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Evolution of granulites from MacRobertson Land, East Antarctica
by
Ian Scrimgeour
Thesis submitted for the degree of Doctor of Philosophy in The University of Adelaide (Faculty of Science)
April, 1994
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Abstract
and Granulite facies rocks in the northern Prince charles N{ountains an extensive Mid surrounding regions of MacRobertson Land form part of of the proterozoic (1000 Ma) regional granulite facies terrain which dominates much granulite terrains in the coast of East Antarctica. As one of the largest exposed processes which result in world, it provides an excellent opportunity to examine the of granulite formation, yet much of the metamorphic and Structural evolution to areas to the MacRobertson Land has remained relatively unknown in comparison east and west.
rocks on Jetty In the northeastern Prince Charles l\{ountlins granulite flcies and post-tectonic Peninsula consist of a supracrustal sequence intruded by syn- andT -7 granites, and preserve evidence for peak metamorphism (Ml) at -800oc '5 were kbars. In metapelitic rocks, M1 garnet- and sillimanite-bearing assemblages and symplectites which overprinted by the development of cordierite-bearing coronas dipping thrusts formed during the development of tight upright fol northern Reaction textures in wollastonite-scapolite calc-silicates from the the prince charles Mountains imply that following peak metamorphism at c.800oc of scapolite to terrain underwent near isobaric cooling, resulting in the breakdown aS the formation calcite and anorthite, and wollastonite to calcite and quartz, as well from of grossular coronas between wollastonite, scapolite and calcite. calc-silicates which is replaced by the Framnes Mountains on the Mawson Coast contain scapolite, and wollastonite-anorthite, anorthite-grossular and calcite-quartz symplectites' reaction textures grossular coronas occur between wollastonite and anorthite. These partially replaced by also imply a cooling-dominated history. Peak wollastonite is forms coronas around calcite and quartz, and then a second generation of wollastonite the calcite and quartz, suggesting that fluids may have played an important role in the evolution the Framnes Mountains. On Fox Ridge, 60 km southwest of Jetty Peninsula, late steeply dipping D2 high strain zones postdate the development of cordierite coronas, and are progressively defined by garnet-cordierite, sillimanite, and biotite-sillimanite assemblages, suggesting that the terrain isobarically cooled late in D2 during ongoing convergen t deform ation. The most plausible explanation for this evolution is that the terrain underwent at least two transient thermal events during a long-lived tectonic event between 1000 and 900 Ma. These relatively transient events were followed by isobaric cooling, and occurred at progressively lower pressures during ongoing convergent deformation, implying that the region underwent decompression during crustal thickening. The extent of decompression during high grade metamorphism decreases from east to west across the Prince Charles Mountains. The region appears to form the southern margin of a broad mountain belt which may have resulted from northeast-directed thrusting of the Proterozoic terrain over the Archean Vestfold Block during the Mid proterozoic. General implications of this study include the apparent transience of the thermal perturbations resulting in low to medium pressure granulite metamorphism, and the likely weakness of the lower crust during granulite metamorphism, which raises the possibility that upper and lower crustal strains may be significantly decoupled during high temperature metamorphism. Contents Abstract Contents Figures Tables. Acknowledgments ... Mineral abbreviations CHAPTER l: Intoductíon...... I 1.1 General 1 1.2 P-T paths in granuliæ tenains 3 1. 3 Proterozoic granulites : geodynamic considerations 4 1.4 The northern Prince Cha¡les Mountains: regional geological setting, previous work and relevant problems. 6 1.5 Thesis outline. 9 CHAPTER 2: The geology and structure of the Jetty Peninsula region, northern Princ e Charle s Mountains 11 2.I Introduction.... 11 2.2 Regional geological context..... 15 2.3 Lithologies..... 15 2.3.1Layered gneiss. 18 2.3.2Intrtsive rocks. 2l 2.3.3 Late Palaeozoic to Mesozoic dykes and sediments . . . .. 29 2.4 Structure 3T 2.4.1DL 31 2.4.2 D2 .. 35 243D3 38 2.4.4 D4 shear zones...... 40 2.4.5 Late brittle structures . 4l 2.5 Sm-Nd geochronology of the garnet granite...... 42 2.6 Geological history and age constraints 44 2.6.1 Mid Proterozoic deformation and metamorphism ...... 44 2.6.2 Pan- Afican thermal event. . 46 2.6.3 Lambert Graben development...... 47 2.7 Conelations and discussion 48 CHAPTER 3: Metamorphic evolution of pelitic granulites from Jetty Peninsula, northern Prince Charles Mountains'.... 53 3. 1 Introduction 53 54 3.2 Petrography of Pelites 3.2.1 Assemblage 1 ...'.. 55 3.2.2 Assemblage 2 ..... 59 3.2.3 Assemblage 3 ...... 60 3.2.4 Assemblage 4 ..... 60 3.2.5 Mineral chemistry .... 62 3.3 Interpretation of mineral parageneses...... 63 3.3.1 Assemblage 3 ...... 64 3.3.2 Assemblage 4 ...... 66 3.3.3 Assemblage 1 ...... 67 3.3.4 Assemblage 2 ..... 68 69 3.3.5 P-T estimates for Jetty Peninsula... ' 3.4 Discussion.. 70 CHAPTER 4: Metamorphic evolution of pelites from Trost Rocks, Amery lce Shelf 74 4.1 Introduction...... 74 4.2 Regional setting and geological relationships...'...... 74 4.3 Petrography 77 4.3. 1 Corundum-absent Pelites 78 4.3.2 Corundum-bearing pelites...... 83 4.3.3 Semi-pelitic assemblages ...... 83 4.3.4 Mineral chemistry ...... 84 4.4 Interpretation of reaction textures 8s 4.4.1 Reactions in the KFMASH and KFMASHTO systems 85 4.4.2 Conelations with Jetty Peninsula. 88 4.4.3T\erole of aÍlz0 in the stabilisation of biotite at Trost Rocks and Jetty Peninsula 89 4.4.4P-T evolution 92 4.5 Discussion .. 94 CHAPTER 5:Wollastonite-scapolite calc-silicates as indicators of the P-T-fluid history in MacRobertson Land ...... 96 5.1 Introduction...... 96 5.2 Calc-silicate gneisses on Else Platform, Jetty Peninsula...... 99 5.2.1 Geological setting .. . 99 5.2.2 Petrography 99 5.2.3 Mineral chemistry .. 103 5.2.4Interpretation of reaction textures and P-T history .... 104 5.3 Calc-silicates from Mt. McCarthy, Porthos Range 108 5.4 Barium silicates in NPCM calc-silicates. 109 5.4.1 Origins of NPCM calc-silicates...... Lt2 5.5 Calc-silicates from Rumdoodle, Framnes Mountains .. '. 113 5.5.1 Geological setting ..... 113 5.5.2 Petrography IT4 5.5,3 Interpretation of reaction textures...... tt7 5.6 The role of the fluid phase t23 5.7 Discussion... ..124 CHAPTER 6: Metamorphic evolution of pelitic shear zones at Fox Ridge, Mcl-oed Massif ...t28 6.I Introduction r28 6.2 Geological setting and structural framework ... t29 6.3 Petrography...... 131 6.3.l Metapeliæs 13T 6.3.2 Biotite-sillimanite rocks... t37 6.3.3 Highly aluminous assemblages... r37 6.3.4 Orthopyroxene-bearing assembalges. 138 6.3.5 Mineral chemistry ...... 139 6.4 Mineral parageneses and a P-T history for Fox Ridge t4L 6.4.1 Interpretation of mineral parageneses . L4L 6.4.2 Pressure-temperature conditions of metamorphism 142 6.4.3 Summary... r44 6.5 Conelation of metamorphism with other areas. t45 6.6 Discussion..... r46 CHAPTER 7: Some geoldynatnic and thermal considerations in granulite terrains .. .148 7 .l Introduction.... .148 7.2 The duration of granulite metamorphism .t49 7.3 Geodynamic implications of decompressional P-T paths in granulite terrams 153 7. Lower crustal channel flow parameters in granulite tenains. r57 7.4.lEvídence for channel flow in modern terrains 158 7 . .ZlJonzontal structures in the mid to lower crust ... 161 7.4.3 Importance of the rheological structure of the crust... 162 7.4.4 Cltannel flow in granulite terrains: some expectations ... 163 7.5 Summary...... 168 CHAPTER 8: Discussion t70 8. 1 Introduction...... t70 8.2 Summa¡y of the metamorphic evolution of the NPCM and sunounding regions. ...17r 8.3 Evidence for the tectonic setting of the 1000 Ma event ...t74 8.4 Thermal evolution of the NPCM. ...175 8.5 Constraints on the timing and duration of the 1000 Ma event ...t77 8.6 Geodynamic implications ... 178 REFERENCES 181 APPENDD( l: Geological obserttations from other regions visited in the northern P rinc e Charles Mountains 41.1 Mount Lanyon.... I1 41.2 MountCollins t-2 A1.3 Mt. Bechervaise / Hunt Nunataks, Athos Range. I-3 A1.4 Stinear Nunataks...... T4 41.5 Mt. McCarthy, Porthos Range r-5 41.6 Crohn Massif, Porthos Range...... I-6 A1.7 Farley Massif and Mt. Jacklyn, Athos Range I6 APPENDX II: Datafor Sm-Nd geochronology . II-1 APPENDD( fr: Mineral aralyses .. ...rn-r-22 APPENDIX IV: P-T estimates using Thermocalc rv-t-22 Figures CHAPTER 1 1.1 Regional geological setting of the northern Prince Charles Mountains 2 CHAPTER 2 2.1 Photos of the terrain in the northern Prince Charles Mountains .... T2 2.2Location of Jetty Peninsula l3 2.3 Geological map of Else Platform.... r6 2.4 Geological map of Kamenistaya Platform. 17 2.5 Lithological field relationships, Jetty Peninsula .. 24 2.6 Block diagram of lithological and structural relationships, Jetty Peninsula . 28 2.7 Structural field relationships, Jeffy Peninsula. . 32- 33 2.8 Structural map of Else Platform ... . 34 2.9 Structural interpretation of Jetty Peninsula 36 2.10 Garnet zonation profile in garnet granite, Else Platform... 42 2.11 Sm-Nd garnet-whole rock isochron, Else Platform garnet granite 43 CHAPTER 3 3.1 Reaction textures in metapelites from Jetty Peninsula .56-58 3.2 Reaction textures in orthopyroxene-bearing pelite, Jetty Peninsula .61 3.3 KFMASHT pseudosections, Jetty Peninsula ....65 3.4 KFMASHTZn pseudosection, Jetty Peninsula.. .67 3.5 P-T diagrams of evolution of Jetty Peninsula pelites .. . . .72 CHAPTER 4 4.1 Location of Trost Rocks 75 4.2 Photos of Trost Rocks 76 4.3 Photomicrographs of reaction textures, Trost Rocks ...... 79-80 4.4 Back-scattered electron images of reaction textures, Trost Rocks. ...81-82 4.5 KFMASH and KFMASHTO grids . ...86 4.6 P-T diagrams for melt-producing reactions .. .90 4.7 KFMASHT pseudosection for Trost Rocks metapelites ...93 CHAPTER 5 5.1 Location of calc-silicate bearing localities 98 5.2 Phtomicrographs of Else Platform calc-silicates...... 101-102 5.3 T-aCO2diagram, Else Platform calc-silicates...... 106 5.4 Back-scattered electron images of Ba-silicates, Else Platform. ...110 5.5 Photomicrographs of Framnes Mountains calc-silicates. . 1 15-l 16 5.6 T-aCOzdiagram, Framnes Mountains . 119 5.7 P-T diagram, Framnes Mountains .lzt 5.8 T-aCO2diagram, Framnes Mountains .122 CHAPTER 6 6.1 Location map of Fox Ridge .. 130 6.2 Photomicrographs of Fox Ridge metapelites . t34-136 6.3 Reaction textures within aluminous boudin, Fox Ridge. 138 6.4 Zoning profile across garnet, Fox Ridge 139 6.5 KFMASH P-T pseudosection, Fox Ridge 142 6.6 Summary of P-T estimates, Fox Ridge. 143 CHAPTER 7 7 .l f c-fl diagrams contoured for potential Moho and mid-crustal temperatures 150 7 .2 P-T path for Prieska Copper Mine, Namaqua-Natal Belt. 152 7 .3 f c-fl diagrams contoured for crustal strength .... 156 7.4 Velocity vectors in crust undergoing channel flow...... 152 7.5 Sketch diagram of channel flow during convergent deformation...... 167 Tables CHAPTER 2 2.1 Summary of the structural and metamorphic history of Jetty Peninsula . . . . 18 CHAPTER 3 3.1 Petrographic and textural relationships in Jetty Peninsula pelites...... 29 CHAPTER 4 4.L Table of assemblages and textural relationships, Trost Rocks pelites 78 CHAPTER 5 5.I Table of assemblages and textural relationships, Else Platform calc-silicates. . . . . 108 5.2 Table of assemblages and textural relationships, Rumdoodle calc-silicates ...... II4 CHAPTER 6 6.1 Assemblages in lithologies from Fox Ridge 132 Acknowledgements Sincere thanks go to my supervisor Mike Sandiford. It was Mike who first fired my enthusiasm for geology as an undergrad, and he's remained an inspiration ever since. Mike, you told me four years ago that I'd inevitably grow to dislike you over the course of my Ph.D. - well, for once, you couldn't have been more wrong' Special thanks also are due to Martin Hand, who was both my co-worker and good friend during the course of my thesis. Many of the important ideas of this thesis were developed during my many fruitful and enlightening conversations with Martin, and I'll always appreciate the contribution he made to my work, as well as the friendship and encouragement he's continually offered. Anyone who is enough of a zoomer to see the lithosphere in a Mars Bar is okay by me' All of rhe members of rhe 1989/90 and l990l9l ANARE Prince Charles Mountains Expeditions are thanked for their excellent company and for their contributions to the success of the two field seasons. My field partners in the first season - Kurt Stüwe, Dennis Arne and Martin Hand, are thanked for teaching a very inexperienced field geologist the ropes, and for their good company. In the second season, my fellow geos Pete Kinny, Wa¡wick Crowe and Rod Brown are thanked for all their co-operation and plenty of good times, and Bob Parker and Pete Colpo for being enthusiastic and entertaining field partners. Special thanks are due to Rob Easther, for his tireless work in ensuring the success of my second field season, and in particular for organising my day trip to Trost Rocks which resulted in Chapter 4. And thanks to the helicopter pitots for taking me to places that helicopters just aren't supposed to go. Isotopic dating was done through the generous co-operation of John Foden, David Bruce, Jo Mawby and Simon Turner. The general staff in the Department of rWayne Geology and Geophysics, and in particular Geoff Trevalyen, Mussared, Nick Spencer and Rick Barret are thanked for their valuable assistance, and Sophie Craddock is thanked for always being helpful, whilst managing to keep me informed on Departmental gossip! And to Fran Parker, for being incredibly efficient whilst still being the friendliest face in the corridor. Huw Rosser at the Centre for Electron Microscopy and Microstructural Analysis at the University of Adetaide was an absolute sensation - his assistance with the electron microprobe was invaluable, and he somehow managed to be endlessly friendly and helpful no matter how frustrating the situation might have been' Ed Grew at the University of Maine provided useful comments as well as generously sending me a number of thin sections from Jetty Peninsula. Robin Oliver also provided thin sections from the Framnes Mountains, and was always good to have around. My fellow Ph.D. students here in Adelaide deserve plenty of thanks for putting up with me and managing to keep me as close to sanity as was possible. People like Graham Moss, Martin Kennedy, Jo Arnold, Simon Turner, Kathy Stewart, Bernd Michealsen, Rob Menpes, Scotty Mildren, Jo Mawby, Bruce Schaefer, Jon Teasdale and Eike Paul all helped make being a Ph.D. student in the department a shared and even enjoyable experience. Andrew Krassay - Krass, what can I say - my number one coffee partner, dancer extraordinaire and a valuable friend. Always the first person I could turn to in a crisis - and if anyone managed to help me through the low points of the last four years it was Krass. And in the good times, there was no one I'd rather be with when 'Rhythm is a Dancer'would come on the jukebox at the Maid. No summary of the last four years of my life would be complete without acknowledging the household at 54 First Avenue (and the green couch therein). To my fellow tenants over the years, especially Krass, Lones, Greg, Matt, Scotty and Martin - thanks for the best times of my life! Plenty of other pretty sensational people deserve credit for the friendship they've given me over the past few years, among others they include Marius, Ben, Mickey, Rory and Jo, my soul-brother Barney, Paul and Lynne, Haggis, Kepler Flöttmann, Sam, Roxy, Danni, Kelli, Craig, the 81ob...... Finally, thanks to my family for their endless support. That may sound like a typical cliche at the end of any acknowledgements, but the support and generosity of my family, who have put up with me being an impoverished Uni student for nearly a decade, means the world to me. Special thanks especially to my parents for always being there for me. Mineral abbreviations used in tables and figures ab = albite alm = almandine an = anorthite arm = armenite bi = biotite cal = calcite cd, cord = cordierite cor, corund = corundum cpx = clinopyroxene czo, zo = clinozoisite gr, grs = grossul¿r gÍ - gamet hyal = hyalophane ilm = ilmenite ksp, kspar = K-feldspa¡ L = silicate melt mei = meionite mt = magneteite opx = orthopyroxene plag, pl- plagiocla.se pr = prehnite py - pyrope pyr = pyrite qz, qtz = euârtz ru = rutile scP = 3s¿P.1iL sill = sillimanite sP = sPinel titan = titanite V = vâpour phase wol = wollastonite Chapter I,Ituroduction Chapter 1 Introduction 1.1 General introduction Metamorphic rocks, and their constituent mineral assemblages and reaction textures, record significant information about the thermal, baric and fluid-interaction history of metamorphic terrains. In Precambrian metamorphic terrains, this information, coupled with structural observations and geochronological and geochemical data, often provides our only clues as to the geodynamic processes which resulted in their formation. Therefore, in investigations of the metamorphic evolution of regional granulite terrains, it is of vital importance to:- ll attempt to constrain the relative timing of assemblages used to construct pressure-temperature paths, and 2l gain an understanding of the thermal and mechanical behavior of the crust during lithospheric deformation and granulite metamorphism, in order to make realistic interpretations of the thermal, baric and structural evolution of the terrain. This thesis deals primarily with petrological observations and interpretations of mineral reaction textures in Mid Proterozoic granulite facies rocks from MacRobertson Land, and in particular the northern Prince Charles Mountains region, which forms part of the extensive East Antarctic Shield (Fig. 1.1). The northern Prince Charles Mountains are almost unique in that, until recently, there had been virtually no detailed geological investigations in the region, and they therefore represented a significant gap in our knowledge of the exposed portions of the East Antarctic Shield. The region is also particularly interesting as it provides the only exposure which extends inland across the southern margin of the extensive Mid Proterozoic granulite terrains which dominate the coastline of East Antarctica (Grew, L982; Tingey, 1982). This study was undertaken in order to provide constraints on 1 Chapter 1, Introduction Napier Complex Mawson Davi Prydz Bay ENDERB FRAMNES Y LAND MOUNTAINS Hills Amery lce Larseman n Hills TROST ROCKS U' .DZ JETTY PENINSULA 2Ei853 æ.<3 FOXRIDGE 'ög N Proterozoic Archean 200 km f the northern Prince Cha¡les Mountains, showing the main locations of t Rocks and the Athos and Porthos Ranges), the extent of the hoterozoic te and greenshist transition zones moving towa¡ds the A¡chean craton in the southem hince Charles Mount¿ins 2 Chapter I, Introduction both the pressure-temperature-time evolution of the terrain, and the broader geodynamic processes which result in such regional granulite facies terrains. The thesis has the following aims: . to describe and map the geological relationships within previously poorly documented or undescribed fragments of the granulite terrain of the northern Prince Charles Mountains; . to analyse the reaction textures within various lithologies to constrain the pressure-temperature evolution of the tenain; . to investigate whether the metamorphic evolution of the region involved a single event or was polycyclic, and to attempt to constrain the magnitude and effect of the 500 Ma event in the northern Prince charles Mountains; . to place the studied regions into the broader context of the Proterozoic terrain of East Antarctica in order to propose a tectonic model for the region; . to investigate general problems relating to the thermal and mechanical evolution of granulite terrains during orogenesis, and to relate these to observations in East Antarctica. 1.2 P-T paths in granulite terrains Our understanding of the thermal and baric evolution of orogenic belts is influenced to a large extent by the interpretation of pressure-temperature (P-T) paths of the constituent metamorphic rocks. In particular, the interpretation of P-T record as being either "clockwise" (near isothermal decompression during and following the metamorphic peak) or "anticlockwise" (near isobaric cooling or compression during cooling) has been seen as being of fundamental importance, and has led to granulite rerrains being subdivided into IBC (isobaric cooting) and ITD (isothermal decompression) granulites on this basis (e.g. Harley, 1989)' A major problem in reconstructing the P-T path of individual rocks anÜor suites of rocks is the apparently fragmented record of the pressure-temperature history 3 Chapter l, Introduction preserved by mineral reaction textures. In lower grade terrains, zoning profiles within minerals can be used to constrain a more continuous P-T evolution, but in granulite terrains, peak conditions exceed the blocking temperatures of diffusion of most cations within minerals, resulting in homogenisation of minerals and erasing of prograde zoning. Therefore any mineral zoning preserved within granulites is likely to merely reflect re-equilibration following cooling beneath the blocking temperature (Selverstone and Chamberlain, 1990). As a result, metamorphic studies of granulite terrains have a reliance on the interpretation of reaction textures. However, the construction of pressure-temperature paths is often done by essentially 'joining the dots' between overprinting assemblages without appropriate geochronological and structural constraints. As a result, 'peak' and 'retrograde' pressure-temperature estimates can often be wrongly interpreted as representing a single P-T continuum, when a more complex, perhaps polycyclic metamorphic evolution is more appropriate. Such simplifications can lead to erroneous interpretations of metamorphic textures, which ultimately result in inappropriate applications of geodynamic models. In addition, differing lithologies within the same sequence can often record differing apparent P-T histories, highlighting the potential difficulties in inferring P-T paths on the basis of a single rock. This thesis examines a number of examples from East Anørctica of such cases where appørent P-T paths are, in reality, partially or entirely misleading, and these are discussed in Chapters 3,4 and 5. 1.3 Proterozoic granulites: geodynamic considerations A general problem in the interpretation of regional low to medium pressure granulite facies tenains concerns the mechanism responsible for heating large regions of the mid-crust to temperatures of 700-1000"C. In recent years, the formation of many regional granulite facies metamorphic terrains has been popularly attributed to crustal thickening during compressional tectonics (e.g. Ellis, 1987; Harley, 1989). Whilst such a process is widely inferred to explain the evolution of many high grade metamorphic terrains (e.g. England and Thompson, 1984), it is difficult to account for 4 Chapter I, Introduction the necessary thermal perturbation during deformation to result in mid-crustal granulite metamorphism by crustal thickening alone. As a result, a number of authors have invoked thinning of the lithospheric mantle during or following crustal thickening, generally accompanied by the advection of melts through the crust (e.g. Loosveld and Etheridge, 1990; Sandiford, 1989; Sandiford and Powell, 1990)' The requirement for magmatic advection of heat suggests that such thermal perturbations may be transient, and this notion clearly has important implications for the models of the thermal evolution of such terrains. For example, many granulite terrains have been interpreted to have undergone near-isothermal decompression from the metamorphic peak, yet this would require unreasonably high erosion rates or significant extensional strain rates if it accompanied a transient thermal perturbation. Evidence for such significant strain is often absent, raising the question of the significance of the reaction textures on which such an interpretation was based. If the thermal perturbation is indeed short in comparison with the timescale of the growth and destruction of orogenic belts, then a cooling-dominated history would appear the most likely scenario for most granulite terrains. This problem is discussed in more detail in chapter 7. Regional granulite facies tenains are often dominated by stn¡ctures and fabrics which record significant strain, and which have been interpreted as being horizontal at the time of formation (Park, 1931) and these have been variously interpreted as reflecting either compressional or extensional deformation (Park, 1981; Sandiford, 19S9). However, a significant problem relating to the thermally perturbed nature of continental crust undergoing granulite metamorphism, is our uncertainty of the behavior of the lower crust during lithospheric deformation under high temperatures. Indeed, it may be inappropriate to infer tectonic models from observed structures and pressure-temperature paths when the hot and ductile mid to lower crust may be decoupled from the upper crust and mantle lithosphere during deformation, as has been suggested by a number of authors including Bird (1991) and Wernicke (1991). 5 Chapter 1, Introduction of This suggests that caution should be used when interpreting the significance paths with structures preserved in granulite terrains, as well as the P-T associated these structures. This problem is addressed further in Chapter 8. In summary, geodynamic questions to be addressed in this thesis include: 1/ Are the thermal perturbations related to granulite facies metamorphism relatively transient on the timescale of orogenesis? 