00 0,09 0,01 0,01 0,02 0,00 0,00 0,02 0,00 MgO 0,04 0,04 0,03 0,03 0,03 0,03 0,01 0,00 0,00 0,00 0,44 0,00 0,00 0,00 0,04 0,01 CaO 0,91 2,61 5,39 7,26 8,44 8,27 8,53 8,30 8,39 8,38 8,02 8,29 8,59 11,13 11,59 11,36 Na2O 11,36 10,17 8,31 7,36 6,72 7,09 6,76 6,97 6,79 7,05 6,80 7,00 6,97 4,75 5,04 5,19 K2O 0,02 0,13 0,37 0,15 0,30 0,13 0,31 0,20 0,19 0,21 0,61 0,24 0,21 0,07 0,12 0,13 Total 100,02 100,26 99,69 100,39 99,77 99,28 99,26 99,07 98,57 98,84 99,94 98,49 98,75 98,62 100,05 99,12 Si 2,96 2,89 2,74 2,63 2,60 2,59 2,59 2,58 2,59 2,58 2,57 2,59 2,57 2,51 2,44 2,46 Ti 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,01 0,00 0,00 0,00 0,00 0,00 Al 1,02 1,10 1,25 1,38 1,39 1,40 1,40 1,41 1,40 1,40 1,39 1,40 1,42 1,48 1,54 1,53 Fe 0,02 0,01 0,02 0,01 0,00 0,00 0,00 0,00 0,00 0,00 0,03 0,00 0,00 0,00 0,02 0,01 Mn 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Mg 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,03 0,00 0,00 0,00 0,00 0,00 Ca 0,04 0,12 0,26 0,35 0,41 0,40 0,41 0,40 0,41 0,41 0,39 0,40 0,42 0,54 0,56 0,56 Na 0,97 0,87 0,72 0,63 0,58 0,62 0,59 0,61 0,60 0,62 0,59 0,62 0,61 0,42 0,44 0,46 K 0,00 0,01 0,02 0,01 0,02 0,01 0,02 0,01 0,01 0,01 0,04 0,01 0,01 0,00 0,01 0,01 Total 5,02 5,00 5,01 5,01 5,00 5,02 5,01 5,02 5,01 5,03 5,04 5,03 5,03 4,96 5,01 5,01 XAn 0,04 0,12 0,26 0,35 0,40 0,39 0,40 0,39 0,40 0,39 0,38 0,39 0,40 0,56 0,56 0,54 XAb 0,96 0,87 0,72 0,64 0,58 0,60 0,58 0,60 0,59 0,60 0,58 0,60 0,59 0,43 0,44 0,45 XOr 0,00 0,01 0,02 0,01 0,02 0,01 0,02 0,01 0,01 0,01 0,03 0,01 0,01 0,00 0,01 0,01 Mineral formula calculated on the basis of 8 oxygens

Garnet has preserved melt related inclusions and textures (Fig 12 A-F). Striking wormlike quartz inclusions occupy garnet cores to outer cores and embayments consisting of quartz and biotite extend into a few garnets (Fig 12 A). These textures along with intergrowths of alkali feldspar and quartz present as inclusions in subhedral garnets (Fig 12 B & D) are interpreted as melt crystallization products. Occasionally such quartz-feldspar inclusions exhibit negative crystal shapes (Fig 12 D). Seemingly pure mica inclusions also exhibit negative crystal shape (Fig 12 C). Other melt related garnet inclusions consist of quartz and feldspar extending from remnant grains of white mica (phengite and paragonite). Commonly these inclusions are surrounded by radial cracks that are likely caused by differential expansion (Fig B & E). In Fig 12 B, melt has penetrated the surrounding cracks before crystallizing and is seen connecting two inclusions in the hosting garnet. Pure quartz inclusions surrounded by radial fractures and polycrystalline quartz inclusions (Fig 12 B-C & F) could represent coesite pseudomorphs as similar textures have been described in relation to the volume expansion occurring at the coesite-quartz phase transition (e.g. Gilotti & Ravna 2002).