2t Can granulite terrains be isothermally decompressed without evidence for significant extensional deformation? 3/ How does the thermally weakened lower crust undergoing granulite metamorphism respond to lithospheric deformation, and does it act as a relatively fluid layer decoupled from the brittle upper crust? The emphasis in this thesis will be on assessing these questions in the light of the metamorphic and structural record of the northern Prince Charles Mountains region of East Antarctica. 1.4 The northern Prince Charles Mountains: regional geological setting, previous work and relevant problems The location and geological setting of the regions studied in this thesis are shown in Fig. 1.1. The basement exposure in the northern Prince Charles Mountains (NpCM) and extending north to the Mawson Coast forms part of a large Mid proterozoic (c.1000 Ma) granulite facies metamorphic tenain that encompasses much of the coast of the Australian Antarctic sector including Kemp, MacRobertson and princess Elizabeth Lands (Sheraton and Collerson, 1983; Sheraton et al., 1984; Black the Rayner et a1., 1987). This terrain, which in Enderby and Kemp Lands is known as Complex, extends to the west beyond the high grade Archean Napier Complex of Enderby Land to the eastern margin of Dronning Maud Land, and is bounded to the east by the Archean Vestfold Hitts Block (Sheraton and Collerson, 1983). The complex extends inland to incorporate the NPCM and is bounded to the south by 6 Chapter l, Introduction Archean rocks in the southern Prince Charles Mountains (Tingey, 1982). In the grade southern part of the northern Prince Charles Mountains, the metamorphic the decreases to amphibolite facies, and continues to decrease to the south towards greenschist facies Archean terrain. Granitic rocks form an important component within the proterozoic granulite tenain. These include the Mawson charnockite, dated by u-Pb (zircon) at 954+L2 and 985+29Ma (Young and Black, 1991), and extensive undated charnockites in the NPCM (Crohn, 1959; Fitzsimons and Thost, lgg}). Much of the NPCM is also dominated by voluminous felsic orthogneisses (Fitzsimons and Thost, 1992). Early Palaeozoic granites outcrop along the western margin of the Lambert Graben and along the Prydz Bay coast, (Tingey, 1981; 1982) and locally derived anatectic granitic bodies in the Larsemann Hills on Prydz Bay have been dated a¡ J$i+27Ma (Zhao et al., 1992). Extensive resetting of Rb-Sr and K-Ar systems has been recorded at approximately 500Ma throughout the terrain (sheraton et al., 1984; Black et al., 1987), and is generally associated to these early palaeozoic granites and minor greenschist facies shear zones that deform the granulites. However, evidence exists for localised regions of high grade metamorphism associated with this early Palaeozoic event (Ren et al.,1992; Dirks et al., 1993). Immediatety to the east of the northern Prince Charles Mountains, the Mid proterozoic granulites are transected by the -80 km wide, north-south trending Lambert Graben, now occupied by the Lambert Glacier and Amery Ice Shelf. This sfiucture is Phanerozoic in age (Kurinin and Grikurov, 1982; Kent, 1991) and is bounded on either side by Proterozoic granulites. An associated parallel graben, the Beaver Lake Graben, separates the Jetty Peninsula region from the rest of the northern prince Charles mountains to the west, and is filled by Permo-Triassic rift-valley sediments belonging to the Amery Group (rffebb and Fielding, 1993). The northern Prince Charles Mountains, due to their relative inaccessibility, have until recently been one of the most poorly documented regions of Antarctica, and their relationship with the surrounding terrain has been unclea¡. They were first 7 Chapter I, Introduction and the first sighted in 1947 by US Navy aircraft during 'operation Highjump" geological investigations were by the 1955 and 1956 Australian National Antarctic Research Expeditions (ANARE) (Crohn, 1959). This first phase of Australian phase geological reconnaissance ended in 1960, and was followed by a second of geological activity in the summers of 1968/69 , 69170 and 70171 by the Australian geological Bureau of Mineral Resources (BMR). This resulted in generalised (Tingey' descriptions and maps of the region accompanied by Rb-Sr geochronology Ig72,Igg2), but these investigations were still essentially reconnaissance in nature, not with no detailed metamorphic or structural studies. Australian geologists did return to the region until the summer of 1987/88 when McKelvey and Stephenson (1990) carried out a geological reconnaissance of the region near Radok Lake, and this was followed by more extensive geological programs in the summers from lggg/gg through to I991t92. Soviet geological activity commenced in the northern prince Cha¡les Mountains in 1965, and has continued almost continuously since journals, (Kamenev et al., 1993). However, due to a lack of publication in western differences in nomenclature and an emphasis on late mafic magmatism and fault tectonics, this work has been little used in Australian geological programs which have concentrated more on metamorphic and geochronological studies of the basement gneisses. published literature on the northern Prince Charles Mountains has derived primarily from the Australian expeditions since 1987/88. Fitzsimons and Thost (lggÌ\ produced the first detailed lithological, structural and metamorphic study of the Aramis, porthos and Athos Ranges, in which they described the tenain as being path dominated by felsic orthogneiss, and they infened an isoba¡ic cooling following peak metamorphic conditions, on the basis of reaction textures in calc-silicates. This notion of a cooling-dominated history has been supported by further studies on the calc-silicates by Fitzsimons and Harley (1994) and is not inconsistent with observations by Thost and Hensen (1992) in metapelites. In spite of these studies, the 8 Chapter l,Ifiroduction significance of these apparent cooling textures, particularly within a terrain which had been assumed to be dominated by isothermal decompression (Harley and Hensen, 1990), has remained uncertain. A further complexity in the interpretation of the metamorphic and structural evolution of the Mid Proterozoic terrain of East Antarctica is the uncertainty as to whether the terrain undenvent a single high grade metamorphic event at c.1000 Ma, or two or more events in a polycyclic history between c.1000 Ma and 500 Ma. A number of authors (Ren et. al., 1992; Zhao et al., 1992; Dirks et al., 1993) have proposed a pan-African (500 Ma) granulite facies event for the Prydz Bay region, and imptied that much of the terrain may have undergone a similar polycyclic evolution. Much of the geochronological evidence for a Pan-African high grade event remains ambiguous, and it is of some imponance to evaluate further evidence for its existence. This is particularly important in the study of the pressure-temperature history of the terrain, to identify whether some reaction textures record evidence for an overprinting event rather than a continuous P-T evolution. 1.5 Thesis outline This thesis results from two seasons of fieldwork in the granulite terrain of the northern Prince Charles Mountains, in which extensive sampling was done and two geological maps prepared. Whilst most of the work is based on samples collected on these expeditions, samples collected in the Framnes Mountains by Robin Oliver on the 1965/66 ANARE expedition, and on Jetty Peninsula by Ed Grew on the 1982183 Soviet Antarctic Expedition have also been used. Mineral analyses were obtained using a JEOL 733 microprobe at the Centre for Electron Microscopy and Microstructural Analysis at the University of Adelaide. For most analyses, the energy-dispersive spectra (EDS) system was used, with an accelerating voltage of 20 keV, although wavelength-dispersive spectra (WDS) analysis was performed on scapolites and barium silicates. All hand specimens, thin sections and polished 9 Chapter l, Iüroduction sections are housed in the Department of Geology and Geophysics at the University of Adelaide under the accession number A964-. Chapter 2 contains a description of the lithological and structural relationships in the Jetty Peninsula region of the northern Prince Char1es, accompanied by a detailed geological map. These relationships are placed within the available geochronological framework, and this includes Sm-Nd geochronology of a syn-metamorphic granite which formed part of this study. Chapter 3 deals with the metamorphic evolution of the pelitic granulites from Jetty Peninsula, and this is followed in Chapter 4 by interpretations of pelitic reaction textures 80 km to the north at Trost Rocks, located on an island in the Amery Ice Shelf. Chapters 3 and 4 suggest that the metamorphic evolution of the region is more complex than a simple P-T 'loop' dominated by isobaric cooling or isothermal decompression. Further constraints on the P-T-fluid evolution of the region are provided by reaction textures in wollastonite- scapolite calc-silicates, from both the northern Prince Charles Mountains and the Framnes Mountains, and these are discussed in Chapter 5. In addition, occunences of rare barium aluminosilicates within these calc-silicates are described and interpreted. In Chapter 6, granulite facies high strain zones which cut previously metamorphosed pelites are described to further constrain the P-T history, and place the assemblages into a more well defined structural framework. Chapter 7 is a discussion of thermal and geodynamic aspects of granulite metamorphism, concentrating in particular on the concept of transience in granulites and problems in invoking isothermal decompression in granulites. In addition, the response of the lower crust to lithospheric deformation during granulite metamorphism is discussed, with an examination of the applicability of models of ductile channel flow of the lower crust, and its potential effects on the thermal, baric and structural evolution of granulites. This leads on to a discussion in Chapter I of the tectonic evolution of the MacRobenson Land region and of Proterozoic granulites in general, along with general conclusions of the studY. 10 Chapter 2, Geology of Jetty Peninsula Chapter 2 The geology and structure of the Jetty Peninsula region, northern Prince Charles Mountains 2.1 Introduction Large regions of the northern Prince Charles Mountains (NPCM) are characterised by rugged peaks, often bounded by cliffs and surrounded by wind scours and crevassed blue ice (Fig. 2.1(a)). This inhospitable tenain, accompanied by strong winds and low temperatures in comparison to those experienced on the coast, poses significant problems for geological investigations, and it is partly for this reason that detailed geological maps of the region have been lacking. However, the region around Beaver Lake, on the southwestern margin of the Amery Ice Shelf, has more accessible, undulating exposure, and its elevation close to sea level means that the climate is significantly milder, to the extent that some authors have proposed the name 'Amery Oasis' for the region (McKelvey and Stephenson, 1990). In particular, the exposure on Jetty Peninsula, a 50 km long north-south trending region of low- lying, undulating exposure, separating Beaver Lake from the southern Amery Ice Shelf (Figs. 2.1(b), 2.2) is particularly suitable for detailed observations, and therefore this region was chosen as being suitable for establishing a geological framework from which other exposures in the NPCM could be considered. This chapter deals with the geological relationships on Else Platform and Kamenistaya Platform, which comprise the northern part of Jetty Peninsula (Figs. 2.1(b), 2.2). These platforms comprise an essentially ice ftee atea,160 km2 in extent, consisting of low knolls and ridges, 10 - 20m in height, separating numerous frozen 11 Chapter 2, Geology of Jetty Peninsula .-& -4 -tt, -a- c 2.1 ßig. looking (a) Aerial view of the Porthos Range, northern Prince charles Mountains, ff;;;. S".-iit of nigher mouitains are al least 1000m above the ice' lh) Aerial view of Else Plãtform, on the northern end of Jetty Peninsula, taken ËJ*;Ïti;ä" or rãóo-, showing rhe relarively flar exposure. The¡rorrhern Amery Ice Shelf is in end 'regrouná and the the 12 x 8 km' (c) olls and ridges' Permanent sno l2 Oo ÈÐa > r^\ ozòo =tt l-¡' 1' + o Éq 0 o ÈRf \ U' A N) ,o (tl t f /,7 a oo OE EF m o Õ"\ /a) (t \ q\ E'1fl" m (/o^o.l \ €68i,ÈÉ65- _z g !R ë ãhÕÀ¡ aa -s o (¡ "õ.ggF' N c¡ ilo o \<. iãg= o) Ì'FÚOÞ x o(Däô¡ 3 çù O) ã ã ae 'r .'.'. NIEA iâ.(Dä=' È x ;4, 3S -'. â å Lr Ef J .'.', =' q ød:¡ N 8H()Þ< z Ê p Þ t I6 F¡u o 9¡ É ô \ ãs=s c Ën U'o :6(DË' z Ðz ò TÞô-nt -{ = ú 0.' c ¡i a [ É hR,' o (t) O.é(D @ = 8:. oo Fã7 m 9 Fä o'¡', EË= D. o<= ,5.- xqu. o an ôrb ) qäß v, o Ëatl s 3 --UlN o s-' \ ÞHh E (\ =Þtai \¡ 6'86 I g-5 N 5f¡ Ð Erþ !P. e a :t c; qa $ (\ (- o o á'3 E ! o m f s o Oa (Þ " -{ ä ig5< { 6 \ J ê (l øo l- I o .o ? D I 5 o õ9 T' d to (\ ä m g U' ---\\ \-..- c v =o z =ø D (D z = (\! ERo @ : õ' c 1- s-- (, =tt, o F Chapter 2, Geology of Jetty Peninsula and open lakes. The outcrop is generally frost shattered, although usually in place (Fig. 2.3), and consists of granulite facies metamorphic rocks of probable Mid Proterozoic age (Tingey, 1982; Manton et al., 1992), with numerous syn- and post- tectonic intrusions and a small area of Permo-Triassic sediments on Kamenistaya Platform. The granulites of Jetty Peninsula are separated from the rest of the NPCM granulite tenain by the Permo-Triassic sediments of the N-S trending Beaver Lake Graben, and are bounded to the east by the more extensive Lambert Graben, which is currently occupied by the Lambert Glacier / Amery Ice Shelf system. Previous geotogical descriptions have shown that the rocks on Jetty Peninsula consist of garnet granulites intercalated with minor sillimanite beæing, mafic and calc-silicate gneisses, intruded by garnet bearing granites and later by mafic and alkalic-ultramafic dykes (Grew 1985, Manton et aI,1992). Zfucon ages suggest that granulite facies metamorphism may have occurred around 1000Ma on Jetty Peninsula and later granite intrusions occurred at approximately 940Ma and 500Ma, (Manton et aI, 1992). Although generalised metamorphic textural relationships have been described by Grew (1986), detailed descriptions of the metamorphic textures and structural relationships are lacking, as are correlations with surrounding fragments of the terrain. Soviet geological investigations in the Jetty Peninsula region have been canied out from the summer base of Soyuz on southern Kamenistaya Platform, although these investigations have concentrated primarily on the Palaeozoic and Mesozoic mafic magmatism and fault tectonics (Andronikov, 1990, 1992; Hofmann, 1991, Mikhalsky et aI, 1992), and have contributed little to our understanding of the metamorphic and structural history of the region. This chapter deals with the lithological and structural relationships on Jetty Peninsula, which provide a basis for a geological history of the region as well as correlations and comparisons between the NPCM and outcrops along the Prydz Bay coastline. It is based on mapping carried out on a I: 20 000 scale over two field t4 Chapter 2, Geology of Jetty Peninsula seasons, with Else Platform being mapped in conjunction with M. Hand (University of Melbourne), D.Arne (Geotrack International) and K. Stüwe (University of Adelaide), whilst Kamenistaya Platform was mapped by the author alone. 2.2 Regional geological context Jetty Peninsula forms part of a large Mid Proterozoic (c.1000 Ma) granulite facies metamorphic terrain that encompasses much of the East Antarctic coast from Enderby Land to the Vestfold Hills, and which was described in detail in section 1.4. The southward transition to amphibotite facies occurs around 80 km to the south of Jerty Peninsula, in the Mt. Meredith region (Tingey, 1982). Jetty Peninsula is the easternmost outcrop in the NPCM, occurring as a fault-bounded basement exposure between the north-south trending, Palaeozoic, Lambert and Beaver Lake grabens (Fig. 2.2). Therefore, the Mid Proterozoic granulite facies metamorphic terrain has been significantly overprinted by structural and magmatic activity related to this graben development. Geological relationships elsewhere in the NPCM have been documenred briefly by McKelvey and Stephenson (1990), and in more detail by Fitzsimons and Thost (1992). Fitzsimons and Thost (1992) describe an orthogneiss- dominated, medium-pressure granulite terrain in the Porthos Range, with syn-tectonic charnockite and leucogneiss intrusions, and east-trending high-grade structures. These are overprinted by pegmatites, low-grade shear zones and late mafic dykes. 2.3 Lithologies The geological maps of Else Platform (Fig. 2.3) and Kamenistaya Platform (Fig.2.Ð show rhe distribution of tithologies within the granulites of Jetty Peninsula, whilst Figure 2.6 schematically summa¡ises lithological relationships. The lithologies of the region can be broadly subdivided into three groupings. The first is a dominantly supracrustal succession of layered granulite facies gneisses. The second are intrusive, dominantly granitic lithologies, which have intruded at various stages during and following the structural and metamorphic evolution of the region, whilst 15 2, Geology of Jetry Peninsula 012 ô N Scale:km ð 67' RJ = n = 0 CISTOJE LAKE t LAKE LEGEND Garnet granite ì^ I c"rnet-b""r¡ng felsic dykes 1 l]]j Megacrystic granite .L : :r Leucogran¡te L l Layered gneiss lE r"l¡,¡" sneiss (P) Orthopyroxene-biotite gne¡ss Homogeneous quartzofeldspath¡c gnetss Pegmat¡te 0 quartz veins ! .r Alkaline maf¡c dykes ( 1) \á'' F M¡gmatitic aranitic dikes (2) 0 \l Calc-silicate grìeiss \9. [ilafic gneiss + Ultrmaf¡c gneiss \Fault â, Permânenl snowfield ..-- À¡oraine ["ig.2.3 Lithological map of Else Platform. Structural relationships are shown in Fig. 2.8. t6 Chopter 2, Geology of Jetty Peninsula 69o00'E ,/ KAMENISTAYA PLATFORM Jetty Peninsula zdso's GEOLOGICAL MAP sl km 0 1 2 N= 42 = 20 sqm6 N 67o /j Beaver Lake Lambert Glacier LITHOLOGICAL UNITS tlecsnt alluv¡um / mora¡ng U I ôç¡ttt P€rmian sandsonsS ./tt I R-T'l and assoc¡ated granitos I IXX Fsls¡c dyk€s I I I Porphyritic granits ffiæ lr:,: ,: ,l Biotite-opx (chamækitic) gneiss l>1 r"' l Gamet-b¡otits gneiss w¡th feldspar garnet ffi l€ucognsiss and rare sill-rich layers Unditfsrentiated pyroxenile, calc-sil¡cat€, bt-opx l...... ¡ and bt-gt gn€isses, w¡th fino grainêd lelsic dykes .D Alkalino malic Pods Malic dykss Soyuz Base Ouarlz ve¡ning (ussw 'Migmal¡tic' biotit€ gran¡te Dip of primary foliation (51) Trend of gneissic layering (Sq-51) Obsry€d Fault lnfsrrsd tan Scrimgeour, APril 1991 orientation ßig.2.4 Geological map ol Kamenistaya Platform. stereonet indicates poles to the S1 gneissic foliation' t7 Chapter 2, Geology of Jetry Peninsula the third is late Paleozoic to Mesozoic sedimentation, mafic magmatism and veining associated with the Lambert Graben development. The abundance 0f metasedimentary rocks and absence of voluminous homogeneous felsic gneisses distinguishes Jetty Peninsula from most other Proterozoic basement outcrops in the NPCM. 2.3.lLayered Gneiss The layered gneiss is the dominant lithology on Jetty Peninsula, and consists predominantly of heterogeneous quartzofeldspathic to semi-pelitic garnet-bearing gneisses. These are interlayered with sillimanite-bea¡ing pelites, orthopyroxene- bearing gneisses and calc-silicates. The lithological layering is defined by units between 1 and 100 metres in width, and is continuous along strike on a scale of hundreds of metres. Where the lithological layering is well developed it is defined by slight variations in the arnount of garnet and biotite, with contacts between adjacent layers being generally gradational. The layered, heterogeneous nature of this sequence together with the existence of pelitic and calc-silicate layers strongly suggests that the gneisses are of a metasedimentary origin. In general, the layered supracrustal gneisses are strongly migmatised with a distinctive banded appearance. Layer parallel, garnet and K-feldspar bearing leucosomes are folded by the gneissic fabric which is paralleled by garnet absent leucosomes in the fold hinges. A later biotite foliation overprinting the garnet bearing assemblages can sometimes be distinguished in the vicinity of hinges of tight to isoclinal folds F2 defined by the gneissic fabric. Complete overprinting of the gneissic fabric in the layered gneisses occurs along high grade east-west trending shear zones. Overprinting the last penetrative high temperature foliations in the layered gneiss are garnets, and less commonly orthopyroxene, which occur within small K-feldspar segregations, reminiscent of those described from the Larsemann Hills by Stüwe and Powell, (1989a). 18 Chapter 2, Geology oÍ Jetty Peninsula Metapelitic gneiss Metapelitic gneisses occur as narrow horizons up to l0 m in width which can be traced for up to 100 m within the layered gneiss. They are relatively uncommon on the southern part of the Else Platform but become progressively more abundant to the north, and are also common on central Kamenistaya Platform. The pelitic gneisses are relatively coarse grained and usually have a banded appearance caused by alternating garnet bearing leucosomes and leucosome poor regions. Leucosomes sometimes define rootless isoclinal folds within the gneissic foliation and are truncated by later leucosomes along the axial surfaces of folds. Less commonly, metapelitic lenses are essentially leucosome free and crop out as dark resistant features. The often complex petrographic relationships within the metapelitic gneisses are described in detail in Chapter 3. The peak assemblage is generally dominated by garnet and sillimanite, with or without spinel, with a later 52 biotite fabric, and cordierite * spinel coronas between garnet and sillimanite. An orthopyroxene-garnet- cordierite-biotite gneiss also occur in a lens approximately one metre in diameter on northeastern Else Platform (Grew, pers. comm., 1991). Darkfelsic gneiss A dark-coloured medium to coarse grained biotite-quartz-feldspartgarnet felsic gneiss is found on northern and central Else Platform and on southern Kamenistaya Platform. It is characterised by dark quartz and green to yellow feldspars, with a poorly defined biotite fabric and variably abundant to absent rounded garnets. This unit is distinctive in that it contains variably abundant lenses of pyroxenite and calc-silicate. The association of these da¡k gneisses with calc-silicate 'lenses suggests a metasedimentary origin for this unit, with the associated pyroxenite possibly being of a metavolcanic origin. 19 Chapter 2, Geology of Jetry Peninsula Calc- silicate gneiss es Iænses and discontinuous layers of calc-silicate gneiss are enclosed within the layered gneiss and are conspicuously concentrated within the dark felsic gneiss. These lenses are usually narrow (1-5 metres) in width and are up to 40 metres long. The pronounced banding within the calc-silicates is defined by differing proportions of diopside, wollastonite, scapolite, calcite and titanite, with only rare marble units. In some lenses, individual layers are characterised by high variance assemblages and are sometimes essentially monomineralic. A grain shape fabric defined elongate diopside and wollastonite parallels the compositional layering. Grossular occurs as narrow rims between wollastonite, scapolite and calcite and postdates the granular fabric, whilst scapolite and wollastonite show breakdown reactions to calcite, anorthite and quartz. A feature of the calc-silicates is the anomalously high proportion of barium silicates, including hyalophane and armenite. More detailed descriptions of the petrography and significance of the calc-silicate gneiss are given in Chapter 5. Mafic gneisses Pyroxene-rich mafic gneisses outcrop as small resistant lenses in the layered gneiss and also form several prominent dark units, concordant with the gneissic layering, striking west and northwest in the central and northern parts of Else Platform and southern Kamenistaya Platform. They contain almost no compositional layering or structural fabrics, and a¡e always less foliated than the surrounding rocks. A characteristic feature of the mafic and ultramafic gneisses is the absence of garnet. A late foliation defined by the orientation of hornblende now seen in a number of the mafic gneisses appears to have formed during late D2 shearing and overprints an earlier orthopyroxene-clinopyroxene assemblage. Coronas of biotite and blue-green amphibole form around the ea¡lier pyroxenes and hornblende. 20 Chapter 2, Geology of Jetry Peninsula O r tho py r o xen e b e ar in g qu ar aofe Id sp athi c g n e i s s Layers of orthopyroxene- and biotite-bearing felsic gneiss occur commonly throughout the layered gneiss. Metre scale lithological heterogeneity is defined by varying proportions of orthopyroxene, biotite, plagioclase and green K-feldspar. Elongate orthopyroxene defines a grain shape fabric with plagioclase and quartz. Biotite partly replaces the orthopyroxene and defines a slightly to strongly discordant foliation. Superficially the orthopyroxene-biotite bearing lithology resembles the non-megacrystic charnockite on Kamenistaya Platform (see below), but its relatively heterogeneous composition, concordant nature and gradational contacts suggest that it may, more likely, be of metasedimentary origin. H o mo g ene ous quartzofe lds p athic gne i s s A unit of homogeneous felsic gneiss approximatety 1 km thick outcrops on either side of an east-trending fold closure in the central part of Else Platform (Fig. 2.3). Mineralogically it is similar to the garnet-biotite-Kfeldspar-quartz variant of the layered gneiss although garnet generally predominates over biotite, and it is much more compositionally uniform. The homogeneous quartzofeldspathic gneiss has a strong fabric defined by elongate garnet with biotite defining a pronounced foliation. This unit is more migmatised than similar rock types in the layered gneiss with abundant garnet-bearing leucosomes transposed into the main fabric and truncated by other, generally garnet absent leucosomes. The contact between the homogeneous quartzofeldspathic gneiss and the surrounding gneiss is difficult to assess due to the extensive frost shattering of the outcrop, however, where observed intact, the contact appears to be gradational over a distance of approúmately 5 metres. 2,3.2 lntrusive Rocks The structural and metamorphic history of the Jetty Peninsula is punctuated with numerous phases of igneous activity, particularly syn-tectonic granites and charnockites. 2l Chapter 2, Geology of Jetry Peninsula Leucogranites Homogeneous leucogranite gneisses containing deformed xenoliths occur within the layered felsic gneisses of the supracrustal sequence. They are often difficult to distinguish from the semi-pelitic layered gneisses which they intrude, but they are characterised by their homogeneous nature and by the presence of elongate deformed xenoliths (Fig.2.7(a)). Their mineralogy is dominantly feldspar and quartz, with abundant small garnets and ra¡e biotite. Due to the diffuse nature of the contacts between the leucogranites and surrounding semi-pelitic gneisses, and the similarity between the two lithologies, the two units proved difficult to distinguish, and hence no attempt was made to map the precise boundaries of the leucogranite bodies. However, the leucogranites are most common on the southernmost regions of Else Platform, south of the megacrystic granite. The presence of xenoliths containing a previous foliation discordant with the 51 foliation in the leucogranites supports the suggestion that they intruded synchronous with D1, possibly as a locally derived partial melt. Ultramafic pods Two isolated pods or boudins of very coarse-grained metamorphosed ultra- mafic pyroxenite outcrop on southeastern Else Platform, and preserve a well defined 52 foliation. The pods are lenticular in shape, up to 30 metres in length and consist primarily of elongated pyroxenes (orthopyroxene and titaniferous augite) up to 40-50 mm in length, with less abundant olivine and hercynitic spinel, and with accessory magnetite and plagioclase. Augite contains abundant exsolution of rutile, and the pyroxenes are often mantled by retrograde hornblende, biotite and quartz. The ultramafic pods are almost certainly of an intrusive origin, and intruded before or early during the deformational history. 22 Chapter 2, Geotogy of Jetry peninsula Megacrystic granite Large bodies of deformed porphyritic granite occur on both Else and Kamenistaya Platform, and although being compositionally variable, all appear to part form of the same suite of intrusions. The largest body is a garnet-biotite porphyritic granite on southern Else Platform, which contains large, elongate megacrysts of white K-feldspar in a matrix which is rich in both biotite and garnet, and it is generally cut by abundant felsic dykes. Porphyritic granites are also found on northern Kamenistaya Platform, although they are usually darker in appearance and contain little or no garnet. The degree of deformation in these granites is quite variable, often within individual bodies, with the large feldspar phenocrysts being subhedral and rectangula¡ with preserved simple twins in low strain regions, whilst in high strain regions they form rounded augens (Fig. 2.5(a,b)). xenoliths of the surrounding layered gneiss are locally present, and often display gradational contacts with the enclosing granite. Fabrics in both the granite and the xenoliths are generally parallel' on a regional scale, the contact between the layered gneisses and the megacrystic granite is broadly discordant to the lithological banding in the layered gneiss, while on a more local scale the contact is often concordant with lithological layering. Charnockite Dark, medium grained charnockite is a the most common lithology on northern Kamenistaya Platform, and contains a mineralogy of quartz, orthopyroxene, yellow-brown potassium feldspars and occasional, usually secondary, biotite, rvith rare occurrences of garnet (Fig. 2.5(a)). In some localities it contains late undeformed felsic segregations containing large euhedral garnets. It is generally spatially æsociated with dark porphyritic granite and augen gneisses, and appears to be of an early igneous origin. The relatively fine grained and non-porphyritic nature of this lithology distinguishes it from the Mt. seaton and Mt. collins charnockites elsewhere 23 Clnpter 2, Geology of Jetty Peninsuln d f 7 :-r !-+ E+ -.-G-. I -+ --- .*qh *Ê 'ü trf,' ':È-\\ Fig. 2.5 Field relationships, Jetty Peninsula. (a) Contact between deformed porphyritic granite and frne grained charnockite, northern Kamenistaya Platform. (b) Less deformed porphyritic granite, with more euhedral feldspar megacrysts, cut by a late pseudotachylite, northern Kamenistaya Platform. (c) Within a fine grained north-south trending biotite granite dyke on Else Platform, showing garnet-bearing leucosomes giving a distinctive migmatitic appearance. (d) Aerial photo, taken at an altitude of c.200 metres above Else Platform, showing swarms of felsic dykes, and a distinctive black mafic dyke. 24 Chapter 2, Geology of Jetty Peninsula in the nofthern Prince Charles Mountains, which are more analogous with the Mawson Charnockite (Young and Ellis, 1991). Felsic d)'kes Swarms of felsic dykes are extremely widespread on Jetty Peninsula, and generally parallel the dominant gneissic layering (Fig. 2.5(d)). They are typically several metres in width, but occasionally form larger bodies up to hundreds of metres in extent. Typically the dykes form discrete bodies that are slightly discordant to the lithological layering in the layered gneiss. Locally however the dykes occur as swarms that occupy upwards of. 50Vo of the outcrop over several square kilometres. In these regions, the original layered gneiss occurs as large xenoliths and enclaves within the felsic dyke material. In addition to these felsic dykes there is also a large E- rW trending body of a leucogranite on northern Kamenistaya Platform which is coarser grained than the felsic dykes but has a very similar mineralogy and is assumed to be related to the later stages of felsic dyke intrusion. This leucogranite hæ been dated by Manton et aI (1992) at 940 + 20 Ma by U-Pb zircon, and at 718 + 32Maby a Rb-Sr mineral isochron. No attempt has been made on Figure 2.3 to identify individual felsic dykes less than approximatety 50 m in width, but where felsic dyke swarms locally comprise greater than 507o of the outcrop area, the area has been mapped as garnet bearing felsic intrusive. The felsic dykes are typified by fine grained quartz and feldspar with deformed quartz phenocrysts up to 15mm long, and almost ubiquitous accessory garnet. The latter often gives the felsic dykes a characteristic spotted appearance. The dykes usually contain a strong foliation defined by elongate quartz and often have a pronounced elongation lineation defined by quarø and garnet. Occasionally, xenoliths of layered gneiss containing a foliation correlated with 51 can be seen in the felsic dykes. The foliation and occasional fold axial surfaces in the xenoliths are often strongly discordant to the quartz fabric in the dykes and are transposed into the quartz 25 Chapter 2, Geology of Jetty Peninsula fabric at the margins of the xenoliths indicating that the garnet bearing felsic dykes intruded the layered gneiss afær D1. Garnet Granite A large body of homogeneous, light grey, medium grained, garnet biotite granite occupies the core of the major F2 regional fold on central Else Platform. It is not cut by the felsic dykes, and its relationship with the gneissic layering is regionally concordant, but locally sharp and discordant. It has a weak to moderate 52 biotite foliation with a strong quartz stretching fabric, and appears to have intruded either as a sill prior to F2 folding or, more probably, during the development of F2. The contact between the garnet granite and the enclosing lithologies is sharp and locally highly irregular (Fig.2.7(d)). A feature of the garnet granite is the spectacular array of associated granitic dykes. These are up to 50m wide and radiate away from the main body (Fi1.2.7(d)). Although these dykes truncate felsic dykes that are folded around the major F2 fold structure they contain the same foliation found in the felsic dykes, albeit less pronounced. This suggests that the garnet granite intruded sometime towa¡d the end of.D2. The results of Sm-Nd isotopic dating of the garnet granite are given in section 2.5. Biotite graníte dykes Northwest to northeast trending biotite granite dykes, around 5-10 metres in width, are found on both Else and Kamenistaya Platforms, and are characterised by containing coarse leucocratic bands and pegmatites which give the dykes a distinctive migmatitic appearance (Fig. 2.5(c)). The dykes crosscut the gneissic foliation but have been subjected to localised folding and shearing and contain at least three generations of leucocratic bands. The earliest generation of leucosomes are tightly folded (Fi9.2.7(Ð), a second generation are axial planar to these folds, and then a third generation appear to postdate deformation. There is a weak foliation within the dykes, defined by leucosomes and biotite, which strikes around 020, dipping 60-700 26 Chapter 2, Geology of Jetty Peninsula to the east, and which is at a slight angle to the dyke margins. The intrusion of these dykes postdates the early (D2) shear zones and the regional E-W folds, but one is offset 50 metres by a late D4 sinistral shear zone on central EIse Platform. These granites have been dated by Manton et al (1992) at -500 Ma, using U-pb on zircon and monazite and Rb-Sr mineral-whole rock isochrons. The sense of transposition of the gneissic layering surrounding the dykes is consistently sinistral in the horizontal plane suggesting that the dykes have intruded along active shear zones. The horizontal offsets on these shear zones are small, usually being less than 30 metres. The dyke foliation is axial surface to abundant isoclinal folds defined by leucosomal veins. Although deformation within the dykes appears to be associated with shearing there is no consistent asymmetry to these folds. Manton et aI. (1992) have shown that the peraluminous geochemical signature of the biotite graniæs is characteristic of high temperature melting of relatively anhydrous granulite facies metasediments. Pegmatites Massive feldspar-quartz-biotite-garnet pegmatites are abundant on Else Platform and found less commonly on Kamenistaya Platform, and range from outcrop scale veins to north-trending bodies up to 2 km in length and 750 m in width. Only the largest pegmatite intrusions, or zones of intense pegmatiæ vein emplacement are shown in Figure 2.3. The pegmatites are cut by late greenschist facies shear zones but a¡e otherwise undeformed. The general NNE-SSW trend of the major pegmatite bodies suggests they are possibly associated with the F3 folding and the NNE-SSW trending foliation and leucosomes within the migmatitic granite. These larger bodies a¡e similarly oriented to the sinuous migmatised dykes and may be related to them in some way. Rare onhopyroxene-bearing and magnetite-bearing pegmatites also occur, and Russian geologists have also reported arsenopyrite and molybdenite bearing pegmatites from Else Platform, and beryl in pegmatites from Kamenistaya Platform (Manton et al., 1992). The high As, Be, B and Mo contents of the pegmatites have led Manton et aI. (1992) to suggest that the source of the pegmatites is different to 27 Chapter 2; Geology of Jetry Peninsula + + + +'+ + + + + +/+ + + + 1'+ + + + +'+ MBG + + +'+ LS 1 F2 U K + + + + + + + + + + '+ r + + + '+ + Is2 + + + + '+ x + .+ + + + l+ + + + + +'+ '+'+ + + h + + + + +'+ '+'+ + + ih + {{ + + +l+ '+'+ + + + + + {.1 + + +l+ '+'+ + /,. + + + '+ + il + + +l+ '+'l + Mctf + J + '+ + + + + +l+ i+i1 + + + + '+ + + + + +'+ ++ + + I + ,+ + t'+ '+'+ + + + + + '+ + GG +'+ '+'+ + + + + + + '+ ++ + r{'+ '+'+ ++++ + ï P CHK + + + ,+ +i+ i + rl'+ '+'+ +*++ + # + + + ,+'+ rJ'+ + + J'+ '+'+ +I++ + + + + + rJ'+ + + rJ'+ '+'+ +T++ + h + D + + + ,+ rJ'+ + '+'+'+ ++++ + + + + + + + + ,+ +'+ + '+'+'+ ++++ + + + + + + + + +'+ ï + '+'+'+ ++++ + + + \ D2 FD NORTH 7 sz PEG Fig.2.6 Schematic block diagram illustrating generalised lithological and structural relationships observed in the Jetty Peninsula region. No scale is implied. Abbreviations: LS layered supracrustals; UMP ulEamafic pod; CHK chamockite; MCG megacrystic granite; GG garnet granite; FD felsic dyke; MBG migmatitic biotite granite; PEG pegmatite;MD Carboniferous mafic dyke; BLG Beaver Lake Graben conüaining Pennian sandstones; AMP Mesozoic alkaline mafic pipe; QV quartz veining. 28 Chapter 2, Geology of Jetry Peninsula that of the biotite granites, and that they were derived from a crustal source previously unaffected by granulite facies metamorphism, as Sandiford (1985) also proposed for late Be-bearing pegmatites in Enderby land. 233Late Palaeozoic to Mesozoic dykes and sediments A feature of the Jetty Peninsula region which is unique in the East Antarctic Shield is the extensive evidence for Carboniferous to Cretaceous magmatism, rifting and sedimentation, which is associated with the Lambert Graben development and the separation of India and Antarctica. Mafic dykes Prominent mafic dykes occur widely throughout the Jetty Peninsula region, generally trending in a N-S direction and dipping steeply to the west. These are far more abundant on Jetty Peninsula than elsewhere in the NPCM and show a strong spatial association with late brittle faults. Most commonly they are fine grained basaltic dykes, approximately 5-10m in thickness, with occasional small amygdales of calcite, and angular xenoliths of granitic country rock. Small phenocrysts of pyroxene and olivine are also present near the centre of some of the wider dykes. There is usually some displacement of the country rock along the N-S trending dykes, which crosscut and sinistrally offset less common E-W trending mafic dykes. The most abundant generation of mafic dykes do not cut the Permian cover sequence immediately to the south and west of Jetty Peninsula and have been assigned a K-Ar age of 308+10 Ma by Hofmann (1991). P ermo -Trias s i c s andstone s Permo-Triassic sediments belonging to the Amery Group (Mond, 1972; Webb and Fielding, 1993), outcrop on the northwestern margin of Kamenistaya Platfonn, on the edge of Beaver Lake near Soyuz and extending south towards Flagstone Bench. In the region of Soyuz these sediments are coarse, well sorted sandstones, probably 29 Chapter2, Geology of Jetty Peninsula belonging to the Bainmedart Coal Measures, which dip at approximately 10o to the east. Small fragments of what appeared to be petrified wood were also observed. The contact between these sediments and the Precambrian basement appears to represent a N-S trending fault, although the exact nature of this contact could not be observed due to the highly shattered outcrop in the region. Alkaline mnfrc dykes and pipes On Kamenistaya Platform there are at least four occurrences of fine grained alkalic-ultramafic pipe-like bodies containing nodules of spinel-garnet lherzolite, which have been described by Soviet geologists (Andronikov, 1990). One of these pipes intrudes the Permian sediments to the south of Soyuz, showing that this is younger than the abundant N-S t¡ending basaltic dykes which intruded prior to Permian sedimentation. These alkaline intn¡sions have been dated using K-Ar at 145- 150 Ma by Andronikov (1990) and 130 Ma (early Cretaceous) by Hofmann (1991). Similar generations of alkalic and mafic intrusive dykes have been described elsewhere in East Antarctica, particularly in the vicinity of the Lambert Graben system, and range in age from Cambrian to Eocene (Sheraton, 1983). Quartzveins On western Else Platform and particularly on Kamenistaya Platform, there are abundant quartz veins, which are usually about a metre wide but which can form bodies up to lkm long and 100 m wide. Aside from occasional coa¡se K-feldspar and minor, secondary calcite, the quartz veins are monomineralic and contain vuggy q\^rtz and banded agate, with occasional larger quartz crystals up to 20cm in length. On Kamenistaya Platform a large area of massive quartz intrudes both the Proterozoic and Permian rocks, indicating that the vein formation is post-Permian. In general, the veins seem to exploit the N-S trending faults which are assumed to be related to the development of the Lambert Graben, and in places on Else Platform they are spatially 30 Chapter 2, Geology ol Jetty Peninsula related to zones of epidote and chlorite alteration. Mafic xenoliths within the quartz veins display minor malachite staining and chlorite alteration. 2.4 Structure A relatively simple structural history is preserved on Jetty Peninsula, and the important structural features are shown in Figures 2.8 and 2.9. Ductile deformation of the metasedimentary and the early intrusive rocks on Jetty Peninsula occurred during two major events, D1 and D2, associated with high grade metamorphism. D1 structures are locally complex and may represent the superposition of multiple events, whose individual features have not been resolved. However, in general, D1 involved isoclinal folding, resulting in an 51 gneissic foliation which is usually parallel to lithological layering. The major recognisable fold structures were formed during D2, which also was responsible for the present structural trends. The garnet bearing felsic dykes are useful for distinguishing between some D1 and D2 structures, since the dykes intruded after D1 but are folded by D1 D3 is manifest as large scale gentle warps and D4 produced minor ductile shear zones that are only important locally. These ductile structures have been overprinted by more brittle structures which are almost certainly associated with the development of the extensive Lambert Graben system. 2.4.1D1 The earliest deformational features recognised on Jetty Peninsula are isoclinally folded garnet bearing felsic segregations (Fig 2.7(b)). An earlier isoclinally folded gneissic foliation has also been recognised in the Radok Lake area (McKelvey and Stephenson, 1990), the Aramis Range (Fitzsimons and Thost,1992), south of the Nemesis Glacier region (Nichols, 1992), and in the Athos Range and Stinear Nunataks (this study). Additional evidence for the existence of an earlier fabric is the presence in early intrusive leucogneisses of xenoliths which contain a foliation which is discordant to the regional 51 gneissic fabric (Fig.2.7(a)). It was not 3l Chnpter 2, Geology of Jetty Peninsula I ßig.2.7 Structural field relationships, Jetty Peninsula. (aiPreviously deformed xenolith within early leucogneiss, Else Platform. The Sl foliation is parallel to the hammer handle. (b) Rootless isoclinal F1 fold of an early felsic segregation within leucogneiss, Kamenistaya Platform. (c) Rare uprigtrt parasitic F2 fold, northeastern Else Platform' 32 Clnpter 2, Geology of Jetty Peninsula d g \.. -.Þ :t--:- - -t ¡-È- -ùra-,. ßig.2.7 (continued) Structural field relationships, Jetty (d) Aerial view looking east (altitude c.1800m) of the intrusion of garnet granite into the core of the major F2 fold, central Else Platform. Whilst the granite contains a¡r Sz foliation, it is locally discordant (arrow A) and the associated radiating dykes crosscut folded felsic dykes (arrow B). A mafic dyke approximately defines the axial trace of the Fz fold. (e) F¡ fold in felsic gneiss with weak axial planar foliation (parallel to hammer handle), southeastern Else Platform. (fl¡ Tight F3a folds of leucosomes within 500 Ma biotite granite, Else Platform. igl Aðrial view looking easr (altitude c.l500m) along major D2 shear zone (Soviet Shear Zone), ióutheastern Else Platform, with intensification of the shea¡ zone (and resultant attenuation of lithological units towa¡ds the bottom of the picture. Width of freld of view is around 2 km. 33 Chapter 2, Geology of Jetty Peninsula sl 84 L1 Ft s2 45 t2 LEGEND + F2 told a¡oa lmportant litholog¡cal boundarY Axial trace ol matot F2 Lineamenls (lrom eorl.l photosr.ph.) (u3u¡lly ,,) Gne¡ss¡c foliation sl) s2 '':t' uj, D2 shear zone a0 ,,2 O3,D4 shear zone lollellon/llñ.ållon 4l/ /<- s1tLl¿¿, S2lL2 ¿às2(sz'tlL2 ax¡al y'ss strt+ N ,r( ,root surrac€ f r' Small scale F2 fold axis ) Mâior F2 fold axis OuÞ Mo"ement sense on D2 shear zong O down l/t t^t. br¡ttte fautt w ,t' ., R"gion of little o¡ no fol¡ation t--Tl Moraine/snowf ¡eld /lake Sl gE reglon ol plBtlorm IA lm Fig. 2.8 Map of Else Platform, showing structural relationships 34 Chapter 2, Geology of Jetry Peninsula possible to distinguish whether these early features are the products of a separate deformational event or a paft of a continuum culminating in the development of the major gneissic foliation (S1). Previous workers (e.g. Fitzsimons and Thost, 1992), have ascribed isoclinally folded leucosomal segregations to a previous deformation, but until it can be established otherwise they will be regarded here as early D1 structures. The D1 deformation on Jetty Peninsula produced the regional gneissic banding, (Sr) and resulted in transposition of early lithological contacts into 51. The fact that infened peak assemblages define 51 indicates that D1 was contemporaneous with the peak metamorphic event. 51 is almost always parallel to the large scale lithological banding and major F1 folds are infened to be isoclinal. No large scale F1 folds have been positively identified, but F1 folds are seen rarely as small scale, generally rootless, isoclinal folds, generally defined by earlier leucosomes. D1 structures have been strongly reoriented during D2, although on Kamenistaya Platform, where D2 is less intense, the orientation of 51 suggests that it may have originally trended approximately N-S, dipping moderately to the east. In most of the layered gneisses, 51 is defined variably by biotite or by leucosomal segregations. In the semi-pelitic gneisses and orthogneisses, biotite defines the 51 fabric, along with elongate granoblastic quartz and feldspar. Pelitic gneisses a¡e similar, except that sillimanite rather than biotite is the mineral fabric element. 51 is less well defined in the mafic gneisses, although pyroxenes define an L1 lineation, and occasional felsic segregations define 51. L1 is an often coarse lineation defined by quanz rodding and mineral aggregates. 2.4.2D2 The second deformational event (D2) on Jetty Peninsula resulted in the development of the dominant east-west trending tight F2 folds and also produced 35 Chapter 2, Geotogy of Jeny Peninsula STRUCTURAL INTERPRETATION OF THE JETTY PENINSULA REGION / F / 2 -,:::P.lätfofrn / D2SZ '.'.':-':¡'.ll'.' '-2t1".--'.'l' " ? LEGEND Permian F2 Proterozoic F2 x D2 shear zones !3ß) D4 shear zones r+Ðo <\J -t Antiformal xo!¡ -1.'' q-e O=4 ìg) Synformal 'r' ¿.4 9) Fault 0 2km Soyuz rig.2,9 Simplified stnùctural interpretation of Jetty Peninsula 36 Chapter 2, Geology of Jetry Peninsula large scale reverse sense shea¡ zones. This event postdates the intrusion of the garnet bearing felsic dykes that intrude Dl structures and was synchronous with or postdated the intrusion of the large homogeneous garnet bearing granite in the core of the major F2 structure on Else Platform (Fig. 2.7(d)). Thus, although 52 is generally similar in orientation to 51, the presence of these post D1 intrusions allows a clear distinction to be made between 51 and 52. In the felsic dykes, 52 is defined by elongate K-feldspar and quartz, whilst in the garnet granite it is defined by biotite, with a steeply plunging L2 lineation defined by elongate quartz. F2 refolds the 51 layering into large upright east plunging folds, which are tight to isoclinal on Else Platform, and more open on Kamenistaya Platform. The hinge regions are often broad and a¡e characterised by apparently low D2 strains and minor 52 development, whilst high D2 strains occur in shear zones along the limbs of the folds, particularly on Else Platform. It therefore appears as though, during F2 fold development, strain was partitioned into the limbs, with resultant development of D2 shear zones. Parasitic F2 folds are common on the limbs of large F2 structures where they are open to isoclinal (Fig. 2.7(c)) and locally possess a well developed upright axial surface foliation that sometimes contains thin leucosomes. On a microscopic scale, 52 deforms the fabrics associated with D1. Semi-pelitic layered gneisses show new biotite foliations, and D1 garnets are sometimes fractured perpendicular to 52 and replaced by biotite. In metapelites, 52 is defined by biotite and locally also by new prismatic sillimanite. Primary sillimanite is strongly extended parallel to L2 and replaced by cordierite-spinel symplectites, and F2 microfolds with axial surface biotite are overgrown by cordierite. In post-D1 orthogneisses, quartz and feldspar define elongate granoblastic textures and garnet forms anhedral grains and grain aggregates elongate nLz. On Kamenistaya Platform, F2 folds are more open, and no high strain zones occur along F2 fold limbs suggesting that D2 strain decreases to the 37 Chapter 2, Geology of Jetry Peninsula south. The 52 fabric is also more poorly developed on Kamenistaya Platform, with only minor 52 biotite fabrics and more static syn-D2 reaction textures. Late durinED2, a series of major east to northeast trending, south-dipping shear zones formed on the limbs of the major F2 folds, with the most important of these being the Soviet Shear Zone (SSZ), on the southern limb of the dominant F2 fold on Else Plaform. The SSZ is up to a kilometre in width, whilst smaller D2 shear zones range in width from -5 - 100 metres. They are characterised by a progressive intensification of 52 toward the shear zones, culminating in the development of strong mylonitic fabrics. In the eastern part of Else Platform the SSZ is a wide (-lkm) region of mylonitic gneisses in which original lithological contacts are easily discernible. To the west the zone narrows to -400m accompanied by increasing strain, as evidenced by localised ultramylonites and the extreme attenuation of lithological units (Fig. 2.7(Ð). The southern edge of the shea¡ zone is more abrupt than the gradational northern edge of the zone. The 52 mylonitic foliation forms an intense east-west trending layering, which is generally defined by elongate mineral grains and grain aggregates. Porphyroblast tails and S-C fabrics indicate a south-up (reverse) sense of movement along a steeply dipping southeasterly plunging stretching lineation defined primarily by elongate qtrartz. The presence of recrystallised cordierite and minor aggregates of garnet in pelites affected by Dz shearing, and pyroxenes and hornblende in sheared mafic lithologies, indicate that D2 shearing occurred at elevated temperatures (>-650- 700"c). 