Filling fractures running through subhedral garnets, surrounding the rims and occasionally forming “pools” adjacent to these garnets are tiny sillimanite needles intergrown with quartz (Fig 12 G). In sample 19a these intergrowths often occupy the intersecting surfaces of two or more neighboring garnet crystals in dense (micro inclusion bearing) garnet bands. The shape and site of occurrence of the intergrowths indicates a melt related origin. In the same section a large prismatic, presumably synkinematic, sillimanite porphyroblast is growing across a large anhedral garnet or cluster of garnets (Fig 12 H). The garnet is elongated in the same direction as the sillimanite grain. Other phases included in this and similar garnets are biotite, rutile, feldspars, graphite, quartz and ilmenite.

Fig 12. BSE images and photomicrographs of (A) wormlike quartz inclusions in garnet with quartz and biotite embayments (B) polyphase inclusion (intergrown quartz-feldspar and white mica) in garnet, melt crystallization product is connecting two inclusions (C) quartz inclusion with radial cracks between negatively shaped quartz and mica inclusions (D) quartz and feldspar inclusion with slightly negative crystal shape (E) polyphase inclusion with quartz and white mica surrounded by radial cracks (F) polycrystalline quartz inclusion (G) sillimanite-quartz intergrowth filling garnet fracture (H) sillimanite grain growing across anhedral garnet.

Micro inclusions Enigmatic micro inclusions in garnet occur in two paragneiss samples. In one anatectic sample (19a) they occur regularly whilst in a finegrained paragneiss (sample 21) more sparsely. Generally, the inclusions are between c. 0.1-10 micrometers in size and often concentrated in swarms within the garnet cores, occasionally outlining a pattern closely resembling the crystal shape of euhedral garnet. The appearance of the inclusions vary between colourless and translucent through brown tinged to almost completely opaque. Some micro inclusions exhibit negative crystal shapes whilst the shapes of others vary between seemingly almost perfect cubes through cuboidal and spheroidal to bleb like shapes (Fig 13). WDS analyses of the micro inclusions yielded insufficient oxide totals (Wt % <60). Results obtained from EDS-analysis indicate that the composition of the micro inclusions vary from carbonate to carbon dominated. A more detailed description and results from micro Raman analysis is presented in the following section.

Fig 13. Photomicrographs with increased magnification (A-B) and close-ups (C and D) of micro inclusions in garnet (sample 19a).

Garnet chemistry Garnets analyzed in two paragneiss samples (19 and 19a) and one calcic schist sample (19b) all have compositions dominated by almandine. The obtained almandine values are c. 54-59 % (Fig 14 A), 63-68 % (Fig 14 B) and 63-72 % (Fig 14 C-E) in the calcic schist and the two paragneiss samples respectively. All garnet profiles show an evenly distributed spessartine content (Fig 14 A-E) which is indicative of high temperature homogenization (Yardley 1977). The calcic schist garnet is comparatively rich in spessartine (c. 6.5-7 %) and grossular (c. 19- 24 %) which is matched by a relatively low pyrope contents of c. 13-18 % (Fig 14 A). The paragneiss samples have a different chemical signature with low grossular and spessartine levels in the matrix of analyzed garnet. The compositions differ slightly between the paragneisses with lower garnet matrix grossular (3-4.5 %) and spessartine (1.5-2 %) levels in one paragneiss (Fig 14 C-E) compared to the second sample (Grs=5.5-8 % and Sps=2-2.5 %), (Fig 14 B). Pyrope contents are c. 17-33 % in garnet from one paragneiss (19a) and c. 25-30 % in the second paragneiss sample (19).

Garnets in the calcic schist generally shows very even compositions with occasional minor elevations in almandine accompanied by a slight decrease of pyrope in the outermost rims (Fig 14 A). The outer rims also display a small increase in grossular that could indicate that the composition of the garnet rims has been altered by the presence of melt (Lang & Gilotti 2007). Such retrogressive and possibly melt related features are clearly evident also in the paragneisses in which some garnet crystals show a comparatively significant increase of Fe/Mg-ratios and grossular levels which is gradational but by appearance has left the cores unaffected (Fig 14 C). This pattern was recognized in both analyzed paragneiss samples but the degree to which the garnets are affected is varying. Several garnets in one paragneiss (19a) where found to contain small coherent areas of elevated grossular (c. 5-12.5 %) and lower pyrope (c. 18-25 %) content compared to the garnet matrix composition (Fig 14 C-E). Whilst some of these likely represent remnants of garnet that grew prior to its hosts there are small irregularities bordering inclusions that might have been caused by reaction between the inclusions and surrounding garnet (Lang & Gilotti 2007).