2.4.3D3 An enigmatic feature of the structural configuration on Jetty Peninsula is a NNE-SSW trending synformal structure on southeastern Else Platform with a east dipping axial plane and a vertical to overturned eastern limb, the effects of which die 38 Chapter 2, Geology of Jetty Peninsula out northward. In this region, 51 and 52 are folded through -110o, from 080 to 190, and the N-S t¡ending St-Sz foliations on the SE edge of Else Platform are interpreted to represent the short limb of the D3 synform. This implies that D3 folds may be kink-like flexures. There are numerous parasitic folds around the hinge, some exhibiting a weak 53 foliation defined by new gro,wth of biotite overprinting D2 mylonitic fabrics (Fig.2.7(e)). The stretching lineations of the north-south trending fold are roughly parallel to the L2 lineation. There are two possible origins for these folds: l/ The N-S fotd may be a kink-like flexure related to D2 deformation associated with the high strain in the adjacent D2 high strain zone. 2l The fold may postdate D2, and is an essentially unrelated feature as part of a third deformation (D¡). The presence of a weakly defined foliation orthogonal to the 52 foliation, and the fact that the structure appears to over print the D2 shear zone and anFz synform in the very south of Else Platform, support the latter alternative. Although the similarity in the stretching lineation and fold a,xis with the F2 structures would tend to suggest that D2 and D3 are temporally related, the parallelism between the foliation in the Pan- African biotite granites and the 53 axial planar fabric implies that the timing of D3 is more likely to be related to these granites, which would therefore place the timing of D3 at c.500 Ma (Manton et aI., 1992). However it must be stressed that this interpretation for D3 remains tentative. Moderately deformed migmatised granitic dykes and large undeformed garnet bearing pegmatites crosscutting pre-D3 structures are approximately parallel to the axial surface of F3 folds. A definitive relationship between the dykes, pegmatites and F3 has not been established, however the similarity in orientations is striking and it is tentatively suggested that the dykes and pegmatites are synchronous with D3. The granitic dykes contain a strong foliation, termed 53¿ to distinguish it from the 53 foliation related to the large scale N-S trending fold, which may or may not be the 39 Chapter 2, Geology of Jefiy Peninsula same foliation. The S3u foliation is defined by biotite and thin leucosomes that form the axial surface to isoclinally folded leucosomes that are occasionally garnet bearing (Fi9.2.7(Ð). S¡a trends 010 and dips steeply to the east and is identical in orientation to 53. F3" folds plunge gently to the north and no consistent sense of asymmetry was found. 53 appears to be associated with shearing since earlier structures are dragged into the dykes in a left lateral and/or east up sense. 2.4,4Dashear zones A second generation of shear zones postdates F3 and most pegmatite intrusions, and are best developed on northern Else Platform. These shear zones, which represent the fourth phase of deformation (Da), are steeply dipping and generally trend in a NW-SE direction. D4 shear zones differ from D2 shear zones in that they are smaller scale, lower grade (- upper greenschist facies), and are generally more finely recrystallised. They are locally ultramylonitic and in places reactivate zones of D2 shear. They generally have a sinistral sense of strike slip movement, as is evidenced by the offset of a migmatitic granite dyke south of Melkoye Lake, but no consistent vertical sense of movement could be distinguished. The Da shear zones occur both as narrow zones (0.3 - 3 metres in width), along which pegmatite veins have often intruded during deformation, and larger scale zones of anastomosing mylonites. No pegmatites have been observed to postdate the mylonites suggesting granitic intrusive activity ceased during D4. Where the pegmatites are deformed, large K-feldspar porphyroclasts have asymmetric recrystallised tails that define Sa with biotite and ribbon quartz. Fine grained biotite, and less commonly muscovite, defines Sa in pelites and semi-pelites, whilst chlorite defines 54 in metabasites. 40 Chapter 2, Geology of Jetty Peninsula 2.4.5 Late Brittle Structures The final deformation involved the development of N-S to NE-SW trending brittle faults, which are particularly well developed on Kamenistaya Platform. These faults are assumed to be related to the development of the Lambert Graben in the late Palaeozoic to Mesozoic (Hofmann, 1991; Webb and Fielding, 1993), and intersect all rock types except the late quartz veins and lenses. There are a number of generations of faults, with the earliest recognisable generation trending north-south. An example of this generation is the fault which separates the granulites from the Permo-Triassic sediments to the west near Soyuz Base, which has an east-up sense of movement. Early north-south faults also exist further to the north on Kamenistaya Platform, but the sense of movement on these faults is unclear. The major north-south trending fault which cuts southeastern Else Platform has a dextral strike-slip component of around 600 metres, with an uncertain magnitude of vertical movement, and smaller scale dextral faults are seen elsewhere on Else Platform. North-south trending Carboniferous dolerite dykes also offset earlier dykes in a dextral sense, suggesting that they may have intruded during development of these faulß. Following the north-south faulting, a series of sinistral NE-SW trending faults offset the earlier faults, and are concentrated on Kamenistaya Platform. These faults appear to postdate Permo-Triassic sedimentation, as they displace the graben sediments to the west, north of Soyuz, perhaps explaining why the sediments which dominate the region south of Soyuz are not found on Else Platform. The spatial association of these faults with the Cretaceous alkaline mafic pipes suggests that the NE-SW trending faults may be of late Mesozoic age. The boundary between Else Platform and Kamenistaya Platform is also infened to follow one of these faults. The origin and significance of these faults is discussed further in section 2.5.3. 4t Chapter 2, Geology of Jetty Peninsula 2.5 Sm-Nd geochronology of the garnet granite The garnet granite on central Else Platform is particularly useful in helping to constrain the timing of D2 on Else Platform, as it is locally discordant to F2 fold structures yet contains an 52 axial planar foliation. It is therefore considered to have intruded during D2. The garnet granite contains subhedral, unzoned garnets (Fig. 2.I0), which are interpreted as being of igneous rather than metamorphic origin. Isotopic dating of the granite was performed using a Finnigan 261 mass spectrometer at the University of Adelaide. Garnet zonation in garnet granite 3 3 û 6) è0 Fe ,, oX N 2 8. û ôÈf 1 I Mg Ca Mn 0 0 1 2 3 4 cm Fig. 2.10. Compositional profile across a typical garnet ftom the Else Platform garnet granite, showing the lack of significant zonation The Sm-Nd isotopic system was chosen due to its relatively high closure temperature in garnet (600 t 30oC, Mezger et al. (1992)), and the anomalously high 1475¡r1¡r44Nd ratio of garnet which allows relatively accurate construction of isochrons, as the garnet point will effectively 'fix' the slope of the isochron. A two point isochron, in spite of its inherent uncertainties, was considered reasonable in this case as other minerals would have had Sm-Nd ratios too close to that of the whole rock to significantly affect the slope of the isochron. Sm-Nd dating using garnets has been shown to be a powerful geochronological tool in high grade terrains for dating both metamorphic and igneous events (e.g. van Breemen and Hawkesworth, 1980; Vance and O'Nions, 1990; Mezger et al., 1992). For the garnet granite, a two-point 42 Chapter 2, Geology of Jetty Peninsula garnet-whole rock Sm-Nd isochron gave a garnet crystallisation age of 939 Ma (Fig. 2.ll). As the isochron has only two points, it is not possible to derive a realistic error on this age; however, the two points have such a large separation in Sm/Nd ratios that, given the good statistics of the analyses (Appendix 2), they still constrain the slope of the isochron with some accuracy. Data used to obtain this isochron are contained in Appendix 2. 0.515 143Nd/ 144Nd Age = 939 Ma 0.514 I.R. = 0.5111801 garnet 0.513 0.5t2 whole rock 0.511 0.0 0.1 0.2 0.3 0.4 0.s 0.6 147Srn/l44Nd Fig. 2.11 Sm-Nd garnet-whole rock isochron for the Else Platforrr garnet granite (sample 964-EP85), indicating at age of c.939 IÙla. The 939 Ma age for the garnet granite is consistent with other geochronological data from the Else Platform region. Manton et al. (1992) have obtained a U-Pb zircon age of 940 + 20 Ma for the leucogneiss on Kamenistaya Platform which is correlated with the felsic dykes on Else Platform. The intrusion of the felsic dykes is interpreted to have occurred prior to garnet granite intrusion, but during the same event, so it would be expected that the ages of both intrusions would be similar recording the latter stages of high grade metarnorphism and deformation in the region. These two ages strongly suggest an age of c.940 Ma for D2 on Jetty Peninsula. A Ndlorra¡ whole rock model age of 1870 Ma for the garnet granite suggests that there was significant incorporation of older crustal material in the melt, and is significantly higher than the Sm/Nd model ages of 1300-1600 Ma obtained by 43 Chapter 2, Geology of Jetty Peninsula Munksgaard et al. (1992) for felsic granulites and charnockites from the Porthos Range of the NPCM. Such a relationship is not surprising considering the l-type nature of the Porthos Range orthogneisses (Munksgaard et aI., 1992) in comparison to the more aluminous S-type nature of the garnet granite. Similarly, the gneissic leucogranite and felsic dykes have S-type geochemical characteristics (Manton et al., 1992). Munksgaard et al. (1992) proposed a significant amount of incorporation of older crustal material into the Porthos Range felsic granulites and charnockites, but the higher model age of the garnet granite probably represent a significantly higher proportion of crustal anatexis, and would suggest that crust of Early Proterozoic or Archean age underlies much of the NPCM region. A further constraint on the evolution of Jetty Peninsula provided by the Sm- Nd isotopic history of the garnet granite is the fact that Sm and Nd apparently were not mobile after 940 Ma, and this places an upper constraint on the thermal evolution of the terrain since the Mid-Proterozoic. The closure temperature of the Sm-Nd system in garnet is believed to be around 600oC, following the experimental work of Mezger et al. (1992), and the closure of the system at 940 Ma implies that this temperature was not exceeded after that time. 2.6 Geological history and age constraints 2.6.1 Mid Proterozoic deformation and metamorphism The geological history of the Jetty Peninsula region, accompanied by age constraints is given in Table 2.1. The timing and geological setting of the deposition of the supracrustal sequence of pelitic, semi-pelitic, calcareous, and possibly minor volcanic sediments is not constrained. Black et al. (1987) demonstrated that parts of the Proterozoic complex in Enderby and Kemp Lands contain reworked Archean material, but no convincing evidence exists for whether such reworked material occurs in the NPCM. Intrusion of leucogneisses followed burial of the supracrustal sequence, probably coincident with the early stages of high grade metamorphism and M Chapter 2, Geology oÍ Jetty Peninsula Age (Ma) Event Deformation Intrusive Activity Meømorphism Eafly ro Mid Proterozoic? Igneous and sedimenøry protoliths Granuliæ c.1000 Ma#ï D1 Formation of a regional gneissic Intrusion of metanorphism @r-¿t) hyering leucogranite and P= -7 kba¡s, Continueddefomration producing megacrystic granite T= 800oC * gneissic fabrics in early graniæ. c.960 Maa Chamockiæ 940!20Mla# Leucogneiss c.920-940Ma40 Srain partitioning into large east- Granulite D2 west shea¡ zones with east Garnet gfanite metamorphism (Ds-oT) plunging lineation, Kilomere- P = 5.5 kba¡s, scale north-rending open folds T=700oC* with minor ærial foliation, -500lvfa# D3 Approximately north rending Granitic dykes. shea¡ zones with upright 53 krgepegmatite amphibolite facies? foliation. bodies. 500485 N4a #+ D¿ Upright, NW-trending mylonite Smallpeguntite greenschist facies. (D¡)t zones. verns. Uplift c.300Ìúa$ North- south rending brittle Mafic dykes. Hydrothermal faults. alteration. Exposure at surface, Permo-Triassic sedimentation c. 140 Mas Brittle faulting associaæd with Alkaline mafic and development of Lanbert Graben ultramafic dykes and stocks. * see Chapters 3 and 5; t Porthos and Aranis Ranges, Fiøsimons and Thost (1992); $ Tingey (1982);  Young and Black (1991); # lvfanton etal. (1992); $ Hotuann (1991); ß And¡onikov (1993). Table 2.1 Summary of the sEuctural and metanorphic history of Jetty Peninisula. 44a Chapter 2, Geology of Jetty Peninsula deformation. Medium pressure granulite facies metamorphism accompanied intense deformation, with at least two stages of progressive isoclinal folding. Porphyritic granites, fine grained pyroxene-bearing gneisses, early garnet granites and ultra-mafic pods intruded during this deformation. The timing of this event is still not tightly constrained. Tingey (1982) reported Rb-Sr isochron ages of 923 + 179 to 1005+ 87 Ma on high grade metamorphic rocks from the NPCM, and interpreted these ages to reflect peak metamorphic conditions, whilst Grew and Manton (1981) obtained a 896 Ma upper intercept age for zircons from a syn-metamorphic pegmatite from the Reinbolt Hills, 100 km to the east of Jetty Peninsula. A U-Pb age of 1000 + 10 Ma determined by Manton et aI. (1992) for metamorphic zircons from a two-pyroxene granulite from Else Platform is perhaps the most reliable estimate yet for peak metamorphism on Jetty Peninsula. Constraints on the nature of the metamorphic evolution of Jetty Peninsula and surrounding regions are given in following chapters. Intn¡sion of fine grained garnet-bearing felsic dykes occurred early within the development of tight to open E-W trending Fz folds and the initiation of shearing in a near vertical direction along an E-W plane, coincident with M2 metamorphism. This was followed by the intrusion of a garnet-bearing granite into the hinge of the fold. As discussed in section 2.4, D2 appears to have occurred at around 940 Ma (Manton et al., 1992; this study). A Rb-Sr mineral isochron from the gneissic leucogranite, dated at940 Ma by U-Pb zircon and correlated with the felsic dykes, gives a cooling age of 717+ 32Ma (Ri = 0.76L) (Manton et al., L992\, although the significance of this age is unclea¡. It is possible that this age may represent a mixing between a940 Ma age and a 500 Ma overprint, although a number of simila¡ dates elsewhere in the tenain (e.9.718+ 10 Ma [Rb-Sr] on a pegmatite from Kemp Land (Clarke, 1987); 770-760 Ma [U-Pb zircon] on post-tectonic igneous intrusions in the Rayner Complex (Black et al., 1987);787 Ma [Rb-Sr] on a post-charnockite leucogneiss and 755 + 6 Ma [K-Ar] on an undeformed metamorphosed dyke, both from the NPCM (Hensen et 45 Chapter 2, Geology of Jetty Peninsula a1.,1992)), suggest that some thermal or igneous event may have occurred around this time. 2,6.2 P an- t¡frican thermal event There is widespread evidence in East Antarctica for the resetting of Rb-Sr isotope systems in mica at approximately 500Ma (Tingey 1982; Sheraton et al., 1984; Black et al., 1987), and evidence for localised regions of high grade metamorphism and anatexis in various fragments of the terrain (Zhao et al., 1992; Ren et a1.,1992; Shiraishi et al., 1992: Dirks et al., 1993), yet the significance of the Pan-African '500 Ma event' in East Antarctica remains controversial. Manton et aL. (1992) were unable to obtain well constrained U-Pb zircon ages from the'migmatitic' granitic dykes from Jetty Peninsula, but their best estimates of crystallisation ages were 524ì,[la for non- porphyritic biotite granite and 565 Ma for porphyritic biotite granite, suggesting that they are Pan-African in age, whilst Rb-Sr mineral isochrons from the same bodies are 487+3 Ma (Ri = 0.724) and 483 + 13 Ma (Ri = 0.714) respectively. The post- deformational pegmatites gave concordant U-Pb ages of 495-507 Ma (Manton et al., L992). The foliated and migmatitic nature of the biotite granites and their geochemical signature reflecting partial melting of anhydrous granulite facies metasediments (Manton et a1., L992) would suggest that temperatures were significantly elevated at the time. The development of asymmetric NNE-SSW trending F3 folds quite probably occurred at a similar time to the intrusion of fine grained biotite granite dykes, as discussed in section 2.3.3., which would imply that these stuctures may also be of Pan-African age. Most rock types on Jetty Peninsula arc affected by some degree of low temperature retrogression, particularly within D4 shear zones where the high grade assemblages are completely replaced by muscovite, biotite and chlorite which define the mylonitic foliation with ribbon quwtz. These upper greenschist facies shear zones and accompanying nalrow pegmatites are correlated with other upper greenschist 46 Chapter 2, Geology of Jetty Peninsula facies shear zones in East Antarctica which have been related to the isotopic resetting and uplift of the tenain during the 500 Ma event. In the case of Jetty Peninsula, where evidence exists for high temperature conditions at c.500 Ma, these shear zones and retrogression probably represent the waning stages of this event. 2.6,3 Lambert Graben development A feature of the Jetty Peninsula region is the abundant evidence for a prolonged history of rifting, magmatism and sedimentation associated with the development of the Lambert Graben system. The Lambert Graben is a deep-seated rift zone, extending inland for a distance of 700 km, which may have been initiated as early as the Cambrian (Manton et a1., 1992). The Lambert Graben was certainly active in the Late Palaeozoic and Mesozoic, as evidenced by magmatism and sedimentation of this age (Hofmann, I99I; Andronikov, 1992; Webb and Fielding, L993) and probably remains a zoîe of crustal weakness today (Kurinin and Grikurov, L982). Deep crustal seismic reflection studies by Soviet geologists have revealed that Jetty Peninsula is located on the surface expression of a deep-seated fault system which extends to the Moho (Ravich et al, 1978; Kurinin and Grikurov, l9g2). The orientation of the 550-500 Ma migmatitic biotite granite dykes on Jetty Peninsula is broadly parallel to the rift zone and associated mafic dykes, suggesting that the granite intrusion may be associated with early development of the Lambert rift structure, an association also proposed by Manton et al. (lgg2), and which Sheraton and Black (19S3) also suggested for the similarly aged Landing Bluff granite on the northeastem margin of the Amery Ice Shelf. However, this association remains highly tentative. Although occasional undeformed mafic dykes are observed elsewhere in the NPCM (e.g. the Porthos Range, Fitzsimons and Thost (L992) and the Athos Range (this study)), the swarms of north-south trending mafic dykes on Jetty peninsula are 47 Chapter 2, Geology of Jetry Peninsula unique in the region in terms of their abundance and parallelism. Abundant vesicles suggest that the dykes were emplaced relatively close to the surface, and therefore their intrusion probably followed significant uplift of the terrain. A K-Ar age of c.300 Ma (Carboniferous) obtained by Hofmann (1991) for these dykes is supported by the observations that these mafic intrusions are not found in the terrestrial Permo- Triassic sediments. The fact that the majority of the dykes are parallel to the graben and occupy sub-vertical fault structures implies that the Carboniferous mafic dykes were emplaced during extension resulting in graben formation. Deposition of sandstones into the Lambert Graben and associated Beaver Lake Graben during the Permian and Triassic occurred in a terrestrial setting within a northward flowing braided river system (Mond, L972; Webb and Fielding, 1993). This was followed by the intrusion of alkaline mafic pods containing lherzolite nodules, which have been dated by K-Ar methods as early Cretaceous (Andronikov, 1990), and which were probably associated with the separation of India and Antarctica (Kent, 1991), when the Lambert Graben was probably reactivated as an aulocagen. Possible continuing movement on the faults resulted in a faulted contact between the sediments and the Proterozoic basement, and hydrothermal activity along the faults resulted in the formation of quartz veins, with associated minor epidote and chlorite alteration. Apatite fission track data (Arne et al., 1993) suggest that the region underwent significant exhumation in the Mesozoic, at around 150 Ma, and this uplift may also be related to the early stages of the breakup of Gondwana. 2.6 Correlations and discussion The geological relationships on Jetty Peninsula form an important link between the northern Prince Charles Mountains and outcrops along Prydz Bay, and therefore correlations with observations by other authors in these sunounding regions is important. Proterozotc relationships in the Aramis and Porthos Ranges in the NPCM to the west of Jetty Peninsula have been described in detail by Fitzsimons and 48 Chapter 2, Geology of Jetty Peninsula Thost (L992) and Thost and Hensen (1992). Along the Prydz Bay coast, specific relationships in the Larsemann Hills have been described by Stüwe et a1., (1989), Stüwe and Powell (1989b) and Dirks et al. (1993), for the Brattsrand Bluffs (Fitzsimons and Harley, 1992a) and the Rauer Islands (Harley 1987, 1988; Harley and Fitzsimons, 1991). In their description of the geology of the Porthos and Aramis Ranges, Fitzsimons and Thost (1992) document a similar structural history to that on Jetty Peninsula, although lithologically, the region is less diverse, being dominantly felsic orthogneiss which is interlayered with minor calc-silicate, semi-pelitic and extremely rare pelitic gneisses. D1 on Jetty Peninsula is correlated with D3 of Fitzsimons and Thost (1992>, which also resulled in the development of the dominant gneissic layering. D1 and D2 of Fitzsimons and Thost (1992) represent folded fabrics in xenoliths and isoclinally folded leucocratic segregations, which were also encountered on Jetty Peninsula but were not classed as a separate deformation. D3 produced recumbent stn¡ctures that are presened in zones of low D6 strain. D4 and D5 stn¡ctures are minor, (D4 related to progressive D3 deformation and D5 a probable earlier stage of D6) and the regional orientation of most features is determined by D6. D6 is heterogeneous and postdates the intrusion of extensive charnockite bodies that have been correlated with the Mawson charnockite which has intrusive ages of 954!12 and 985+29Ma (Young and Black 1990) implying that D6 is younger than this. It is likely that the Jetty Peninsula charnockites a¡e also of a similar age to the Mawson Charnockite. Fitzsimons and Thost (1992) considered D6 to have occurred at around 920 IÙla, on the basis of correlation with the Mawson Coast, where Young and Black (1991) have dated a metamorphic rim on a zircon reflecting post- charnockite deformation (D3, Cla¡ke, 1988) at92I+ 19 Ma. D6 high strain zones are characterised by intense upright fabrics and an easterly plunging lineation that obliterates earlier structures. Between the high strain regions flat lying D3 structures are preserved. Low D6 strain results in the formation of easterly plunging antiforms 49 Chapter 2, Geology of Jetty Peninsula and synforms. D3 structures can be readily correlated with D1 in the structural framework developed for Jetty Peninsula since each produced the early regional gneissic banding. D6 structures in the Aramis and Porthos Ranges have the same These orientation as D2 on Jetty Peninsula and are considered here to be contiguous. in correlations are based in part on work undertaken as part of this study elsewhere as D2 on the NPCM (Appendix 1). If D6 of Fitzsimons and Thost (1992) is the same the Jetty peninsula, then their estimate of 920 Ma for this event agrees well with estimate of 940 Ma for the same event on Jetty Peninsula' Similarly a tentative correlation is made between D2 on Jetty Peninsula and D3 on the Mawson Coast' Large scale open folding about steeply ptunging N-S axis overprints the east-west rrending D6 folds (Nichols, lgg2), and this is conelated with D3 on Jetty Peninsula' (1991) In the Reinbolt Hills, 80 km east of Jetty Peninsula, Nichols and Berry described a gneissic fabric associated with D1 deformation, folded by tight to isoclinal east-west trending F2 folds. The broad similarity between these deformations and D1 and D2 on Jetty Peninsula suggests a direct correlation between these events. F3 in the Reinbolt Hills is similar in its orientation to F3 on Jetty peninsula, although Nichols and Berry (1991) interpreted it as being coaxial to F2 and therefore part of the same continuous event as F2. The syn-tectonic Reinbolt Charnockite can tentatively be correlated with the Jetty Peninsula, Crohn Massif and Mawson Charnockites on the basis of its relationship with the deformational history. This charnockite was dated by Grew and Manton (1981) using U-Pb in zircons at 896 Ma, although this age is somewhat dubious, as most points on the chord used to obtain the age were from zircons in a pegmatite lens. Evidence for magmatic and thermal activity at c.500 Ma is widespread (1992) throughout East Antarctica. Similar ages to those obtained by Manton et aL for the Jetty peninsula biotite granites have been derived for small syenogranite dykes + 495-9 in the Larsemann Hills (563 + 27,529 +15/-16 Ma [zircon Pb-Pb], 494.3 1'3, 50 Chapter 2, Geology of Jetry Peninsula (1992) considered the + 4.6 Ma ¡a04¡7394¡l) by Zhao etaI. (lgg2) and Ren etar\. who bodies to be derived from local anatexis, a proposal supported by geochemical information (Stüwe et al., 1989). Pegmatites cross cutting the syenogranites have et a1., been dated ar 535 + 18 and 52I + 11 Ma (zircon Pb-Pb; Zhao et aL., 1992; Ren 1991), which are similar to the dates obtained by Manton et aJ. (1992) for the Jetty peninsula pegmatites. However, granulite facies conditions inferred for the Larsemann Hills at c.500 Ma by Ren et al. (1992) and Dirks et 41. (1993) suggest that the event was significantly more intense there than in most other regions such as Jetty peninsula. Two discrete high grade metamorphic events at c.1000Ma and c.525Ma have also been inferred from the Lutzow-Holm Complex in the western part of East Antarctica (Shiraishi et al., 1992)' The most significant feature of the observations of the geological evolution of Jetty Peninsula is the relative ease of correlation of the deformational and intrusive events with other fragments of the very broad and extensive Proterozoic terrain. In pafticular, the outcrops along eastern margin of the Amery Ice Shelf, the entire NPCM, and the rest of MacRobertson Land to the Mawson Coast all appear to have a very similar history. In all these regions, a gneissic foliation related to peak metamorphism, probably at c.1000 Ma, is intruded by charnockites and overprinted by large scale, generally east-west trending folds at c.950-900 Ma, with a slightly lower grade metamorphic overprint. A minor thermal event may have occurred at c.750 Ma, followed by a more extensive and loca[y intense thermal and intrusive event at 550-500 Ma, which locatly resulted in further high grade metamorphism. The phanerozoic history of the region is dominated by the uplift, rifting and mafic magmatism associated with the Lambert Graben development and breakup of Gondwana. The unique feature of Jetty Peninsula is that it provides perhaps the most complete and well exposed record of this history and hence is ideal for producing a correlative structural and metamorphic framework for the sunounding terrain' In the following chapters, constraints on the metamorphic history are developed in order to 51 Chøpter 2, Geotogy of Jetry Peninsula tectonic cycle determine whether the granulite met¿rmorphism represents a continuous the or whether the region has undergone a more complex history, and to constrain geodynamic processes which resulted in the observed structural and metamo{phic evolution of the region. 52 Chapter j, Jetry Peninsula m.etannrphism Chapter 3 Metamorphic evolution of pelitic granulites from Jetty Peninsula, northern Prince Charles Mountains 3.1 Introduction The geology of Jetty Peninsula, as described in Chapter 2, is unlike many other regions in the northern Prince Charles Mountains in that it is dominated by a sequence of layered metasedimentary gneisses, and therefore contains many lithologies which are useful in determining the metamorphic evolution of the tenain. The metamorphic evolution of Jetty Peninsula is of significant interest in the overall investigation of the evolution of the Proterozoic of East Antarctica as it is one of the only regions in MacRobertson Land in which detailed geological mapping of a Ithologically diverse supracrustal sequence may be undertaken. It is therefore an important link in relating observations from the Prydz Bay coast with those from areas in MacRobertson Land and the Rayner Complex to the west. No detailed metamorphic study has previously been made of Jetty Peninsula, although Grew (1986) briefly described the petrography of pelitic, ultramafic and calc-silicate lithologies from the region, concluding that they reflected moderate pressure (< 7 kbars) granulite facies metamorphism of dominantly sedimentary precursors. The only other published detailed metamorphic studies in the northern Prince Charles Mountains have been by Fitzsimons and Thost (1992) and Fitzsimons and Harley (1994), who have inferred isobaric cooling histories from c.800oC and 7 kbar for granulites from the Aramis and Porthos Ranges, 80 km to the west. The nearest outcrop to the east of Jetty Peninsula are the Reinbolt Hills, 80 km across the Amery Ice Shelf, for which Nichols and Berry (1991) inferred a near-isothermal decompressional evolution from peak conditions of around 800oC and 7 kbars. Jetty 53 Chapter 3, Jetry Peninsula metanorphism Peninsula therefore provides a link between these two regions with vastly different apparent retrograde P-T histories. This chapter deals primarily with textural relationships within pelitic lithologies from Jetty Peninsula, whilst the calc-silicates are dealt with separately in Chapter 5. Much of the work for this chapter was done in collaboration with co- workers at the University of Melbourne, and forms the basis for a published manuscript (Hand et al., 1994). Collaborative research is referenced where appropriate throughout the following discussion, and additional observations and findings not from my own research are referenced to Hand et al. (1994). 3.2 Petrography of pelites The pelites from Jetty Peninsula contain a variably well developed 51 foliation, which is defined by phases forming part of the peak metamorphic assemblage. An early coarse grained assemblage is defined by garnet and sillimanite, and generally also contains ilmenite, plagioclase, K-feldspar and qrrarlz, with or without biotite, rutile and spinel. This assemblage is associated with D1 deformation, and sillimanite defines the 51 fabric. This assemblage will therefore be termed Mt. This assemblage is variably consumed by biotite fabrics, and subsequent cordierite- bearing coronas and symplectites, which grew during the D2 deformation, and which hereafter will be referred to as M2 assemblages. The terms M1 and M2 are intended to relate the assemblages to observed structures, and do not necessarily imply unrelated metamorphic events. The pelites have been subdivided into four different assemblages which reflect differences in bulk composition. In assemblage l, garnet, sillimanite and biotite are separated by symplectites of cordierite and spinel. Assemblage 2 is similar to assemblage 1, but lacks significant biotite, whilst in assemblage 3 the cordierite coronas do not contain spinel. The fourth assemblage contains a coarse garnet- 54 Chapter 3, Jetty Peninsula metamorphisrn orthopyroxene-biotite assemblage which is partially consumed by cordierite and seconda-ry orthopyroxene. Generalised petrographic relationships within these four assemblages are summarised in Table 3.1. st sill sp bi cd opx ksp qz pl ru ilm 1 PRo PRzs R¿s Rtz Rz¡ P P P - IIPRa 2a PRe PRz IrzR¿ r Rz¡ PR¿ r 11 I1PR+ 2b P P IzR¿ ra R23 p PRz¿ P PIT PR¿ 3 P P r R P PRz¿ PRz+ -p 4 P IrRrr. Rr.¡ PRz¿ P D PRr¿ Table 3.1 Generalised petrographic and textural relationships in selected samples from the four assemblages. Assemblagel-964.41 ;Assemblage2(a)-9&-6068(b)96a-60a; Assemblage3-964-624; Assemblage 4 -9&-665; P = pffi of the inferred peak assemblage; Il = inclusion in garnet; 12 = inclusion in sillimanite; Rl = defining 52 fabric; R2 = consuming primary garnet; R3 = consuming garnet and sillimanite and/or biotite; R+ = symplectitic intergrowths with cordierite consuming garnet and sillimanite and/or biotite; R5 = coronas a¡ound 52 biotite; R¡ = euhed¡al garnets in cordierite coronas adjacent to primary gfirnet; R? = fibres overgrowing primary sillimanite; R8 = coronas between cordierite and spineVilmenite. (lower case denotes minor phasss) 3.2.1Assemblage 1 In these rocks, the primary assemblage contains garnet, sillimanite, ilmenite, quartz, plagioclase and K-feldspar. Sillimanite often contains inclusions of spinel, whilst garnet commonly contains inclusions of biotite and ilmenite, and less commonly spinel and/or sillimanite. This peak assemblage is consumed and enveloped by a biotite fabric which defines 52 and which also forms within pull- aparts of gamets which have been extended along theL2lineation. Where sillimanite has been deformed and dynamically recrystallised during D2, aggregates of biotite occur within hinge lines, suggesting that biotite and sillimanite were stable together early durin E Dz. With increasing strain, sillimanite and garnet were pulled apart parallel to L2. Where sillimanite has been pulled apart, symplectites of cordierite and spinel, with or without ilmenite, replace sillimanite (Fig. 3.1(b)). In garnet the pull- aparts were initially infilled by biotite, although cordierite is locally developed along the contact between garnet and secondary biotite (Fig. 3.1(d)). Garnet, sillimanite and 55 Chapter 3, Jetty Peninsub metanrphism .t d a, ':l "cd t -{, rJ I /it _-L Fig. 3.1 Reaction textures in pelites from Jetty Peninsula, sample 964-61 (Assemblage 1). In all cases the width of the field of view is 2 mm. (a) Symplectites of cordierite and spinel separating biotite (defining the 52 fabric) and sillimanite. Arrow indicates inclusion of biotite within spinel. (b) Sillimanite extended along the L2 lineation, and separated from peak gamet and Sz biotite by symplectites of spinel, cordierite and ilmenite. (c) Sillimanite separated from Sz biotite by cordierite and spinel. (d) Fracture in primary gamet containing biotite, which is partly consumed by cordierite. Note growth of ilmenite at edge of biotite, and small euhedral secondary gamets within the cordierite. 56 Chapter i, Jeny Peninsuln metnmarphism e ( å r I 3q i a.t t. ¡ "l o,\ \ ¡ , { - -{D t . "--- 4, g r.:t Fig. 3.1 (continued) Reaction textures in pelites from Jetty Peninsula. In all cases the long axis of the field of view is 2 mm. (e) Sillimanite, biotite and garnet separated by cordierite and spinel with minor ilmenite. Arrows indicate spinel overgrowing 52 biotite, and late gamet overgrowing primary gamet. Sample 964-67 (f) Symplectite of cordierite and spinel separating gamet and sillimanite. Note that in the absence of biotite, ilmenite is a reactant rather than a product. Sample 964-83 (g,h) Cordierite recrystallised in 52 high strain fabric, with secondary garnet and spinel growing within recrystallised cordierite. Sample 964-59. (g) - plane polarised light; (h) - crossed polars 57 Chapter 3, Jetty Penirculn metamorphism F( Fig. 3.1 (continued) Reaction textures in pelites from Jetty Peninsula. In all cases the long axis of the field of view is 2 mm. (i) Biotite-sillimanite fabric defining 52 separated from primary gamet by cordierite. Sample 964-51 (Assemblage 3). () Symplectite of cordierite and quartz with minor biotite consuming gamet. Sample 964-51 (Assemblage 3) (k) Þrimary gamet and sillimanite separated by cordierite. Secondary sillimanite (sillz) forms fibres and coronas at contacts of sillimanite and cordierite, and secondary gamet (gtz) forms narrow rims on primary garnet. Sample 964-59 (Assemblage 1). (l) Euhedral secondary gamets overgrowing cordierite corona between gamet and sillimanite. Sample 964-83 (Assemblage 2). 58 Chapter 3, Jetry Peninsula metamorphisrn 52 biotite are generally separated by cordierite and spinel, with the spinel generally occurring as small grains or symplectitic intergrowths adjacent to sillimanite, or as rims on biotite (Fig. 3.1(a-c, e)). Coarse-grained spinel and ilmenite inclusions in fractured primary garnets are commonly surrounded by coronas of cordierite. The occurrence of ilmenite as a phase within the cordierite coronas appears to be dependent on the abundance of biotite as a reactant, with secondary ilmenite being spatially associated with biotite (Fig. 3.1 (b, d, Ð). In places, a second generation of garnet occurs as fine grained (< 50pm) euhedral grains within cordierite coronas and along the contacts between cordierite and garnet (Fig.3.1(d, 1)). In addition, a second generation of sillimanite occurs as fibres oriented parallel toL2 on the ends of deformed 51 sillimanite (Fig.3.1 (k)). This secondary sillimanite also often occurs as narrow coronas around ilmenite and less commonly spinel, where they are in contact with cordierite. 3.2.2 Assemblage 2 These rocks, which are particularly common on Kamenistaya Platform, generally have a granoblastic appearance, and are characterised by symplectites of cordierite and spinel separating primary ilmenite, garnet and sillimanite, with little or no biotite present. Due to the lack of biotite, little or no 52 fabric is evident in this assemblage. Feldspar content is variable, and is often a relatively minor phase. For example, in some rocks (e.g.964-606E), K-feldspar is absent and plagioclase is rare, with the matrix consisting primarily of quartz, cordierite, ilmenite and minor rutile. Ilmenite is usually a reactant within the cordierite-producing reactions (Fig. 3.1 (f)), although where biotite is present, small euhedral ilmenites occur within the cordierite coronas. In some specimens (e.g. 964-604), rare coronas and symplectites of cordierite and quartz form around garnet at contacts with plagioclase. Where K- feldspar is also present, biotite forms within these cordierite-quartz symplectites. Fine-grained (20-25¡rm), euhedral secondary garnets occur locally within the s9 Chapter 3, Jetty Peninsula metamorphism cordierite and at the margins of larger primary garnets. Rutile often occurs within this assemblage, typically as a matrix phase although also as an inclusion within coarse- grained gamet, K-feldspar, quartz and S1 sillimanite. Within D2 shear zones, 52 is defined by a coarse fabric consisting of elongate cordierite and quartz. Garnet, sillimanite and ilmenite are locally separated by cordierite, with the cordierite recrystallised by rhe D2 shea¡ fabric (Fig. 3.1(g,h)). The occulrence of fine grained spinel within this recrystallised cordierite suggests that reaction between primary garnet, sillimanite and biotite continued during waning D2 deformation. 3.2.3 Assemblage 3 This assemblage is much the same as assemblage I, with the exception that there is no spinel in the cordierite coronas between garnet and sillimanite (Fig. 3.1(Ð). In this assemblage, cordierite coronas which locatly separate sillimanite, garnet and biotite often contain ilmenite, particularly when biotite is a reactant. A feature of this assemblage is the abundance of quartz-bearing coronas around garnet. In particular, these coronas are symplectites of cordierite and quartz between garnet and plagioclase, although where K-feldspar is present, biotite also exists within the coronas (Fig. 3.10)). In other places, garnet is consumed by symplectites of plagioclase and quartz, with or without cordierite. 3.2.4 Assemblage 4 An orthopyroxene-bearing pelitic assemblage (assemblage 4) exists at only one known locality on Jetty Peninsula, in a lens on northeastern Else Platform (sample 964-665; E. Grew, 1991, pers. comm.). In this sample, a dominant biotite fabric envelops a coarse primary garnet-orthopyroxene assemblage, and partially consumes the orthopyroxene (Fig. 3.2(b)). Biotite also exists as inclusions within the garnet and clinopyroxene. A second generation of orthopyroxene forms symplectites 60 Chnpter 3, Jetty Peninsula metamorphism a a I Ì-' gt I Fig. 3.2 Reaction textures in orthopyroxene-bearing pelites from Jetty Peninsula. Sample 964-665, Assemblage 4. Width of field of view is 2 mm. (a) Symplectite of cordierite and orthopyroxene with minor plagioclase, separating primary gamet from Sz biotite. (b) Optically continuous primary orthopyroxene consumed by 52 biotite. (c) Cordierite separating corroded Sz biotìtes, accompanied by the growth of secondary orthopyroxene (arrowed). 61 Chapter 3, Jetty Peninsula m,etamorphism with cordierite and plagioclase between garnet and biotite (Fig. 3.2(a)). In the inner part of these symplectites, immediately adjacent to garnet, plagioclase is the dominant phase accompanying cordierite, whilst orthopyroxene is the dominant phase adjacent to biotite. Cordierite and orthopyroxene also consume biotite in the matrix (Fig. 3.2(c)). 3.2.5 Mineral Chemistry Garnet on Jetty Peninsula is generally almandine-rich, with Xp" most commonly being between 0.65 and 0.78, although cores of garnets in the most magnesian (orthopyroxene-bearing) assemblages have Xp" as low as 0.55. The remaining content of the garnets is dominantly pyrope, with grossular contents of 2 to 6 mol%o, and spessartine between 1 to 3 molVo. Late small secondary garnets are typically slightly more almandine-rich than the primary garnets within the same rock. The garnets are typically zoned, with a marked increase in X¡s towards the rim, and a corresponding increase in spessartine and almandine components. This zonation probably represents a retrograde effect due to the consumption of the garnet by more magnesian phases such as cordierite and biotite. Cordierite does not exhibit any significant zonation, and Xp. values range from 0.33 in sample 964-56 down to 0.16 in sample 964-83, with little variation within specimens. However, in general, XFe values drop marginally adjacent to garnets, with a decrease in Xps of no more than 0.02. Spinel is hercynite-rich but in all samples analysed is consistently more magnesian than co-existing garnets, with X¡, values generally ranging between 0.65 and 0.75. Most spinels have significant gahnite and chromite components, with up to 6.4 wt%o ZnO (Hand et a1., 1994) and with Cr2O3 contents typically ranging between 0.5 and 2.0wtVo, although one exceptional, corroded primary spinel in sample 964- 606E has around l4wtvo Cr2o3 and lÙwtvo zno. Tio2,Fe2o3 and Mno contents of 62 Chapter 3, Jetty Peninsula metamorphism the spinels are consistently very low. In general, prograde spinels have higher gahnite and chromite contents than the secondary spinel in symplectites with cordierite. Biotite has very variable composition between specimens, ranging from an Xps of 0.19 ro 0.25 in 964-665 (44) ro 0.45 to 0.50 in 964-56 (42). The variation is also wide within samples, with D2 biotites generally having higher Xps values than biotiæ inclusions within primary garnets. Orthopyroxene compositions in sample 964-665 range in Xpe between 0.35 and 0.38, with little variation between the coarse primary orthopyroxene and the finer secondary orthopyroxene. However, the primary orthopyroxenes have a slight increase in Xpe at the margin, where they are consumed by cordierite and biotite. lúzOZ contents vary from up to 4.5 wtVo in the primary orthopyroxene, to 3.0-3.5 wtVo in secondary orthopyroxene. A rimward decrease in Al2O3 conesponds to the rimward decrease in Xps in the primary grains. Plagioclase feldspar is characterised by anorthite contents of 38-50 mo|Vo, with plagioclase resulting from the breakdown of garnet often having a slightly higher anorthite content than plagioclase within the matrix. K-feldspar has orthoclase components of 90-95 mol%o, and ilmenite and rutile are both almost pure, with ilmenite having only minor Mn substitution. 3.3 Interpretation of mineral parageneses As shown in previous studies (e.g. Powell and Sandiford, 1988; Clarke et al., 1989), the Ti-bearing phases, rutile, ilmenite and Ti-hematite, are important in the mineral parageneses of many metapelites. On Jetty Peninsula, the behavior of ilmenite is variable, behaving as both a reactant and a product during cordierite- producing reactions, and this is intimately related to the abundance of biotite. The incorporation of significant TiO2 into biotite has meant that rutile was not involved in 63 Chapter 3, Jetty Peninsula metamorphism the ilmenite-bearing reactions. In order to consider ilmenite, magnetite and rutile bearing assemblages, Clarke et al. (1989) extended the petrogenetic grids for the KFMASH (K2o-Feo-Mgo-Al2o3-sio2-H2o) model sysrem developed by Granr (1985) and Hensen (1986) by also considering TiO2 and Fe2O3. This resulted in the development of a grid in the KFMASHTO system (where T is Tio2 and o is o2), which has since been implemented and shown to be useful in explaining pelitic reaction textures from a number of low to medium pressure granulite facies terrains (e.g. Stüwe and Powell, 1989b,c; Cla¡ke et al., 1990). In the case of the pelites from Jetty Peninsula, only ilmenite-bearing reactions are relevant, as magnetite is absent and rutile is not always present. As a result, only rutile-absent reactions in the KFMASHT system need to be considered. The most important mineral textures are those which formed during the D2 deformation and which record the P-T evolution following M1. These textures can be divided into two main groups, on the basis of whether or not they involve the development of secondary spinel. Initially, reactions not involving spinel witl be considered, as the abundance of zinc in the secondary spinel introduces added complexities, which shall be discussed further in section 3.3.3. 3.3.1 Assemblage 3 (Spinel-absent aluminous M2 assemblages) Pressure-temperature pseudosections in the KFMASHT system appropriate for the pelites on Jetty Peninsula are shown in Fig. 3.3. Pseudosections have been used here instead of simple petrogenetic grids because of the dependence of bulk composition on the reaction history of the va¡ious assemblages. The pseudosections are based on the KFMASHT grid shown in Fig. 3.3(a), which is projecred from qtrartz, K-feldspar, ilmenite and melt (after Clarke et al., 1989). The inferred peak assemblages are characterised by the association of garnet and sillimanite with the absence of cordierite, and these are overprinted by cordierite 64 Chapter 3, Jetty Peninsula metamorphism ASHT + ilm ! x + +l P CLo gt sill bi o(r gt cd sill cd sp sill o(s) ll bi cd (a) Feo opx Mgo P gt sill bi ô gt cd sill opx cd gt gt cd sp bi gt cd bi cd M1 assemblage M2 cordierite-bearing coronas and symplectites Figure 3.3 quAitative KFMASHT P-T pseudosections for compositions appropriaûe to Jetty peninsula metapelites, based on KFMASH univariant equilibria (after Clarke & Powell, 1991). (a) KFMASHT grid, and AFM diagran showing bulk compositions of assemblages in pseudo-sections. Boxes labelled b,c,d on the grid indicate the approximate position of pseudosections in Figs 6b, &,6. (b) Pseudosection for bulk composition (1) (the most aluminous on the AFM diagram). Peak conditions are cha¡acterized by the stability of gamet and sillimanite and the absence of cordierite. The position of syn-D2 cordierite-bea¡ing assemblages are indicated at lower pressures, and the shaded arrow is the infened syn-D2 P-T path. The solid arrow represents the apparent path indicated by textures in bi-absent rocks (e.g. 93583). (c) Pseudosection for bulk composition (2). Shaded and solid arrows a¡e the same as for @g. 6b). (d) Pseudosection for bulk composition (3), which approximates the composition of assemblag e 4 (964665). The peak assemblage is gamet-orthopyroxene. Conditions during D2 are initially cha¡acterized by the stability ofgarnet and biotite. These a¡e locally repiaced by ofhopyroxene- cordierite assemblages. Adapted from Hand et â1. (1994) 65 Chapter 3, Jetty Peninsula metwnorphism which developed at the expense of primary garnet and sillimanite. It is clear from the pseudosections in Fig. 3.3 that the stabilisation of cordierite at the expense of garnet and sillimanite reflects lower pressure conditions, and that the association of garnet and sillimanite without cordierite is unstable at these conditions for rocks of this bulk composition. However, although the cordierite coronas between garnet and sillimanite give the appearance of decompression, the stabilisation of biotite between these two assemblages introduces complexity to this evolution. It therefore must be taken into account that, sometime between the stabilisation of the M1 assemblages and the M2 coronas, the assemblage biotite-sillimanite was stable. During D2, biotite-sillimanite and garnet reacted to produce cordierite and ilmenite, resulting in the stable association of either sillimanite-cordierite-ilmenite or biotite-cordierite-ilmenite with primary garnet. The development of these two assemblages appears to reflect grain scale variations in bulk composition, with biotite-absent assemblages developing in parts of the rock with abundant sillimanite, and biotite parageneses being stable close to garnet. The two assemblages can be found within the same thin section suggesting that equilibration volumes are small. The stabilisation of cordierite and spinel following the development of biotite fabrics can be best explained by heating during D2. This apparent heating path during D2 is indicated by the horizontal ¿urows on the P-T pseudosections in Figs. 3.3(b, c and d). There are no constraints as to whether this heating was isobaric or not, and it is conceivable that this heating occuned during a continuous decompressional evolution from peak conditions. 