Fig 14. Compositional profiles showing variation in pyrope, grossular, almandine, spessartine component and Fe/(Fe+Mg) in (A) anhedral garnet in calcic schist sample 19b (B) subhedral inclusion free garnet in paragneiss sample 19 (C) subhedral garnet hosting micro inclusions in the core region in paragneiss sample 19a (D) cluster of garnet hosting remnants of older garnet in paragneiss sample 19a (E) large inclusion rich garnet in paragneiss sample 19a. Gaps in the profile are due to inclusions and/or analyses yielding insufficient totals. BSE images of corresponding garnets with profiles outlined are shown to the right.

5.2. Raman spectroscopy Raman spectra obtained from micro inclusions in garnet cores of the paragneiss reveal variations in mineral content. Results, presented in the following section, show the inclusions to be mostly made up of carbonates, carbon allotropes and phase composites. Several analyzed inclusions could not be positively identified due to fluorescence and analytical difficulties caused by the small grain size. To avoid any confusion regarding the position of peaks and luminescence all presented spectra for carbon allotropes were shot using a green (532 nm) laser.

Carbonates Micrometer sized inclusions (usually <10 µm) consisting of carbonate were identified by the, apart from peaks related to hosting garnet, dominating presence of Raman active modes at -1 -2 1084-1090 cm corresponding to the symmetric stretching (ν1) of (CO3) (Fig 15). Minor additional bands at 302-314 and c. 730, often partly obscured by the overlap with host garnet spectra, together with compositional indication from EDS spectra makes it possible to identify the carbonate inclusions as mainly consisting of magnesite-siderite solid solution. Quantitative compositions of the carbonate species could however not be assigned. Inclusions -1 found to be composites of carbonate, CO2 indicated by the ν1 mode present at 1282 cm and -1 -1 the ν2 at 1384 cm together with N2 indicated by the sharp peak at 2327 cm were also found (Fig 15; see also Andersen et al. 1990; Mposkos et al. 2009; Van den Kerkhof & Olsen 1990).

Fig 15. Representative Raman spectra (633 nm laser) of magnesite-siderite solid solution dominated, carbonate micro inclusion (top) and of carbonate, CO2, N2 composite inclusion (bottom). Peaks are labeled with wavenumber (cm-1) and corresponding phase.

Graphite Micro inclusions found to consist mainly or partly of graphite commonly appear brown-black in plane transmitted light and crystals are typically cuboidal in shape. Sizes vary from less than 1 µm up to c. 15 µm. These were identified in multiple spectra by the presence of the G active mode at 1580-1585 cm-1 and usually a substantial degree of disordering characterized by the D1 peak at c. 1350 cm-1 (Fig 16). The D1 mode is dispersive, meaning its position varies with excitation energy (laser wavelength). The D1 peak in disordered graphite is in several spectra accompanied by a shoulder on the high wavenumber side of the G peak (c. 1610 cm-1) often referred to as the D2 peak (not pictured) which is also indicative of disordered graphite (Ferrari & Robertson 2000; Pimenta et al. 2007; Pócsik et al. 1998). Some inclusions were found to contain disordered graphite occurring together with carbonate (Fig 16).

Fig 16. Representative Raman spectra (532 nm laser) of disordered graphite (G- and D1-mode) together with carbonate. Peaks are labeled with wavenumber (cm-1) and corresponding phase.

Diamond In total 3 diamond inclusions in garnet, sized 3.5-7 µm across, were identified in different garnets by Raman active modes spanning 1328-1330 cm-1 (Fig 17). The spectra reveal a varying degree of graphitization indicated by the intensity of the G mode of graphite. Common for all 3 diamond inclusions is that the hosting garnets exhibit relatively thick inclusion free rims. These rims surround inclusion swarms in the core region of the garnets within which single diamonds are found. The diamonds appear brown when viewed in plane transmitted light and are of a cuboidal shape (Fig 17).