3.3.2 Assemblage 4 (Orthopyroxene-bearing pelites) The orthopyroxene-bearing sample (964-665) differs in bulk composition from the other pelites in that it is less aluminous and more magnesian. As described earlier, primary orthopyroxene and garnet are overprinted by a strong 52 biotite foliation. This 52 biotite largely replaces orthopyroxene but does not appear to replace garnet, indicating that garnet and biotite were stabilised with respect to 66 Chapter 3, Jetty Peninsula metarnorphism garnet-orthopyroxene. The formation of cordierite and fine-grained orthopyroxene around garnet is consistent with the M2 inferred path from more aluminous pelites. Phase relationships appropriate for this bulk composition are portrayed in Figure 3.3(d). For compositions progressively away from garnet, the sequence of assemblages is orthopyroxene-cordierite-garnet, orthopyroxene-cordierite and orthopyroxene-cordierite-biotite. Thus, close to garnet, orthopyroxene and cordierite coexist with garnet and in the matrix away from garnet, orthopyroxene-cordierite and biotite coexist. cd bi sill bisill gt cd sp sill gt bi sp sill Zn HT +Zn +ilm+q+ksp+l Syn-D2 cordieriþ-bearing coronas and symplectiÞr M Fig. 3.4 KFMASHTZn pseudosection, indicating the qualiøtive effect of adding zinc to KFMASHT for a 4{k composition (indicated on the AFMZn tetrahedron in the inset). fne Oasnø ünes indicate the KFMASHT divariant fields from Fig. 3.3(a). The stability field oi spinel is expanded by the addition of zinc and, as a rgsllL spinel-bearing fields will move to lo\ryeitemperatutes and higher pressures. For the indicaæd M2 P-T path inferred for Jetty Peninsula, the addition of zinc will sablise spinel-bearing assemblages at the expense of ea¡lier M2 sillimanite and biotite. The reaction: bi + sill = gt + cd + sp is indicaæd in rhe inset. Adapted from Hand et al. (1994). 3.3.3. Assemblage I (spinel- and biotite-bearing reaction textures) An interesting feature of the pelites from Jetty Peninsula is the fact that the occurrence of spinel does not appear to be wholly dependent on the aluminous or ferrous nature of the bulk composition. Although spinel-bearing assemblages are 67 Chapter 3, Jetty Peninsula metamorphism favoured by locally aluminous bulk compositions, (i.e. close to sillimanite), assemblages that lack spinel within reaction textures are also locally sillimanite-rich, suggesting that the growth of spinel is influenced by additional factors. A feature of the assemblages which contain secondary spinel is the significant amount of Zn within the spinel. Zinc has been shown to exert a significant influence on the stabitity of spinel-bearing assemblages (e.g. Shulters and Bohlen, 1989; Nichols et al., 1992), as almost all zinc in a rock will be preferentially incorporated into spinel. To account for the effect of zinc in metapelitic rocks, Hand et al. (1994) expanded the KFMASHT system shown in Fig. 3.3 to include zinc. The resulting I(FMASHTZn pseudosection is shown in Fig. 3.4, and this shows the significant expansion of the stability fields of spinel-bearing assemblages with the addition of zinc into the system. For the infened post-peak P-T path, reactions which consume biotite and sillimanite will produce spinel in addition to cordierite when zinc is part of the equilibration volume, whereas in a zinc-poor system, biotite and sillimanite consumption produces cordierite without spinel (assemblage 3). 3.3.4. Assemblage 2 (Biotite-absent assemblages) In rocks without significant syn-D2 biotite (assemblage 2), the peak assemblage of garnet-sillimanite-ilmenite (t rutile) breaks down directly to coronas of cordierite + spinel + quartz, which represents a broadly decompressional P-T vector. With a P-T evolution as indicated, intersection of the garnet-sillimanite-cordierite- ilmenite field leads to the formation of cordierite coronas around garnet. For the same conditions, zinc-bearing bulk compositions will potentially develop spinel in addition to cordierite as the stability field of spinel-bearing assemblages moves to lower temperatures and higher pressures. This explains the presence of spinel in some cordierite coronas between garnet and sillimanite and its absence from others. For bulk compositions close to garnet, sillimanite is destabilised across the KFMASHT univariant, and garnet + spinel + cordierite would be the stable 68 Chapter 3, Jetty Peninsula metamorphism assemblage. For bulk compositions closer to sillimanite, garnet reacts out, and the assemblage sillimanite-cordierite-spinel becomes stable. 3.3.5 P-T estimates for Jetty Peninsula Textural complexity and extensive overprinting by reaction textures make it difficult to identify an equilibrated assemblage reflecting M1 metamorphic conditions for quantitative P-T estimates. Sample 964-41contains only weakly zoned garnets, lacks an 52 foliation and contains domains in which there is little visible effect of syn- D2 metamorphism. Therefore it was chosen as being most representative of peak conditions. Calculations using the core composition of garnet (almandine component), rutile, sillimanite, ilmenite and quartz, with a¿¡¡ = 0.38 (Berman, 1990) and ai61 = 0.96, using the internally-consistent data set of Holland and Powell, (1990), give pressures of 7 .l - 7.9 + 1 kbar over the temperature interval 750-850oC (Hand et al., 1994). Core compositions of primary garnet and orthopyroxene, assumed to have been in equilibrium in sample 665 yield pressures in the range 6.5- 7.5kbar (Hand et al., 1994) using a variety of garnet-orthopyroxene barometers (wood, 1974; Harley and Green, 1982; Harley, 1984). These estimates probably represent a minimum since the resorption of the orthopyroxene by 52 biotite will have the effect of increasing Kp6e-vg¡ between garnet and orthopyroxene, which has been shown to lower pressure estimates via this approach (Harley, 1984). Some indication of peak metamorphic temperatures may be gained from Fe-Mn compositions of coexisting primary ilmenite and garnet (Pownceby et al., 1987). Average core compositions of garnets in sample 904567 together with ilmenite inclusions give a temperature of 775 + 50oC. Since the minerals will continue to equilibrate with falling temperature, this temperature is likely to be a minimum estimate. These p-T estimates are very common for granulites in general (Harley, 1989) and compare with 6-7kbu,700-800"C for elsewhere in the northem Prince Charles Mountains (Buick et al., 1990; Fitzsimons and rhost, 1992: Fitzsimons and Harley, 1994) and 7kbar, 69 chapter 3, Jetty Peninsula metamorphism 800oC for the Reinbolt Hills directly to the east across the Lambert Glacier, (Nichols and Berry, 1991). Assemblages within the syn-D2 reaction textures are well-preserved and contain phases that allow successful application of the average pressure-temperature approach described by Powell and Holland (1988), using the internally consistent data set of Holland and Powell, (1990). Taken together, average pressure calculations on the M2 assemblages sillimanite-quartz-cordierite-spinel and quartz-spinel-garnet- cordierite in samples 964-67,964-41and 964-83 give 5.9 +0.4 kbar ar 700oC and aH2O = 0.1. Pressure estimates using the corona assemblage garnet-orthopyroxene- cordierite-plagioclase-quartz in sample 964-665 (Table 4) give 5.7 + 0.4 kbar. The pressures results increase slightly with increasing temperature and water activity. The cordierite-spinel-quartz-sillimanite barometer of Nichols et al. (1992) takes into account Zn-beanng spinel compositions. Results of this ba¡ometer using cordierite- spinel assemblages gives a maximum pressure of 5.7 t 0.3 kbar and a minimum pressure of 4.3 + 0.7 kbar at 700oC. Although these pressures seem somewhat low, the difference in pressure from the peak assemblàges to the overprinting coronas and symplectites is consistent with the relative sense of change derived from Fig. 3.3. 3.4 Discussion The metamorphic evolution of the metapelites of Jetty Peninsula can not be easily summarised as a simple decompressional or cooling-dominated p-T path from peak conditions. The stabilisation of biotite-bearing assemblages between the peak and cordieriæ-bearing assemblages can be most easily interpreted as a period of cooling following peak conditions, with later heating during D2. However this leaves three main possibilities: 1) The region underwent two metamorphic events, M1 and M2, which reflect two separate tectonothermal events. In this case, the retrograde path from peak conditions 70 Chapter 3, Jetty Peninsula metamorphism (Mr) is unconstrained, and after a possible return to the geotherm, a heating path ted to peak M2 conditions at lower pressures than peak M1 condirions (Fig. 3.5(l)). 2) The entire metamorphic evolution occurred within a single event, with peak metamorphic conditions being followed by a period of cooling and decompression, then being followed by a second thermal event at lower pressure, resulting in the development of cordierite coronas (Fig. 3.5(2)). 3) Fluctuations in other factors, particularly aH2O, may have increased the stability field of biotite-bearing assemblages, resulting in the stabilisation of biotite-sillimanite during near-isothermal decompression. The stabilisation of biotite in an originally almost anhydrous assemblage suggests that water activity may have increased following M1, and this may well be due to crystallisation of partial melts within the rock. The role of water activity in the stabilisation of biotite-bearing assemblages in Jetty Peninsula and neighbouring regions, and its implication for the P-T evolution of the terrain is discussed in detail in the following chapter. A number of factors, particularly geochronological and structural evidence, must be examined to determine the most likely of these possibilities. I) Geochronology As described in Chapter 2, geochronological evidence from the NPCM and elsewhere in MacRobertson Land suggests a single tectonothermal event at 1000-920 Ma (Tingey, 1982; Young and Black, l99l; Manton et a1., 1992; this study), with no evidence for the granulite metamorphism at 500 Ma which has recently been documented in parts of the Prydz Bay coastline (Ren et al., 1992; Zhao et al., 1992; Dirks et al., 1993). In particular, the dating of syn-D2 intrusions at c.940 Ma (Manton et aL., 1992) and c.939 Ma (this study, see section 2.4), supports the notion that D2 and the accompanying granulite metamorphism occuned as part of the same event as peak metamorphism during D1. Therefore, on the available evidence, scenario (2), in which there is only a single metamorphic event, appears to be the most likely, although more precise geochronology is required before a polycyclic evolution can be entirely discounted. 7t Chapter 3; Jetry Peninsula metamorphism (1) 9 I M 1 7 É (ú Ë -o s -:¿ fL 6 M2 5 4 400 500 600 700 800 Toc M1 assemblage M2 cordierite-bearing coronas and symplectites (2) I I ú L 7 (ú t t t¿o fL 6 5 4 400 500 600 700 800 Toc Fig. 3.5 P-T diagram showing the two scenarios for the P-T evolution of Jetty Peninsula as discussed in section 3.4. (1) Two separate events, Ml and M2, with peak M2 conditions at lower pressure than peak Ml conditions. (2) A single event, with the syn-D2 cordierite-bearing assemblages reflecting a thermal pulse during a decompressional history. Discussion in text. 72 Chapter 3, Jetty Península metamorphism 2) Structural evidence The correlation of the peak metamorphism with D1 and the development of the biotite-bearing and cordierite-bearing assemblages with D2 is well established (see section 2.3). Significantly, both of these deformations had fabrics with an east plunging lineation, which is characteristic of 1000 Ma fabrics throughout MacRobertson Land and Prydz Bay (Dirks et al., 1993, M. Hand, 1993, pers. comm.). The coaxial nature of the deformations associated with the peak and retrograde metamorphism supports, but is not sufficient to demonstrate, the notion that the metamorphic evolution was all related to the same event. On the basis of the existing geochronological evidence, it seems that the entire high grade metamorphic evolution of Jetty Peninsula occurred within a single Mid Proterozoic event. During this event, between 1000 and 940 Ma, M1 metamorphism of c.800oC and 7 -7 .5 kbars was followed by an overall decompressional history, with a second thermal event occurring at lower pressures resulting in the development of M2 biotite- and cordierite-bearing assemblages. 73 Chapter 4, Trost Rocks Chapter 4 Metamorphic evolution of pelites from Trost Rocks, Single Island, Amery Ice Shelf 4,1 Introduction In the previous chapter, textural evidence from the metapelites of Jetty Peninsula was presented to highlight the apparent complexity in the pressure-temperature evolution of the region. Significantly, the region along the western margin of the Amery Ice Shelf appears to form a transitional zone between the Prydz Bay region, which has been widely interpreted to have undergone near-isothermal decompression (Harley and Hensen, 1990) and regions to the west, which have inferred isobaric cooling histories (e.g. clarke et a1., 1989; Fitzsimons and rhosr, rgg2) (Fig. a.1). It is rherefore important to look at other exposed fragments of this part of the tenain, to test whether the apparently complex evolution of Jetty Peninsula is similarly preserved elsewhere. In this chapter, very fine-grained reaction textures in pelites from Trost Rocks, a remote rock exposure on Single Island in the Amery Ice Shelf, will be described in order both to help constrain the metamorphic evolution of this previously undescribed fragment of the East Antarctic granulite terrain, and to examine the relevance of this evolution in the context of the surrounding tenain. 4.2 Regional setting and geological relationships Single Island is a large, almost entirely ice-covered island in the Amery Ice Shelf, approximately 100 km northeast of the northern Prince Charles Mountains, and approximately 200 km west of the Prydz Bay coastline (Fig.4,1). It is roughly circular with a diameter of around 30 km, but has only two small exposed outcrops of rock; Dodson Rocks, on the southeastern corner of the island, and Trost Rocks, which 74 Chapter 4, Trosr Rocks Mawson 10 I 1,2 200 km 6 Cape Bruce 789 N 1 MAWSON I 1 3 I 6 6 4 4 78 VIS TROST Rauer 789 Depot Prydz . ROCKS Group Peak I Amery Bay Hills lce Shelf a 1 Single I lsland 1 7 6 LarsemannB 4 - To Hills 6 o a- :t o 6 Reinbolt Hills 4 7 89 ,a at=. NORTHERtt tr qt PRINCE CHARLES I MOUNTAINS I 6 t- \ / 4 !d 789 I / r¡ / !'ig. a.1- Location of Trost Rocks, on the northeast€rn corner of Single Island in the Amery Ice Shelf. P-T paths constructed for su¡rounding terrains show that isobaric coõling paths predominate o the west, w-hilst paths have been -decompression-dominated inferred for regions tothe easi. / White and Clarke, 1992;2 Cla¡ke et al, 1988; 3 Clarke and Norman, lt 93; 4 Fitzslmons and Thost, 1992; 5 Fiøsimons ædllarley, 1993;6l{arley, 1988;7 StüweandPowell,1989; 8NicholsandBerry,1991. 75 Chapter 4, Trost Rocks r .{i- b Fig.4.2 Two views of Trost Rocks, Amery Ice Shelf (a) Aerial view of Trost Rocks, looking southwest from a distance of 1 km. Ice dome of Single ls_land rising in the background. Cliff-face is a¡ound 100 m high. A smaller second cliff face occurs off the left hand side of the photo. Arrow indicates location of helicopter in (b). (b) Pholo taken looking northwest along the cliff-face at Trost Rocks, with the Amery Ice Shelf behind. Accessible outcrop is extremely limited on the clifftop. Helicopter ior scale. 76 Chapter 4, Trost Rocks comprises two 50-70 metre high sheer cliff-faces, separated by approximately 1 kilometre on the northeastern edge of the island (Fig. a.Ð. This chapter deals with the sequence at Trost Rocks, which consists of layered granulite facies supracrustals and syn-tectonic felsic intrusions, with late pegmatite intrusions. The layered supracrustals include pelitic, biotite-garnet-spinel, garnerorthopyroxene, biotite-garnet, and biotite- orthopyroxene layers. This chapter involves the descrþtion of rocks collected during a brief helicopter reconnaissance, and the restricted time and limited and inaccessible outcrop did not allow detailed structural analysis, although it was observed that the dominant gneissic foliation is vertical, and uends at 0200. No geochronological data exists for Trost Rocks, but its central location within what is considered to be a broad Mid Proterozoic (c.1000 Ma) metamorphic terrain (Tingey, L972: Sheraton et al, 1984; Manton et al., 1992; Kinny et al, 1993) and its general similarity to other supracrustals within that terrain supports the assumption that it is of Mid Proterozoíc age. However, recent geochronology from the Prydz Bay region (Ren et aI., 1992; Zhao et al., L992) has identified the possibility of high grade metarnorphism at c.500 Ma in the Larsemann Hills, and therefore it remains possible that at least some of the granuliæ metamorphism at Trost Rocks may have occurred at this time. 4.3 Petrography The petrography of the Trost Rocks metapelites is complex, and the reaction textures are extremely fine-grained, to the extent that back-scattered electron images on the microprobe were the only way of viewing many of the textural relations between the minerals. All pelitic assemblages at Trost Rocks are characteizedby a relatively coarse grained peak assemblage dominated by garnet and K-feldspar, with minor spinel and ilmenite. The garnets are usually approximately 0.2 - 1.0 cm in diameter and contain inclusions of biotite, spinel and ilmenite. In sample 964TR4, corundum also forms part of the peak assemblage, whilst in samples 964TR8 and 964TR11 the peak assemblage contains sillimanite. This coarse assemblage has extensively reacted to a series of overprinting assemblages defining the retrograde thermal and baric history. 77 Chapter 4, Trost Rocks Rock samet soinel biotite sill Ksoa¡ olas corund ilm oüd ouilþ. TRl. 3 PRz IRz IRz -PR2 I-IP TR4 PRz IRrzo IRrz R3 PRs R2 P IRs R4 IRs TR8 PRz IRzo IRz PR¡ PRs PRz IRs R4 P TR 11 PRz IRze IRr PR¡ PRs P IRs R4 P Table 4.1 Assemblages in semi-pelitic (TRl & 3), and pelitic (TR4, 8, 11) lithologies from Trost Rocks; P = pd of the coa¡se 'peak'assemblage; I = inclusions within garnet; I = rinming corundum; 2 = consuming garnet and K-feldspar; 3 = associated with late M2 assemblage overgrowing biotite-spinel assemblage and rirming ilmenite; 4 = conSuming M1 garnet and M2 sillimanite and biotite; 5 = forms symplectites with cordierite between M2 biotite and M1 garnet; 6 = fonns symplectites with cordierite around M2 sillimanite proximal to M1 garnet; 7 = srnall late garnets overgrowing spinel and cordierite; 8 = associated with late garnet overgrowths. 4.3.1 Corundum-absent pelites In samples 964TR8 and 964TRl1 the infened peak assemblage is garnet-spinel- sillimanite-Kfeldspar-plagioclase-quartz with accessory rutile, and with abundant biotite, spinel and ilmenite inclusions within gamet. This assemblage is overgrown by biotite and spinel, which define an Sz fabric enveloping and consuming primary garnet, spinel and K-feldspar. Primary garnets are extended along the L2 lineation and have developed fractures perpendicular to L2 which are infilled with biotite + spinel (Fig.4.3a). Biotite and spinel also occur along boundaries between K-feldspar grains, particularly proximal to garnet. The biotite-spinel assemblage is overgrown by sillimanite, which often exists as coronas around the secondary spinel and ilmenite, or as intergrowths with secondary biotite (Fig. 4.3(c,d)). This biotite-sillimanite assemblage is best developed where it envelops primary garnet. Garnet is almost always separated from the secondary biotite, sillimanite and spinel by cordierite-bearing coronas and symplectites. The most common of these textures are symplectites of cordierite and K-feldspar between garnet and biotite (Figs. 4.3(b), 4.4(a,b)). Cordierite also forms coronas between primary garnet and secondary sillimanite, and growth of small vermicular spinel is locally developed around sillimaniæ within these coronas (Fig. 4.4(b,c)). W'here biotite is proximal to sillimanite, it is also 78 Chapter 4, Trost Rocks ij ¡i.'s Fig. 4.3 Reaction textures in pelites from Trost Rocks. V/idth of field of view is 2 mm. (a) Garnet pulled apart along the L2 lineation, with the pull-aparts infilled with biotite and spinel. Sample TR4. (b) SZ biotite and spinel in garnet pull-apart with cordierite-kspar symplectites (sym) separating the primary garnet from the secondary biotite. Sample TR4. (c) Garnet consumed by biotite and spinel, with sillimanite forming coronas around ilmenite and spinel and sillimanite fibres overgrowing 52 biotite. Arrow indicates small euhedral garnets overgrowing primary garnet and secondary spinel. Sample TR8. 79 Chapter 4, Trost Rocks å Fig. 4.3 (continued) Reaction textures in pelites from Trost Rocks. (d) Primary garnets consumed by biotite, spinel and sillimanite, with cordierite consuming biotite, garnet and sillimanite. Sample TR8. Width of field of view is 2 mm. (e) Primary sillimanite consumed by 52 biotite and then by extremely fine grained spinel- cordierite symplectites. Arrow indicates late garnet overgrowing biotite. Sample TRl1. Width of field of view is 2 mm. (f) Late small garnets (gt2)overgrowing corroded primary garnet (gtl) and secondary spinel. Trost Rocks, sample TR8. V/idth of field of view is 1 mm. 80 Chapter 4, Trost Rocks gt gt (a) (b) Fig.4.4. Back-scattered electron images of reaction textures from Trost Rocks. Scale ba¡ is 100¡tm. (a) Symplectite of cordierite and K-feldspar due to a reaction be¡veen biotite and primary garnet. Sample TR8. (b) Reaction between primary garnet and 52 sillimanite and biotite, with cordierite - K- feldspar symplectites between biotite and garnet, and cordierite-spinel symplectites around sillimanite. Sample TR8. (c) Cordierite consuming biotite, sillimanite and garnet with vermicular intergrowths of spinel around sillimanite, and K-felspar a¡ound garnet adjacent to biotite. Note also the consumption of 52 biotite by spinel. Sample TR8. (c) 81 Chapter 4, Trost Rocks I (d) (e) Fig. 4.4.(cont.) Back-scattered electron images of reaction textures from Trost Rocks. Scale bar is 100Pm. (d) Spinel and biotite extensively consumed cor by cordierite, with ilmenite also being a product. Cordierite also separates corundum from K-feldspar. SamPle TR4. (e) Symplectite of cordierite and K-feldspar between biotite and garnet, with spinel also consuming biotite. SamPle TR4. (Ð Corundum, initially partially consumed by spinel and biotite, and then enveloped by a narrow corona of cordierite. Sample TR4. (Ð 82 Chapter 4, Trost Rocks often rimmed by spinel (Fig. 4.4(c,e)). Ilmenite is often also present in the cordierite- bearing coronas and symplectites (Fig. 4.4(d)). Small, almost euhedral garnets, no more than 0.05 mm in diameter occasionally occur clustered around the conoded primary gamets, and often overgrowing spinel and the cordierite coronas (Figs 4.3(e,Ð). Where late garnet overgrows primary garnet, it often forms intergrowths with qvaÍtz, whilst when it grows within cordierite it forms clusters of euhedral grains. 4.3.2 Corundum.bearing pelites Sample 964TR4 is similar to the peliæs described above, with the exception that corundum exists instead of sillimanite as part of the peak assemblage, and in places is still in contact with primary garnet. Corundum also appears to have initially been in contact with primary spinel. Corundum, in the presence of K-feldspar and garnet, is consumed by secondary spinel and, to a lesser extent, by biotite (Fig. a.aÐ. As in the other pelites a biotite-sillimanite assemblage overgrows the biotite-spinel assemblage, although sillimanite is never found adjacent to corundum. Exûemely narrow coronas of cordierite separate corundum from K-feldspar (Fig. .aÐ. 4.3.3 Semi-pelitic assemblages In semi-pelitic lithologies on Trost Rocks, the infened peak assemblage is garnet - K-feldspar - quartz - ilmenite, with abundant spinel, ilmenite and biotite inclusions within the garnets. This assemblage is enveloped by a strong fabric defined by biotite, spinel and ilmenite, which consume primary garnet and K-feldspar. The biotite is significantly more abundant in the semi-pelites than in the pelites, and the secondary spinel contains abundant exsolution of magnetite. As in the pelites, the garnets are extended along theL2lineation and these pull-aparu are infilled by biotite and spinel. Biotite and spinel also form along K-feldspar grain boundaries. Small, euhedral gamets overgrow these biotite-spinel fabrics, and also form around the margins of primary 83 Chapter 4, Trost Rocks garnets. 