Raman peaks were fitted using a Voigt function in Fityk 0.9.8 (Wojdyr 2010). The diamond spectra display significant peak widening with FWHM 9.5-16 cm-1 as well as downshifting of peak positions from the ideal position at 1332 cm-1 (Solin & Ramdas 1970). These are two features common to metamorphic diamond and can be caused by small diamond grain sizes (Yoshikawa et al. 1995), compressive or tensile stress in the diamond crystal lattice, some degree of structural disorder (Kalbitzer 2001) and/or to some extend by chemical impurities such as nitrogen (Surovtsev et al. 1999).

A closer examination of the diamond inclusions using BSE and SE imaging (Fig 18 A-B) reveals additional characteristic and important features. The hardness contrast between diamond and hosting garnet cause relief and smooth depressions surrounding the diamond inclusions, visible in SE images, after polishing. BSE images show at least one diamond bearing inclusion to partly consist of another bright colored phase (Fig 18 B) from Raman spectroscopy concluded to be carbonate.

Fig 17. Representative Raman spectra (532 nm laser) of 3 analyzed diamond inclusions (left) with corresponding photomicrographs (diamonds centered) in reflected light (middle) and transmitted light (right). The spectra reveal a varying degree of graphitization and downshift of diamond peaks. Peaks are labeled with wavenumber (cm-1) and corresponding phase.

Fig 18. BSE images (left) and SE images (right) of diamond inclusions. (A) Largest identified diamond (c. 7 µm across) with smooth diamond-garnet contact (B) Large inclusion, to the left of the image, consisting of diamond and minor carbonate (photo: J. Majka).

6. Geothermometry Cation exchange reactions between coexisting garnet and biotite were used to assess equilibration temperatures. Representative homogenized garnet core compositions were selected from profiled garnets to minimize the effect of local disequilibria. The temperatures were calculated at a pressure of 1 GPa which was selected as a minimum estimate based on observations of kyanite and phengite in the selected samples. Calculation results for seven commonly used calibrations are presented below.

6.1. Garnet-Biotite thermometer Calculated temperatures (Table 5.) for the garnet-biotite exchange at 1 GPa in the paragneisses show inconsistencies between different calibrations and mineral analyses. Extreme temperatures obtained from biotite replacing phengite are likely unrealistic, possibly due to disruption by high Ti substitution, and will therefore be disregarded further in this section. Biotite enclosed in garnet consistently gave lower temperatures (540-690oC) than matrix biotite adjacent to garnet (650-1090oC). Out of the seven calibrations, the Bhattacharya et al. (1992) and Dasgupta et al. (1991) can presumably be preferred since these calibrations rely on the most recent thermodynamical data which considers non-ideal mixing in garnet along with Ti and Al interaction in the octahedral sites of biotite. The Bhattacharya et al. (1992) calibration produced equilibration temperatures in the range of 560-810oC. The commonly cited calibration by Perchuk & Lavrent’eva (1983) yielded very similar results for all biotites. This calibration proved comparatively accurate when eleven calibrations were empirically tested by Wu & Cheng (2006). Temperatures calculated using the calibration of Dasgupta et al. (1991) are generally slightly higher, ranging between 550-920oC, whilst the highest temperatures (>930 oC) were obtained using the older calibrations of Thompson (1976), Ferry & Spear (1978) and Hodges & Spear (1982). The latter three calibrations consistently yielded c. 100oC higher temperatures for biotite adjacent to garnet than newer calibrations.

Table 5. Calculated temperatures for garnet-biotite exchange at 1.0 GPa using analysed biotites, listed by occurrence, and representative homogenized garnet core compositions in samle 19 and 19a with the calibrations of Thompson (1976), Holdaway & Lee (1977), Ferry & Spear (1978), Hodges & Spear (1982), Perchuk & Lavrent’eva (1983), Dasgupta et al. (1991) and Bhattacharya et al. (1992). Temperatures calculated with the latest calibration (Bhattacharya et al. 1992) using two non-ideal mixing parameters for garnet (Ganguly & Saxena 1984; Hackler and Wood 1989) are listed separately.

7. Thermodynamic phase equilibrium modeling In order to evaluate favorable PT-conditions and the timing of different mineral growth stages thermodynamic modeling using Perple_X software was employed. Results from calculations carried out in two independent steps in the NCKFMASTHC- (model 1) and NCKFMASTH

(model 2) system using bulk rock chemistry (Wt %: SiO2=49.8 Al2O3=24.6 CaO=2.45

FeO=9.63 K2O=4.25 MgO=2.94 Na2O=1.98 Ti=1.2 LOI=0.3) and quantitative mineral chemistry analysis (WDS) from one paragneiss sample are presented below. Pseudosections and modelled phase equilibria are presented stepwise in section 7.1-7.2 starting with the peak mineral assemblage and onset of exhumation. The deduced exhumation path is presented in section 7.3.