4.3.4 Mineral Chemistry Mineral analyses were obtained using a JEOL 733 microprobe at the Centre for Electron Microscopy and Microstructural Analysis at the University of Adelaide, using the energy-dispersive spectra (EDS) system with an accelerating voltage of 20 keV. Representative analyses of minerals are given in Appendix 3. Garnets are consistently almandine-rich with Xp. generally between 0.6 and 0.7, and Xotr between 0.25 and 0.32, whilst grossular and spessartine components generally total less than 0.08. In most cases there is little compositional zonation within the garnets, except within 0.1-0.2mm of the rim, where, at contacts with ferromagnesian phases, there is an increase in the Xps and corresponding decrease in XMg, which is probably a function of reequilibration during cooling. However, in places, what appear to be late overgrowths have formed on the rims of these garnets, with slightly higher Xp. values than the large M1 garnets. Small, apparently late garnets overgrowing Ml garnets and consuming biotite and spinel also have high Xps of up to o.76. Biotite at Trost Rocks generally has an Xps of between 0.45 and 0.54, although one biotite in the secondary spinel-biotite assemblage in the semi-pelites had an Xps of only 0.32. Primary biotites, which are found as inclusions within garnet, generally have slightly lower X¡s than M2 biotites. TiO2 content of the biotites is variable, generally ranging between 0.25 and 0.4 atoms per formula unit (assuming 11 oxygens), and is generally inversely proportional to the Mg2+ content. The correlation of Ti-rich biotites with Fe-rich biotites is not unexpecúed as the source of the Ti is presumably the consumption of ilmenite. Spinel from pelites at Trost Rocks is predominantly hercynite-spinel solid 84 Chapter 4, Trost Rocks solution, with only minor components of chromite and magnetite, and minimal to negligible Mn,Zn and Ti substitution. The Xps of the spinels is generally between 0.68 and 0.74, with between 0.40 and 2.0 wt%o Cr2o3. In the semi-pelites (e.g. 964TRl), there is a significant magnetite component in the spinel, wittr around 2.5mo\7o magnetite as well as magnetite exsolution, and a much lower chromite component. Cordierite is generally only found as n¿uro\ry coronas and symplectites at Trost Rocks, and has a consistent composition both between and within samples, with an Xps of 0.22-0.24. No consistent compositional zonation was detected within the cordierite coronas, although it was generally difficult to get profiles across coronas due to the higtrly altered nature of most of the cordierite. Feldspar in the Trost Rocks metapelites is almost all K-feldspar, which generally has a composition of around Or9gAb7An3. Plagioclase composition is more variable. Plagioclase occurring in the matrix as a minor constituent with K-feldspar usually has an intermediate composition between anorthite and albite, whilst plagioclase forming from the consumption of garnet usually has a significantly more calcic composition of Angg, to incorporate the grossular component of the garnet. 4.4 Interpretation of reaction textures 4.4.1 Reactions in the KFMASH and KFMASHTO systems The layered gneisses at Trost Rocks are characterised by being rich in oxides, particularly ilmenite and spinel, and contain abundant biotite. Therefore the most appropriate system to use when investigating the metamorphic evolution of these rocks is the KFMASHTo (K2o-Feo-Mgo-Al2o3-Sio2-H2o-Tioz-oùsysrem. Fig. 4.5 shows petrogenetic grids in the KFMASH and KFMASHTO system, adapted after Clarke et al (1989). The topology of these grids requires the assumption that Xp"sp > XFeBt, which is consistently true for coexisting spinel and garnet at Trost Rocks. In the KFMASH grid, the reactions of interest at Trost Rocks are those emanating from the 85 Chapter 4,Trost Rocl KFMASH ! ()) xo +q+ksp+l o dñ q lspl sittI sp x lopxl -o o ô a Þo o, .\t '\-(/) a.\- f, ç¡t KFMASHTO gt sù\ Ø sP +q+ksp+melt .o c), [opx ru] ùro g\ sù\ r..l sQ r þ\ [opx mt] õ E øQ o. a./) þ\ 5 ttl o O) sQ a þ. o, arò øQ Fig. 4.5. Petrogenetic grids in the KFMASII and KFMASIITO systems adapted from Clarke and Powell (1991). The univariant reaction lretween the [opx mt] and [opx ru] invariant points in the KFIVIASFITO grid Íìccounts for the observed cordierite-bearing symplectites at Trost Rocks. 86 Chapter 4, Trost Rocks orthopyroxene-absent ([opx]) invariant point, involving sillimanite, biotite, spinel, garnet and cordierite, with K-feldspar and quartz in excess. Expanding this grid into the KFMASHTO system, the [opx] invariant point becomes two invariant points [opx mt] and [opx ru], joined by the univariant reaction biotite + garnet + sillimanite = spinel + cordierite + ilmenite, which is a typical reaction at Trost Rocks. The most common reaction texture in the Trost Rocks metapelites is the breakdown of the M1 garnet-K-feldspar-ilmeniæ + sillimanite assemblage to spinel and biotite. In the KFMASH system, this reaction could be explained by the (opx cd) reaction sillimaniæ+garnet+K-feldspar = 'biotite+spinel (1) The biotiæ and spinel producing reaction therefore is; garnet + K-feldspar + ilmenite + sillimanite = spinel + Ti-biotite (2) In the corundum-bearing pelite 964TR8, corundum is often also consumed by spinel and biotite, through the reaction garnet + corundum + K-feldspar* ilmenite = spinel + biotite (3) Following this, the spinel-biotite assemblage became unstable, and through a continuous reaction the stable assemblage became sillimanite-biotite. Following the development of the biotite-bearing assemblages, was the stabilisation of cordierite-bearing assemblages. These are predominantly cordierite - K- feldspar * ilmenite symplectites between gamet and biotite. In addition, cordierite formed between garnet and sillimanite, with growth of vermicular spinel around the secondary sillimanite and biotite. The reactions which produce these symplectiæs a¡e- garnetl + sillimanite2 = cordieriæ + spinel (4) garnetl + biotite2 = cordierite + K-feldspar + ilmenite (5) When considering the progress of the cordierite-producing reactions in the KFMASIIT system, the univariant (opx mt ru) reaction: biotite + garnet + sillimanite = spinel + cordierite + ilmenite + K-feldspar (6) 87 Chapter 4, Trost Rocks adequately accounts for all of the observed textures, including the consumption of biotite by spinel (Figs. 4.4c,e). This reaction occurs with heating and/or decompression, although it important to note that this is not the breakdown of a stable biotite-garnet-sillimanite assemblage, but the reaction between primary garnet and a later biotite-sillimanite assemblage. The growth of late euhedral garnets overgrowing spinel, cordierite and primary garnet, is of significant importance in the interpretation of the overall P-T evolution of Trost Rocks. It can be seen on both the KFMASH and KFMASHTO grids that all reactions which produce garnet at the expense of cordierite require cooling and/or compression, and this precludes a decompression-dominated history following the stabilisation of cordierite. The P-T evolution of Trost Rocks is examined further in section 4.4.4. 4.4.2 Correlations with Jetty Peninsula Any interpretation of the significance of the metamorphic textures observed at Trost Rocks requires conelations with neighbouring fragments of the tenain for which more detailed structural observations and geochronological data exist. The peak metamorphic assemblage at Trost Rocks presumably represents the same peak metamorphic event throughout MacRobertson Land and the Prydz Bay coast, which occurred around 1000 Ma (Tingey, 1972; Sheraton et al, 1984; Kinny et al., 1993; Manton et al., 1992), and is directly correlated with peak metamorphism on Jetty Peninsula. On Jetty Peninsula, the primary assemblage was overprinted by a biotite fabric, which was subsequently overprinted by cordierite-bearing assemblages. The secondary assemblages on Trost Rocks are also dominated by an early biotite fabric which is overprinted by cordierite-bearing assemblages, and therefore the biotite-rich and cordierite-rich assemblages on Trost Rocks are correlated with the similar assemblages on Jetty Peninsula. On Jetty Peninsula, the evidence for the stability of sillimanite in an assemblage with biotite, was scarce, and relied mainly on the existence 88 Chapter 4, Trost Rocks of biotite within dynamically recrystallised sillimanite, whilst there are abundant intergrowths of biotite and sillimanite at Trost Rocks, which clearly define a stable assemblage overgrowing earlier biotite-spinel assemblages. On Trost Rocks, there is no convincing evidence that the biotite-bearing and cordierite-bearing assemblages form part of the same structural-metamorphic event. However, on Jetty Peninsula there was convincing structural evidence that they defined the same fabric (see Chapter 2), and therefore it can reasonably be assumed that the development of the retrograde assemblages on Trost Rocks also formed part of the same event. Similarly, the late garnet growth at Trost Rocks is conelated with the late garnets on Jetty Peninsula. 4.4.3 The role of øIJ2O in the stabilisation of biotite at Trost Rocks and Jetty Peninsula Fig. 4.6 is a pseudosection in the KFMASHTO system for pelitic compositions similar to those at Trost Rocks, adapted from Hand et. al. (1994). It is clear from this pseudosection that the overprinting of the original garnet-sillimanitelrutile assemblage by cordierite + spinel can occur with near-isothermal decompression. However, the intermediate biotite-spinel and biotite-sillimanite assemblages require a more complex metamorphic evolution. As discussed in section 3.5, the varying scenarios which could account for the occunence of these biotiæ-bearing assemblages are: 1) A period of cooling following peak conditions, followed by a second thermal event with the development of cordierite-bearing assemblages, or 2) An increase in the activity of H2O, at relatively constant temperature. To determine which scenario is more likely to account for the stabilisation of biotite at Trost Rocks and Jetty Peninsula, it is necessary to consider the processes which control a(HzO) during granulite facies metamorphism, and in particular the role of panial melts. This is particularly important because partial melts are usually present in pelitic rocks from this region which have developed a biotite fabric. 89 Chapter 4, Trost Rocks b cd b¡ gt gt V L v bi L cd v L L gt v cd L v gt cd v KFMASH + qz + ksp +sill bi g L I gcdL bi Lv L ,1, cdL P bi v cdL V T Fig.4.6 (a) P-T diagram adapted from Ellis (1986), showing a schematic representation of partial melting reactions in the KFMASH system, using the assumption that Mgliquid>garnet. (b) P-T pseudosection using the same projection as (a) for an intermediate XFe. Arrows indicate the direction in which the proportion of melt will increase across a divariant field. (Adapted from Powell and Downes, 1990) 90 Chapter 4, Trost Rocks In relatively anhydrous granulite facies rocks, the most plausible way to account for an increase in a(HzO) is the crystallisation of a melt. As such, it is tempting to account for the stabilisation of biotite following Mt by invoking an increase in a(H2O) due to crystallisation of partial melts in the rock. Figure 4.6(a) is a P-T grid for partiat melting in the KFMASH system adapted from Ellis (1986), and Figure 4.6(b) is a P-T pseudosection based on this grid by Powell and Downes (1990). The arrows in Figure 4.6(b) represent the direction in which melt proportion increases across the divariant fields. For an initial assemblage in the garnet-bearing trivariant field, it can be seen that two divariant reactions control the proportion of melt in the rock. A decrease in pressure will result in an increase in the proportion of melt in the rock, whilst a decrease in temperature will result in crystallisation of melt. It therefore seems inconsistent to suggest that a rock could undergo an increase in a(H2O) due to crystallisation of melts during near-isothermal decompression. A more plausible explanation is that the rock must undergo cooling in order to get crystallisation of the melt phase and a resultant increase in a(H2O). This suggests that, although an increase in a(HzO) may be required for the stabilisation of biotite, it is likely that cooling following M1 would be required to account for this increase. The notion that biotite production is related to the crystallisation of partial melts is supported by the observation that more anhydrous pelitic assemblages such as 964-83 on Jetty Peninsula, which do not contain leucosomes, did not develop a biotite fabric following M1 metamorphism. These rocks may represent residues from partial melting or alternatively may never have undergone melting. In either case, ttris suggests that the removal of melt from the rock was a crucial factor in determining whether different rock types stabilised biotite following cooling from the metamorphic peak. The recognition that melt crystallisation is a key factor in biotite formation implies that some degree of cooling must have occurred following M1 to enable partial melts to crystallise. 9L Chapter 4, Trost Rocks 4.4.4 P-T evolution Pressure-temperature estimates for the pelitic assemblages at Trost Rocks were made using the average pressure-temperature approach of Powell and Holland (1985, 1988) and using the internally consistent dataset of Holland and Powell (1990). The complexity of the metamorphic textures made it difficult at times to identify phases which were in equilibrium, and the very fine grained nature of the secondary phases introduced doubts regarding possible re-equilibration following crystallisation. The X2 test within the THERMOCALC program of Powell and Holland (1988), allowed the identification and removal of "suspect" phases from the calculations, but even so, the pressure-temperature estimates for Trost Rocks remain poorly constrained. THERMOCALC output for the Trost Rocks metapelites is shown in Appendix 4. For the M1 assemblage, average pressure calculations consistently had best statistics for temperatures of approximately 750-800oC, which is generally consistent with peak temperatures in neighbouring regions in the NPCM. V/ithin this temperature range, average pressure estimates range from 6.0-6.3 t 0.5 for the infened peak assemblage garnet-sillimanite-quartz-ilmenite-rutile-spinel-plagioclase in sample 964TR8. For the corundum-bearing M1 assemblage in 964TR4, pressure estimates of 6.8-7.0 + 0.5 kbar were obtained. For the overprinting assemblages, pressure- temperature estimates are less well constrained, due to the extremely fine grained nature of the reaction textures. However, the KFMASHTO pseudo-section in Fíg.4.7 shows the apparent relative changes in pressure, which suggest lower pressures than peak conditions, probably also at lower temperatures, similar to the overprinting syn-D2 assemblage on Jeny Peninsula. Pressure-temperature estimates were also obtained using the experimental results of Nichols et al (L992) for the KFMASHZn system. For the M2 cordierite-spinel- sillimanite-quartz assemblage in sample 964TR11, pressures of 5.3-6.5 kbars were obtained in the temperature range 600-700oC, assuming low a(H2O), with the 92 Chapter 4, Trost Rocks KFMASHTO + qz + ilm + ksp + melt (o ? Eu, ¿ gt sill cd ru sill cd cdspru gt sill cd sp sill cd sp ru A T sill fro ru ru F M M2 cordierite-bearing M1 metamorphism coronas and symplectites 93 Chapter 4, Trost Rocks geothermometer giving what would appear to be unrealistically low temperatures of 603-622'C over the interval4-8 kbars. The inaccuracies with this method may be partly due to the scarcity of quartz in contact with sillimanite, spinel and cordierite. An interpretation of the thermal and baric evolution of Trost Rocks is shown by the arrows on the KFMASHTO pseudosection in Fig. 4.7. In summary, Ml metamorphic conditions appear to be in the region of 750-800oC and 6-7 kbars. The immediate retrograde evolution following M1 is not constrained, but stabilisation of biotite-bearing assemblages probably occurred at conditions around l00oC and I-2 kbars below Mt conditions, and were followed by heating, possibly with some accompanying compression, to a second thermal peak during which cordierite was stable. No reliable evidence exists at Trost Rocks for the evolution following cordierite- corona development, although late garnet growth may record some degree of near- isobaric cooling or cooling accompanied by compression, and would seem to preclude a decompression- dominated evolution following syn-D2 metamorphism. 4.5 Discussion The reaction textures preserved in the metapelites of Trost Rocks and Jetty Peninsula reflect broadly simila¡ metamorphic histories, for which two heating events at different pressures appears to be the most likely interpretation. It is quite possible that these could reflect thermal pulses in a decompression-dominated history from peak conditions during the c.1000 Ma event, but the lack of an ísothermøl decompression path removes any time constraint which would require the peak and cordierite-bearing assemblages to represent part of the one event. If the decompression could be shown to have occurred isothermally, then it would be necessary for the time span between M1 and M2 to be relatively short, due the probable transient nature of thermal perturbations within the mid crust (Sandiford and Powell, 1991). If this was the case, a 'clockwise' P-T path such as has been inferred for Prydz Bay (e.g. Harley and Hensen, 1990) could be assumed to be appropriate. However, if cooling occurred between the two events, 94 Chapter 4, Trost Rocks then there is no constraint on the time span, and decompression need not have been rapid. Clearly, when viewing the reaction textures at Trost Rocks in isolation, they can reveal little about the true metamorphic evolution of the region due to the uncertainty of the timing of the development of the overprinting assemblages. This highlights the necessity of integrating structural and geochronological information with metamorphic petrology, and only through correlations with neighbouring regions can an understanding of the nature of the pressure-temperature evolution be obtained. The metamorphic evolution of Trost Rocks will be further correlated and integrated with other fragments of the terrain in following chapters. 95 Clnpter 5, Wollastonite- scapolite calc - silicate s Chapter 5 Wollastonite-scapolite calc-silicates as indicators of the P-T-fluid history in MacRobertson Land 5.1 Introduction Our understanding of the P-T-t evolution of granulite terrains has been based largely on the interpretation of reaction textures in pelitic, and less commonly, mafic and felsic granulites. Whilst often providing good constraints on pressure-temperature conditions, textures in pelitic rocks frequently reveal little about the role that fluids played in the evolution of the terrain, and in particular the role of CO2-rich fluids in granulite terrains. In comparison, calc-silicate assemblages have largely been neglected in studies of the evolution of granulite terrains, although recently several authors (Schenk, 1984; Moecher and Essene, 1990a; Harley and Buick, L992; Fitzsimons and Harley,1994) have demonstrated the potential of these rocks in helping to constrain the P-T-fluid history of granulites. These studies have shown that diagnostic reaction textures involving wollastonite and scapolite are relatively common, and can provide important additional consEaints on the retrograde pressure, temperature and fluid evolution of high grade terrains. The use of calc-silicates in P-T-fluid studies has been largely made possible by recent developments in understanding the phase relationships in the CaO-Al2O3- SiO2-vapour (CASV) system, particularly through recent studies by Harley and Buick (1992). Perhaps the best examples of the use of phase relations in the CASV system in constraining the P-T-fluid evolution of wollastonite-scapolite 96 Clnpte r 5, W ollastonite - scapolite calc-s ilicate s calc-silicates are studies by Fitzsimons and Harley (1994), who have inferred a near-isobaric cooling path for the Nemesis Glacier region of the northern Prince Cha¡les Mountains on the basis of grossular-producing reactions in wollastonite- scapolite calc-silicates, and Harley and Buick (1992), who inferred a decompressional P-T path for the Rauer Islands on the basis of wollastonite- producing reaction textures. These studies have shown that wollastonite-scapolite calc-silicates are potentially as useful as pelites in constraining granulite P-T paths, whilst having the additional benefit of constraining the fluid composition during metamorphism. In addition to providing constraints on P-T-fluid paths for localised terrains, studies of wollastonite-scapolite calc-silicates are also important in addressing more general problems relating to the role of fluids in the generation of granulites. Many authors (e.g. Newton, 1986; Santosh and Yoshida, 1992) have proposed that granulite formation results from pervasive streaming of COz-rich fluids through the terrain; an interpretation based largely on CO2-rich fluid inclusions in granulites and incipient charnockitisation in southern Indian and Sri Lankan granuliæs. Although it has recently become clea¡ that many granulites have formed in fluid-absent or internally-buffered Co2-poor environments (e.g. Lamb and Valley, 1988; Moecher and Essene, 1990b), the role of the fluid phase in granulite metamorphism remains inadequately understood. The ability of calc- silicate assemblages to constrain a(COz) during metÍrmorphism therefore makes them important tools in our attempts to understand the role that fluids play in granulite formation. This chapter deals with reaction textures in calc-silicates from a number of localities in the northern Prince Charles Mountains (NPCM) in order both to provide further constraints on the metamorphic history of the region, and specifically to provide additional data on the role that the fluid phase plays in the 97 Chapte r 5, Wo llas tonit e -s c apolite c alc - s i lic ate s 500E 600E Complex Mawson f Co mp lex Prv6, Bay r Grou Amery lce Shelf Northeln Prlnce ' Larsemann Gharles Mountalns /l-, Hills Reinbolt Hills 200km . show_ing the_location of the three calc-silicate localities Iig. lf -Map described in Chapter 5: Else Platform and Mt. McCarthy, northern Prince Cha¡les Mountains, and Rumdoodle i-n the Framnes Mountains. The location of the Nemesis Glacier region of the NPCM, where Fitzsimons & Harley (1993) also described wollastonite-scapolite calc-silicates, is also shown. 98 Clnpter 5, Wollastonite- scapo lite c alc-silicate s evolution of granulites. Occunences of rare barium-aluminium silicates from these calc-silicates are also described. Apart from their 'curio' value, these rare assemblages also place further constraints on the fluid history of these rocks. Finally, calc-silicates from the Framnes Mountains, on the Mawson Coast of MacRobertson Land, both as a basis of compæison with the NPCM, and also to place further constraints on the evolution of the terrain. Location of regions discussed in this chapter are shown in Fig. 5.1. 5.2 Calc-silicate gneisses on Else Platform, Jetty Peninsula 5.2.1 Geological setting Calc-silicate gneisses occur on the northern part of Else Platform and generally form lenses within pyroxenites or are interlayered with pyroxenites on the scale of only a few metres. Detailed locations of calc-silicate localities on Else Platform are shown in the geological map of Else platform (Fig. 2.3), with the most interesting assemblages being found in the north of the platform. Calc- silicate and pyroxenite lithologies are enclosed within the biotite-orthopyroxene gneiss and are restricted to narrow (1-5m) discontinuous lenses up to 30m long. A grain shape fabric defined by elongate diopside and wollastonite often parallels the banding. Calc-silicates also occur on southern Kamenistaya Platform in the vicinity of Soyuz Base, where they are also interlayered with, and form lenses within pyroxenites. 5.2.2 Petrography All calc-silicates discussed here include scapolite and clinopyroxene as part of the coarse grained fabric which defines the 'peak' assemblage, and all but one contain wollastonite. The peak assemblage is generally granoblastic with equant grains, although in some samples elongate wollastonite and diopside weakly define a fabric. Although peak assemblages are similar between specimens, significant variations exist in the bulk chemistry, resulting in variable retrograde reaction 99 C lnpt e r 5, W ol last onit e - s c ap o lite c alc - s il ic at e s textures. Mineral assemblages are summarised in Table 5.1 In the more calcic samples (964656C, D and 964MH1;Assemblage A), the peak assemblage comprises wollastonite, scapolite, calcite and diopside, with minor titanite and rare hyalophane (in 964656D). This assemblage is characterised by the breakdown of scapolite to fine grained symplectites of anorthite and calcite, with abundant hyalophane and minor quartz (Figs.5.2(a); 5.a(a)). \Vollastonite is often partially consumed by calcite and quartz, particularly along wollastonite cleavage planes (Fig. 5.2(b)). Rims of diopside have also formed on the wollastonite, particularly at contacts with calcite. Assemblage wol scp cDx gr plag cal qz czo titan hval A: 656C&D; MHI PR P PRz -R1 PRrz Rrz R4 P PRr B: MH3, MH64 PPP R3 R1 PRrz¡ R3-P c;628 P P - PRr Rr PR4P Mt. McCarthy PPP R1 Rrz Rrz PRr Table 5.1 Assemblages in calc-silicates from Else Plaúorm and Mt. McCarthy, showing minerals which fonn part of the peak assemblage (P) and retrograde reaction textures (R). 1 = in symplectites replacing scapolite; 2 = consuming wollastonite; 3 = coronas between scapolite, wollasûonite + calcite; 4 = consuming anorthite, calcite and/or diopside. Samples 964MH3 and 964MH64 (Assemblage B), have the same peak assemblage as assemblage A, but with a more sodic bulk composition. In these rocks, narrow rims of grossular garnet occur at contacts of wollastonite and calcite with scapolite (Fig. 5.2 (c)). In places, where the grossular rim is between wollastonite and scapolite, the rims also contain minor q\artz. A second, more grossular-rich generation of garnet forms rims on the original garnet coronas at contacts with scapolite. Breakdown of wollastonite to calcite and quartz is observed to have occuned subsequent to the formation of the garnet coronas. Sample 964628 is wollastonite-absent, with the peak assemblage scapolite- 100 Chnpter 5, Fts. S.2la) Svmplectites of anorthite and calcite (an-cal) with minor quartz, consumrng pãpófit.