The Mn component is ignored to avoid inconsistences in the garnet solution model. H2O is treated as a saturated fluid phase in model 1 and XCO2 is set to 0.05, enough to stabilize carbonates together with garnet on the high pressure low temperature side of the pseudosection. SiO2 is treated as a saturated phase component in both models.

Prograde and peak pressure stages could not be modeled using garnet compositions. In order to constrain lowest melting temperatures and to evaluate mineral reactions involving melt the second model was constructed. The water contents in model 2 is set to the bulk rock loss on ignition value (0.3 %) which is estimated to represent the maximum water content at final melt crystallization. Calculations are presented in the 0.3-1.4 GPa, 600-1000oC range, complying to specified limitations of the melt(HP) solution model (White et al. 2001).

7.1. Model 1 The calculated peak assemblage is Ph Grt Ky Jd Rt Coe Mgs-Sid (Dia) considering a minimum pressure of c. 3.1 GPa to stabilize diamond at temperatures exceeding 650oC

(McClelland & Lapen 2013) and assuming a maximum pressure of 4.5 GPa. The onset of exhumation is inferred from the Si a.p.f.u. in phengite with predicted simultaneous volume decrease of kyanite and the calculated stability of carbonates together with garnet (Fig 19). A further pressure decrease to the lower temperature side of the Bt Pl Ph Grt Rt field, is indicated by anomalous garnet compositions interpreted to be older cores hosted in younger garnet stable under similar conditions as phengite rims (Fig 20). The size and occurrence of retained garnet cores are considered too small to cause any significant change to the effective bulk chemistry. Equilibrated garnet matrix compositions are not stable in this model.

7.2. Model 2 Melting is predicted at 850-890oC (Fig 21) with a calculated melt volume of 5-15 % (Fig 22 A) stable with paragenetic garnet- (XMg and XCa) biotite- (XFe Bt) and plagioclase- (XAn) compositions in the melt Pl Grt Sil San Rt Ilm field (Fig 22 B-D). Modeled peak temperature of 880-910 oC at 1.08-1.15 GPa is inferred from respresentative equilibrated Ca and Mg garnet compositions. Final exhumation and melt crystallization following temperature decrease is constrained by XFe for biotite in quartz-biotite symplectites and the high XAn for matrix plagioclase (Fig 22 C).

7.3. Exhumation PT path The presented exhumation path (Fig 23) is derived from combining modeled phase equilibrium of garnet, phengite, biotite and plagioclase in model 1 and model 2. Considering fluid saturation at high pressures followed by melting and fluid escape upon exhumation, as indicated by mineral textures and chemistry, the two models can be combined. The exhumation path is defined by reequlibration of phengite, remnant garnet core-, representative garnet matrix-, biotite- and plagioclase compositions in agreement with observed mineral assemblages and analyzed mineral chemistry.

The modeled exhumation started in the stability field of diamond with the peak assemblage Phg Mgs-Sid Grt Ky Jd Rt Coe (Dia) and continued to the Bt Pl Phg Grt Rt (Qz) field, likely passing through assemblages including paragonite. Between c. 1.5-1.1 GPa the temperature increased to c. 880-910oC causing extensive partial melting, resetting of garnet chemistry and causing almost complete consumption of phengite. At the final stage of exhumation growing matrix biotite and plagioclase were continuously equilibrated.

Fig 19. P-T pseudosection calculated for paragneiss bulk chemistry in the NCKFMASTHC system. H2O specified as a saturated fluid phase, XCO2=0.05. Si is specified as a saturated component. The peak assemblage field (Ph Grt Ky Jd Rt Coe Mgs-Sid (Dia)) is outlined in red. Numbers correspond to: (1) Bt Phg Pa Grt Gl Ab (2) Bt Pl Phg Grt Ky San (3) Bt Pl Phg Pa Grt Zo Ilm (4) Bt Pl Phg Pa Ma Grt Ilm (5) Bt Pl Phg Pa Grt Ky Ilm. Diamond-Graphite- and Coesite-Quartz reaction lines from McClelland & Lapen (2013).

Fig 20. Calculated mineral composition curves (isopleths) for representative and highest XCa garnet composition anomalies interpreted to represent early garnet equilibration, phengite Si a.p.f.u. and predicted vol.% of kyanite. XCa, XMg in garnet and Si a.p.f.u in phengite rims are stable at 720-790 oC and 1.28-1.48 GPa together in the stability field of Bt Pl Phg Grt Rt. (Left) Calculated peak assemblage is Ph Grt Ky Jd Rt Coe (Mgs-Sid). Grey marks the predicted stability field of Bt, dark grey of Bt+Sil and blue of garnet together with carbonates on the calculated pseudosection for model 1 (Right).

Fig 21. P-T pesudosection (NCKFMASTH system) including melt solution model. H2O=0.3 wt %, Si specified as saturated component. Numbers corresponds to: (1) Bt Pl Grt Ky San Rt Ilm (2) Bt melt Pl Grt Sil San Rt Ilm (3) Bt melt Pl Grt Sil Crd San Ilm.

Fig 22. (A) Predicted melt volume in P-T space of second model, melt first stable at c. 850oC (B) Representative flat garnet matrix composition intersecting in the Bt melt Pl Grt Sil Sa Rt stability field (C) Predicted stability of XAn in plagioclase and XFe in biotite representative for mineral analysis. (D) Typical analyzed garnet compositions intersecting at 880-910oC and 1.08-1.15 GPa with analysed high XFe (0.44-0.54) matrix biotite in the Bt melt Pl Grt Sil San Rt Ilm field. XAn of matrix plagioclase is stable at slightly lower pressures.

Fig 23. PT path for exhumation of Snasahögarna paragneiss shown on model 1 with marked PT space of model 2. (1) corresponds to stability of maximum Si a.p.f.u. in phengite intersecting mineral assemblages including garnet and carbonates (2) corresponds to Ca and Mg phase equilibrium of remnant cores and phengite rim compositions. (3) corresponds to representative garnet (XMg and XCa), biotite (XFe) and plagioclase compositions stable with additional sillimanite, rutile and melt.

8. Discussion

8.1.Ultrahigh pressure metamorphism (UHP) UHP metamorphic provinces are found in association with major collisional orogens around the world. Examples include northeastern Greenland, the western Alps and the Himalayas. In the Scandinavian Caledonides areas where UHP rocks occur include the famous Western Gneiss Region (WGR) and Tromsø. In rocks subjected to extreme pressures related to subduction zones and continental collision peak metamorphic events can be recorded by UHP index mineral phases including coesite and metamorphogenic diamond (Gilotti 2013). The recent UHP constraints on eclogites of northern Jämtland (Janák et al. 2012) combined with the find of metamorphogenic diamond presented in this thesis points to a widespread UHP influence within the Seve Nappe Complex.

8.2. Metamorphogenic diamond

Microdiamond occuring in metamorphic rocks was first reported from the Kokchetav Massif in Kazakhstan (Rozen et al. 1972; Sobolev & Shatsky 1990) and later from China (e.g. Shutong et al 1992). Described occurrences in southern Europe include microdiamond found in the Greek Rhodope (Mposkos & Kostopoulos 2001) and in central Europe microdiamonds in gneisses of Erzgebirge, Germany (Nasdala & Massone 2000). In the Scandinavian Caledonides microdiamonds were initially described from the world famous UHP province in the western gneiss region of Norway (Dobrzhinetskaya, et al. 1995) and recently also in the Tromsø region of northern Norway (Janák et al. 2013). Common for many of these localities is that the diamonds occur as inclusions in rigid metamorphic minerals (mostly garnet or zircon), are commonly to some degree graphitized and often associated with carbonate phases. Micro Raman spectroscopy has proven to be a crucial tool in the identification and interpretation of metamorphogenic diamond (e.g. Perraki et al. 2009).

Diamonds found in the Snasahögarna gneiss are generally consistent with petrological, visual and analytical characteristics of metamorphic microdiamond in gneisses described from worldwide UHP localities. Appearance, variation in size, site of occurrence and Raman spectroscopy signal traits exclude confusion with artificial externally introduced diamond. The intimate association of diamond and disordered graphite with carbonate-gaseous composite inclusions indicates that the graphite-diamond garnet inclusions formed by precipitation from carbon saturated volatiles that were entrapped in garnet at UHP conditions. A similar process was recently suggested for low temperature and ultrahigh pressure, diamond-bearing sediments found in the Alps (Frezzotti et al. 2011).

The dominating occurrence of disordered graphite as opposed to diamond and partial graphitization of diamond described in this thesis indicates a strong graphitization process acting during exhumation. It seems highly likely that a high temperature stage causing partial melting would have contributed to such a process. The thick inclusions free rim surrounding all 3 identified diamond bearing garnet cores would have helped preserve the diamond inclusions.

8.3. Metamorphic evolution Fluid levels and fluid composition play a key role in metamorphic reactions. From microscopic and analytical studies a few central observations should be stressed regarding the modeled sample. The composition and appearance of micro inclusions occurring in garnet points to water and carbon saturation at UHP conditions. Phengite, which is stable at great depths, would have acted as an important carrier of H2O during subduction. On the contrary, later paragenesis suggests low levels of H2O and CO2 in the studied paragneisses with a very restricted occurrence of carbonates in the rock matrix, no observed chloritization of biotite and only tiny remaining remnants of metastable phengite. This indicates that fluids present in the rock at the UHP stage largely escaped during exhumation.

Partial melting is inferred from macroscopic anatectic segregations and microscopic textures including garnet inclusions and quartz-biotite embayments in garnet along with melt related symplectites observed in the paragneiss matrix. A substantial part of the melt likely formed due to dehydration reactions involving phengite as remaining grains has largely been consumed and are commonly observed in conjunction with melt related textures. Comparing the two produced models (Fig 19; Fig 21) it is clear that the petrology of the samples in general are best modeled when considering the possibility of melt in a fluid restricted reaction environment. At such conditions kyanite and alkali feldspar stability is favored over that of white mica and ilmenite is stable at higher temperatures. Only minor mineralogical remnants from the presumed peak pressure conditions are preserved in the form of very restricted amounts of phengite and kyanite together with garnet hosting diamond bearing micro inclusions. Jadeite has not been identified but likely reacted to form albite and quartz upon pressure decrease. High albite plagioclase and quartz found together enclosed in garnet could represent such reaction products that were restricted from further reequilibration by the hosting garnet. Polycrystalline quartz inclusions in garnet surrounded by radial fractures is interpreted to represent previously occurring coesite. The following melt formation would have contributed to the conversion of coesite to quartz explaining why no coesite has yet been identified. Direct and indirect observations of the peak pressure mineral assemblage are generally in agreement with model 1 (Fig 19) continuing down through the coesite-quartz and across the jadeite-albite (+quartz) reaction curves.

Peak metamorphic conditions for the studied samples could not be constrained using thermodynamic modeling of analyzed garnet composition. This could have several possible explanations:

1. The obtained chemical bulk composition is not representative of the effective composition that was available for reactions involving garnet at the peak pressure stage. This problem can be resolved by calculating the local chemical composition in thin section portions of interest using gridded EDS analyses.

2. The garnet never equilibrated at peak pressure conditions.

3. The garnet chemistry was reequilibrated at a later stage. Reequilibration of garnet chemistry as a consequence of high temperatures is a well-known and early described process (Yardley 1977; Florence & Spear 1991), which possibly in combination with explanation 1 and/or 2, is favored by the author of this thesis. All profiled garnets from rocks of varying lithologies show even spessartine contents across the compositional profiles, a key trait of high temperature resetting. This scenario is supported by modeled phase equlibrium of garnet, biotite and plagioclase indicating temperatures of >880oC. Temperatures approaching this range are not contradicted nor fully confirmed by calculated garnet-biotite exchange temperatures although results are generally consistent with earlier published data (outlined in section 3.1) indicating equilibration at 660-770oC. The present study and previous geothermometrical studies obtained individual temperatures well exceeding 800oC. The modeled peak temperature (model 2) is in the high pressure end of the biotite stability field where highest XFe analysed biotite is in predicted equilibrium with garnet. Moreover, the sheer range of calculated garnet-biotite temperatures and analysed biotite compositions suggests either retrogression or equilibration at different temperatures. Textural evidence of melting and garnet rim composition indicating late melt interaction suggest that melt was still present at the final stage of garnet growth. Supporting this are sillimanite porphyroblasts grown syn-post anhedral garnet with even spessartine distribution containing inclusions of ilmenite and biotite. Intergrowths of sillimanite with quartz surrounding garnet rims and filling garnet fractures suggests retained high temperatures (at least >700oC based on paragenetic assemblage) post the growth of this garnet. Rare Ba-mica found in associated calcic schist may also indicate high temperatures. The composition resembles that of

47 oxykinoshitalite, earlier found in ultramafic volcanics (Kogarko et al. 2005; Manuella et al. 2012).

8.4. Exhumation processes Recently Hacker et al. (2013) published a review article on the formation and exhumation of ultra-high pressure terranes including numerical models proposed for selected provinces with varying characteristics around the world. Although these modern tectonic models provide possible explanations for the transportation of continental crust back from ultra-high pressure depths towards the surface, it is not clear how a late temperature increase to >880oC such as the one proposed for Snasahögarna in this text can arise. A similar although not as drastic temperature increase to >800oC causing partial melting during decompression of related units on Åreskutan was recently described by Klonowska et al. (2013). Their proposed explanation for Åreskutan was that the HP-UHP event was very short lived not leaving time for the rock to heat up until subsequent exhumation stages. This could possibly be valid also for the Snasahögarna gneisses.

A very recent model for the exhumation of young blueschist in the Yuli belt Taiwan suggests a highly interesting scenario (Sandmann et al. 2013). In this model, supported by geochronological evidence, the subducted ocean-continent boundary of the Chinese passive margin (Yuli belt) was rapidly exhumed as the overlying forearc lithosphere was extracted downwards. On a speculative basis such a scenario occurring in the Scandinavian Caledonides would allow mantle moved by suction forces to come into contact with partly exhumed units (e.g. the Seve Nappe Complex) from below causing a late rapid temperature increase affecting these units. In turn this could pose an alternative explanation to the occurrence of voluminous spinel peridotites in the Seve Nappe Complex which is not in contradiction with previous partial exhumation of UHP units by ductile return flow similar to that suggested for the Himalayan UHP terranes by Hacker et al. (2013) based on numerical models by Warren et al. (2006). This could possibly be tested with accurate temperature and geochronologial constraints coupled with trace element- and isotopic data from various Seve- and related lithological units.

8.5. Timing of subduction The in situ identification of metamorphogenic diamond in gneiss from the Snasahögarna area proves that the Middle Seve Nappe found in central Jämtland was subducted to depths of c. 120 km below the surface. Published and unpublished geochronological data reviewed here

48 points to subduction of these units in the latest Ordovician as earlier proposed by Majka et al. (2012); Ladenberger et al. (2012) and Klonowska et al. (2013). Subduction of continental crustal material to depths below 100 km is usually associated with continent collisions suggesting a similar setting for the peak pressure metamorphic stage observed in the Snasahögarna gneisses.

9. Conclusions The first discovery of metamorphogenic diamond in Sweden occurring in the garnet-kyanite- bearing gneiss from the Snasahögarna area is presented. This pelitic gneiss was subjected to ultrahigh pressure conditions reaching the diamond stability field. Mineralogy and thermodynamic modeling and bulk rock chemistry attest to a metamorphic history of fluctuations in fluid and gaseous components. The fluid saturated UHP stage was followed by later heating in a relatively dry environment. Temperatures reaching above 880oC at c. 1.1 GPa and simultaneous partial melting is constrained by mineral textures and thermodynamic modeling. Regionally, these results together with published geochronological data are in agreement with continental collisional episode causing subduction of the Seve Nappe Complex in Jämtland to depths up to c. 120 km at the end of the Ordovician.

Acknowledgements I am grateful to Jarek Majka who supervised this project and has been an invaluable continuous support and problem solver throughout this and other adventures. Maciej Manecki for most kindly lending me his knowledge, time, office space and access to state of the art analytical equipment at the AGH University in Kraków without which this thesis could not have been finished. David Gee for first introducing me to Caledonian geology and for his advising support. Students and former student colleagues which during the last years and especially during this project has been invaluable with your encouragement and companionship. This study is part of the Metamorphic Map of Sweden project funded by SGU.

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