ìi'rci). ^nrce Platforrn, sample 656D. Width of freld of view is 2 nnm. -r: - î , both on grain by scaPolite 656D. Width 101 Chapte r 5, Wollastonite- scapolite calc-silicates al É Fig.5.2 (c) Rim of grossular garnet__(gr) aryuld scapolite (scp) at contacts with wõtlastonite (wo) anã calcite (¿al). White flecks within the grossular are quartz. Else Platform, sample MH3. rWidth of field of view is lmm. (. :l å Ê I a ,t f; I ' { Fig.5.2 (d) Clinozoisite (zo)-consuming anorthite and cal.cite (an-cc), which_resulted frõm the'earlier breakdown < f scapolite (scp). Diopside (di) remains essentially unreactive. Else Platform, sample 628. Field of view is 2 mm. rcz Clapter 5, Wollastonite-scapolite calc-silicates clinopyroxene-quartz-anorthite, with minor titanite. In this sample, scapolite has extensively reacted to anorthite and calcite, and this secondary anorthite-calcite assemblage is then partially consumed by clinozoisire (Fig. 5.2 (d)). 5.2.3 Mineral Chemistry Mineral analyses were obtained using a JEOL 733 electron microprobe at the Centre for Electron Microscopy and Microstructural Analysis at the University of Adelaide. For most analyses, the energy-dispersive spectra (EDS) system was used, with an accelerating voltage of 20 keV, although wavelength-dispersive spectra (WDS) analysis was performed on scapolites and barium silicates. Representative mineral analyses are given in Appendix 3. Scapolite varies widely in composition between specimens, with equivalent anorthite (EqAn: 100(Al-3)/3 for scapolite normalised to (Al+Si)=12; Orville, 1975; Ellis 1978) contents varying between 70 and 84, corresponding to a range in X-r (mole fraction of meionite) of 0.78 to 0.91. Within specimens, EqAn varies by up to 4, but no consistent zoning is apparent within individual scapolites. The breakdown of scapolite to anorthite-calcite symplectites is clearly dependent on scapolite composition, as only scapolites with a EqAn of 78 or above have passed through this reaction. In some samples, substitution of K for Na has resulted in potassium comprising up to 20Vo of the monovalent cations, explaining why K- bearing feldspars occur in breakdown symplectites after scapolite. WDS analysis has shown SO+, Ba and Cl contents of the scapolites to be negligible. Garnet is never found as part of the primary assemblage in wollastonite-scapoliæ calc-silicates from Else Platform, and only occurs as naffow rims between scapolite, wollastonite and calcite in samples MH3 and P64. These rims are almost pure grossular-andradite solid solutions, with between 65 and 75 mo|Vo grossular. The almandine component varies between I and 7 molVo, with 103 Clnpter 5, Wollastonite- scapo lite calc-silicate s approximately 1 mo|Vo spessartine and negligible pyrope content. An inverse conelation exists between andradite and almandine contents within specimens. No zoning was observed within the rims, although in sample MH3, secondary garnet rims have developed between grossular-andradite and scapolite. These secondary rims have almost pure (98 to 99.5 molØo) grossular compositions, with minor spessartine and almandine. Clinopyroxene compositions show considerable va¡iation between specimens with X6¡ ranging from 0.40 in MHI to 0.84 in 656C, and with variations in X¿¡ of up to 0.1 within specimens. Esseneite ((Mg, Fe)-rSLrAlFs3+¡ substitutions predominate over tschermaks ((Mg, Fe)-tSi-tAlz) substitutions, and generally account for most Al within the pyroxenes. Esseneite substitutions appear to decrease towards the rims of the clinopyroxenes, as also described by Fiøsimons and Harley (1994) in calc-silicates from the Nemesis Glacier region. Plagioclase compositions are invariably extemely anorthite rich with X¿¡¡ varying between 0.96 and 0.99, with negligible orthoclase or celsian components. Compositional variation within hyalophane (Ba-K feldspars) is extreme both within samples and between different samples. Celsian (Ba-feldspar) contents of these hyalophanes range from Cn13 to Cn¿¡ (Cn - Ba / (Ba+K+Na+Ca)), with the remainder dominantly orthoclase. Na content is consistently in the range Abz-s whilst Ca content is extremely variable between Anç1 and 4n13, with a consistent positive correlation between An and Cn components" 5.2,4 lnterpretation of reaction textures and P-T history The reaction textures in calc-silicates from Else Platform can all be understood in terms of relatively simple reactions. The development of anorthite- calcite symplectites after scapolite is consistent with the breakdown of scapolite, viz. to4 Clwpter 5, Wo llastonite - scapo lite calc - silicate s scapolite =3 anorthite +calcite (1) The breakdown of wollastonite to calcite and quartz in most wollastonite-bearing assemblages on Else Plaform is consistent with the progress of the reaction: wollastonite + COz = calcite + quartz (2) The formation of grossular rims between wollastonite, scapolite and calcite in assemblage B can be att¡ibuted to the reaction: scapolite + 3 wollætonite + 2calcite = 3 grossular +3COz (3) Blebs of quartz in grossular coronas between wollastonite and scapolite suggest that a second grossular forming reaction took place. ,scapolite +wollastonite = 3 grossular +2quutz+COz Ø) The development of clinozoisite after anorthite and calcite seen in assemblage C can be an¡ibuted to the hydration reaction: 3 anorthite + calcite + H2O =2chnozoisite + CØ (5) Therefore, all of the reactions can be understood in terms of the CASV (CaO-AlzO3-SiO2-vapour) system, following the approach of Harley and Buick (1992). Relevant petrogenetic grids for the CASV system for granulite facies calc- silicates of similar compositions to those on Else Platform have been constructed by Fitzsimons and Harley (1994) for calc-silicates in the Nemesis Glacier region, and these have been used to constrain the P-T-fluid history of the Else Platform region. Figure 5.3 is an T-a(COù gr'd constructed by Fitzsimons and Harley (L994) for compositions which approximate those in assemblage A. The stability field of the peak wollastonite-scapolite-calcite assemblage suggests that peak conditions in Else Plaform were in excess of 800oC at a pressure of around 7 kbar, which is consistent with P-T estimates from pelitic lithologies. The stability field of this assemblage also constrains a(CO2) to have been greater than about 0.4, although the lack of garnet in assemblage A introduces considerable uncertainty into the relative placing of the grid in T-a(COz) space, as the stability fietd of the assemblage is partially dependent on grossular actvity. Higher grossular activities 105 Chapter 5, Wollas tonite-scapolite calc-silicates (3) o o U) E 2 o ¿ o o ! õ Wo An 850 wo Grs Qtz Wo Scp Grs CalQtz -\ 800 P--. / T \ I YC / ("c) o& / / EqAn = 81 + --- (1) 750 / (2) / h;f / EqAn 78 scp = (1) \ An Cal 700 Else Platform EqAn = 78 Xgrs = 0.76 P=7kbar 0.2 0.3 0.4 0.5 0.6 cI COz Fig. 53. T-aCO2 grid for the CASV system at 7 kbar adapted from Fitzsimons & Harley (1994) showing cooling reactions observed in calc-silicates from Else Platform for compositions approximating that of æsemblage A. T and aCO2 values are only approximate, due to uncertainties in the p,ressure and variations in the activities of grossular and meionite. The dashed line shows the effect on the grid topology of 1 kbar of decompression. The horizontal shading indicates the stability field of assemblage A. Numbers refer to reactions mentioned in the text 106 Clnpter 5, Wollastonite- scapolite calc-silicates push the stability field of the peak assemblage towards higher a(COz) values, by aiding the progress of reaction (3). For an Xgr of 0.76, the wollastonite-scapolite- calcite assemblage would only be stable above an a(CO2) of 0.5 (Fig.5.3), although high andradite contents in garnets in other calc-silicate assemblages from Else Plaform would suggest thatXgr may be lower. The breakdown of scapolite to calcite and quartz (reaction (1)), is a common reaction in granulite facies calc-silicates, and is interpreted to be a consequence of cooling, virtually independent of pressure and fluid composition (Goldsmith and Newton, 1977 Huckenholz and Seiberl, 1990). In the case of assemblage A, in which scapolite compositions are EqAnz¡-gt and plagioclase compositions approximate An9g, the scapolite stability relations of Goldsmith and Newton (1977) suggest that the breakdown of scapolite would have occurred during cooling through temperatures of between 72O-760'C. However, the scapolites in assemblage B, with compositions of EqAn72-75, show no record of reaction to anorthite and calcite, indicating that the cooling history preserved in the rocks does not extend below approximately 680oC. The breakdown of wollastonite to calcite and quartz (reaction (2)) is also consistent with cooling and/or an increase in a(COz) (Fig. 5.3), and is also pressure dependent, occurring at lower temperatures a,s pressure decreases (Schmulo vich, 197 7 ). The formation of coronas of grossular in wollastonite-scapolite calc- silicates (Reactions 3 & 4) has recently been described in detail by Fitzsimons and Harley (L994) using samples from the Nemesis Glacier region as an exarnple. Similar grossular coronas have been recorded in a number of other tenains (e.g. Schenk, 1984; Warren et al., 1987; Motoyoshi et al., l99l; Dasgupta, 1993), but their significance in terms of the P-T-fluid history has often been unclear. Wanen et al. (1987) interpreted grossular coronas between scapolite and wollastonite * calcite from the Arunta Block as reflecting either a decrease in a(COz) and/or a 107 Chapte r 5, Wollastonite- sc apolite calc - silicate s decrease in pressure. However, more recent studies (Moecher and Essene, 1990a; Harley and Buick, 1992, Fitzsimons and Harley , 1994) have established that this decarbonation reaction has a negative dT/da(COz), and therefore proceeds with decreasing temperature and/or increasing pressure and/or H2O infiltration. In particular, Fitzsimons and Harley (1994) have shown that decarbonation during cooling is a likely explanation for the grossular rims in both the Arunta Btock and Nemesis Glacier calc-silicates. The very steep negative slope of reaction (3) in T- a(COz) space (Fig. 5.3) would suggest that at least some decarbonation (probably through infiltration of HzO) occurred at high temperatures during cooling in Else Platrorm calc-silicates, followed by continuing cooling through reaction (4). In summary, peak metamorphism in the Else Platform region was in the vicinity of 800oC at around 7 kbar, and was followed by a period of near isoba¡ic cooling to 700oC or less. The sequence of reactions observed at Else Plaform and the inferred P-T-fluid history is consistent with observations elsewhere in the northern Prince Charles Mountains (Buick et al., 1991; Fitzsimons and Thost, 1992; Fitzsimons and Harley,1994), which have inferred isobaric cooling paths on the basis of calc-silicate textures. Peak conditions of 830oC at 6-7 kbar followed by isobaric cooling were infened by Fitzsimons and Thost (lgg2) for the Aramis Range, although no subsequent hydration phase was inferred. Similarly, Fitzsimons and Harley (1994) suggest peak metamorphic conditions of 800oC and 7 kbar for the calc-silicates of the Nemesis Glacier region, followed by cooling, with little or no preceding decompression. 5.3 Calc-silicates from Mt. McCarthy, porthos Range Calc-silicates containing the primary assemblage diopside-wollastonite- scapolite -calcite-quartz (964MM2) arc found as a layer within garnet-bearing pyroxenites near the western end of Mt. MccaÍhy in the eastern Porthos Range, 100 km to the east of Else Platform (Fig. 5.1). In these rocks, calcic scapolite 108 Clnpte r 5, Wo llastonite- scapolite calc- s ilicate s (Meqo) has extensively reacted to symplectites of calcite, anorthite, quartz and hyalophane, and the existence of abundant hyalophane in these symplectites suggests that the Ba-rich nature of scapolitic calc-silicates is a feature not unique to the Else Platform region. Breakdown reactions of wollastonite to calcite, q\artz and minor diopside (which incorporates the small Fe-Mg component of the wollastonite) have also occuned. A similar cooling history to Else Platform is infened for ML McCarthy, which is also similar to ttrat infened by Fitzsimons and Thost (1992) from calc-silicates in the nearby Aramis Range. 5.4 Ba-silicates in NPCM calc-silicates An unusual feature of the calc-silicates from Else Platform and Mt. McCarthy is the occurrence of barium silicates at both localities. Most recorded occunences of Ba-silicates, particularly Ba-feldspars, in high grade metamorphic rocks have been within Mn-rich metasediments and have been interpreted as metamorphosed submarine exhalites, generally associated with significant Mn and Ba mineralisation (e.g. in wilkes Land, Antarctica (Plimer and Lovering, 1983); Broken Hill, Australia (Mason, 1987); Kodura, India (Sivaprakash, 1980) and Aberfeldy, scotland (coats et al., 1984)). In comparison, Ba-feldspars are only rarely found in non-mineralised Mn-poor metasediments. Therefore, the occurrence of Ba-silicates in calc-silicates from the northern Prince Charles Mountains is of some interest. There æe three main occurrences of Ba-silicates in NPCM calc-silicates. The most common is as hyalophane within fine-grained symplectites of calcite and anorthite after scapolite at Else Platform and Mt. McCarthy (Fig. 5.4 (a)), although hyalophanes also occur as grains up to lmm in diameter in sample 964656D (Fig. 5.a 0)) and in this case appear to be parr of the peak assemblage. Veins with a composition approximating that of armenite, a rare hydraæd barium silicate, also occur within scapolites in sample 964MH1. 109 Chapte r 5, Wo llastonite- scøp olite calc-silicates lt t å / yol t¡: t s a hya I t -l Í Lt t { (a) (b) Fig. 5.4. Back-scattered elecnon images of barium silicates in calc-silicates from Else Platform. Scale bar is 100Pm. I (a) Symplectite of anorthite, calcite, quartz scapolite. and hyalophane consuming f.. 1 Sample 656C. (b) Symplectiæ of wollastonite and anorthite, with quartz, separating primary wollastonite and hyalophane. Sample 656D. ìr (c) Veins of a hydrous Ba-aluminosilicate (armenite) cutting scapolite. Brighter fl armenite is more Ba-rich, whilst darker armenite has higher alkati substitutions. -$ l5 ¿9 1 ti0EE.r.E-lE[,ELit Sample MHl. (c) 110 Clapter 5, Wollastonite - scapo lite calc - silicate s Specimen 964656D has hyalophanes up to one millimetre in diameter, with the average composition Or66Cn29Ab5 which have broken down in the presence of wollastonite and clinopyroxene to wollastonite-anorthite-calcite-quartz (Fíg. 5.4 (b)). This seemingly highly unlikely reaction can only be explained by high temperature mobilisation of barium and potassium, possibly associated with metasomatic addition of calcium. The presence of appæent peak hyalophane in this rock suggests that barium may have originally been present in localised parts of the sequence, and been subsequently remobilised to appear in veins and symplectites elsewhere within the calc-silicates. In specimen 964MH1, veins of a hydrated Ba-Ca aluminosilicate cut scapolite-rich assemblages (Fig. 5.a (c)) in which the breakdown of Ba-absent scapolite has also resulted in the formation of anorthite-calcite-hyalophane-quartz assemblages. Electron microprobe analyses of this mineral (Appendix 3) have shown that it has a composition which approximates that of armenite [BaCa2Al6Si9O3g.2H2O], although the occurrences were too fine grained for confirmation of this identification by X-ray diffraction analysis. Assuming this identification is correct, then to the writer's knowledge, this is only the seventh reported occurrence of armenite. It was first described as forming pseudo- hexagonal prisms in a calcite vein at the Armen silver mine in Kongsberg, Norway, associated with Ag-Ni-co-As mineralisation (Neumann, 1941). At Broken Hill, Australia, it has been identified as a major constituent of a calc-silicate rock with minor celsian, and as a minor constituent in a b¡owniæ and celsian rich aplitic gneiss (Mason, 1987). Armenite also occurs in veins cutting metasomatised diorite at Rémigny, Quebec (Pouliot et al., 1984), and a similar occunence has been reported from the Tokovian granite massif, USSR (Semenenko et al., 1987). At the Su Zufurin Ba-F mine in Sardinia, armenite forms bands and veins at the contact of skarns and hornfels (Balassone et al., 1989), and it forms porphyroblasts within sulphidic calc-silicate-rich quartz-rock in the Dalradian of 111 Clapter 5, Wollastonite - scapolite calc-silicate s Scotland (Coats et al., 1984; Fortey et al., 1991). The armenite in veins in 964MHl has variable mineral chemistry, and has significant Na and K substitution. Previous reported occwrences of armenite have had only minor alkali substitutions [(Na+,K+)Si4+ <=> (Ba2+,çv2+)Rtl+1 but the most alkali-rich armenite in the Else Platform calc-silicates has the composition (Ba¡.5sCa1.62Na6.33K0.¡¿)Als.¡zSi9.73.xH2O. In Figure 5.4(c) ir can be seen thar alkali-poor armenite generally occurs as rims between alkali-rich armenite and scapolite, and also forms nalrow veins. Mineral analyses of armenites using the WDS system on the electron microprobe are given in Appendix 3. The Ba-absent, K-poor nature of the peak wollastonite-scapolite- clinopyroxene assemblage in all samples except 964656D, and the relatively Ba- and K-rich retrograde symplectites suggests that the barium was introduced to these assemblages by fluids under relatively high temperature conditions, postdating the peak of metamorphism. In the same rock, narrow rims of wollastonite form between quartz and calcite, whilst n¿urotry hyalophane rims exist along nearby calcite boundaries at contacts with anorthite. As the hyalophane in this sample is also found in symplectites resulting from the pressure and fluid independent cooling reaction scapolite => calcite + anorthite, then the reaction calcite + quartz => wollastonite would presumably represent an increase in a(H2O). It therefore appears likely that fluid flow following the metamorphic peak resulted in mobilisation of ba¡ium and potassium during cooling in the NpcM. 5.4.1 origins of NPCM wollastonite-scapolite calc-siricates The occurrence of barium silicates in both of the wollastonite-scapolite calc- silicates described above from the NPCM, raises the question of why barium silicates should occur in calc-silicates over such a wide area. Although separated by 100 km, both occulrences are rema¡kably similar in their mineralogy and T12 Clwpte r 5, Wollastonite- scapolite calc-silicate s lithological associations, suggesting they are of a similar origin, and are perhaps the same unit. The two most important similarities are the occurence of barium feldspars in symplectites after scapolite, and the spatial association of the calc- silicates with pyroxenites. The abundance of barium in the sequence, associated with mafics in a supracrustal setting, would be consistent with the calc-silicates being deposited as carbonate-rich sediments in a submarine setting with volcanic exhalitive input. Alternatively, it is possible that the barium could have been metasomatically introduced to the sequence by fluids released during crystallisation of melts. Irrespective of the origins of the barium it was clearly mobilised, along with potassium, during high grade metamorphism in the NpCM. 5.5 calc-silicates from Rumdoodle, Framnes Mountains 5.5.1. Geological setting The Framnes Mountains, on the Mawson Coast of MacRobertson Land, 250-300 km north-northwest of the NpcM, form another relatively poorly understood fragment of the Late Proterozoic granulite terrain of East Anta¡ctica. The Framnes Mountains üe comprised of a supracrustal sequence of layered felsic gneisses, pelites and calc-silicates known as the Painted Gneiss (Trail, Lg70), which a¡e extensively intruded by the Mawson Charnockiûe, which has been dated at954+ 12Ma and 985 + 29 Ma (young and Black, 1991). clarke er al. (19g9) inferred an isobaric cooling path from pelitic reaction textures in the painted Gneiss, whilst conceding that this path was poorly constrained and did not preclude earlier decompression. In the following section, calc-silicates from the Painted Gneiss on Rumdoodle, a mountain in the northern Masson Range of the Framnes Mountains, will be briefly described to help constrain this metamorphic history, and compare it with the NPCM. The samples were collected from the northwestern face of Rumdoodle, to the north of the Rumdoodle Hut, on an ANARE expedition in 7965166. ll3 Clapte r 5, Wollastonite - scap o lite calc - silicate s 5.5.2 Petrography All three calc-silicate assemblages collected from Rumdoodle have wollastonite and diopside preserved in the peak assemblage, along with minor titanite, and are summarised in Table 5.2. The peak assemblages are granoblastic, with no apparent prefened orientation of grains. Sample wol scp cpx grs Dlas æ qrz ûtan 4288-91 PRr P P R+s P PRs P 4288-87/r PRz P P R¿s R¿o Roz PRsz P 4288-87t3 PR¡ (P?) P R¿s R4 Ra Rs P Table 52 Assemblages in calc-silicates from Rumdoodle, showing minerals which form part of the peak assemblage @) and rerograde reaction textures (R); 1 = rims between scapolite and quarø; 2 = rims between calcite and quartz; 3 = pd of late shear fabric; 4 = rims and symplectites between wollastonite and scapolite; 5 = rims between wollastoniæ and anorthite; 6 = symplectites after scapolite; 7 = oonsuming wollasúonite sample 4288-91 has the peak assemblage clinopyroxene(Xdi=60)- scapoliteBqAn6g) -wollastonite-anorthite with minor titanite and quartz. In places, scapolite has abundant inclusions of quartz, and where this occurs, scapolite has reacted extensively to intergrowths of blades of wollastonite with less abundant anorthite (Fig. 5.5(a)). Rims of grossular (Xg¡=93) with minor quartz occur between wolastonite and anorthite, and at contacts of these minerals with scapolite (Fig. 5.5(a, b)), and in places almost completely pseudomorph the wollastonite- anorthite intergrowths (Fig. 5.5(b)). These same reactions can locally be seen occurring on a smaller scale in some quartz inclusions within scapolite, where a narrow rim of wollastonite forms between the quartz and scapolite, and then grossular occurs between the wollastonite and the primary scapolite. Calcite, quartz and minor diopside are observed to partially consume large primary wollastonite. rt4 Cløpter 5, Wollnstonite-scapolite calc-silicates :-*f J ',/{r.Çl .:q /,Á,\ v "J., Fig. 5.5 Reaction textures in calc-silicates from Rumdoodle. In all cases the width of the field of view is 2 mm. (a) Scapolite, containing rounded inclusions of quafz, being replaced by intergrowths of wollastonite and anorthite, with coronas of grossular subsequently forming between wollastonite and anorthite and wollastonite and scapolite. Sample 4288-491. (b). Optically continuous scapolite being replaced by wollastonite and anorthite with subsequent development of coronas of grossular around wollastonite. In bottom left corner garnet has almost entirely replaces scapolite, wollastonite and anorthite. Sample 4288-91. (c) Scapolite reacting to form symplectites of anorthite and calcite where it is adjacent to diopside, and to symplectites of grossular a¡rd a¡rorthite where it is adjacent to wollastonite. Sample A288-87 ll 115 Chnpter 5, Wollastonite-scapolite calc-silicates o ; wol f. wol I Fig. 5.5 (continued) Reaction textures in calc-silicates from Rumdoodle. (d) Primary wollastonite (wol1) partially consumed by calcite and quartz. Secondary wollastonite (wol2) with minor prehnite, rims quartz at contacts with calcite. Crossed polars, sample A288-8711. Width of the field of view is 2 mm. (e) Back-scattered electron image of wollastonite rims between quartz and calcite, with minor prehnite and pyrite. Sample A288-8711. Scale bar is 100¡rm. (Ð Very fine grained symplectites of grossular and anorthite consuming wollastonite. No scapolite remains, but was almost certainly originally present. Plane polarised light, sample A288-87/1. Width of the field of view is 2 mm. (g) Back scattered electron image of the same grossular-anorthite symplectite shown in (f). Sample A288-8711. Scale bar is 100pm. (h) Shear fabric containing wollastonite, with minor anorthite, calcite and pyrite, with trails of hne grained grossular. Crossed polars, sample A288-8713. Width of the field of view is 2 mm. 116 C løpte r 5, W o I last onit e - s c ap o lite c alc - s i lic at e s In sample A288-87/1, wollastonite, scapolite, clinopyroxene and r titanitform the peak assemblage. In parts of this specimen, scapolite breaks down in the presence of wollastonite to form symplectites of grossular and anorthite. However, elsewhere in the rock, scapolite is consumed by symplectites of anorthite and calcite, with minor wollastonite. In places, a single scapolite is consumed by both grossular-anorthite and calcite-anorthite symplectites, with the grossular-bearing symplectites occuning proximal to wollastonite (Fig. 5.5(c)). Where grossular-anorthite symplectites have formed, grossular-quartz rims exist between anorthite and wollastonite. Wollastonite occurs as large primary grains, which are often partly consumed by calcite and quartz. However, wollastonite is also found forming rims between secondary calcite and quartz which resulted from the breakdown of primary wollastonite (Fig. 5.5(d,e)). This secondary wollastonite is intergrown with prehnite, and is spatialty associated with clusters of fine grained pyrite. In sample A288-8713, the inferred peak assemblage is wollastonite- scapolite-clinopyroxene-quartz with minor titanite. Extensive symplectites of grossular(Xgr=92) and anorthite occur a¡ound wollastonite (Fig. 5.5(f,g)), and only minor amounts of scapolite remain in the rock. Scapolite is generally pseudomorphed by wollastonite-anorthite intergrowths, simila¡ to those 4288-91, as well as the grossular-anorthite symplectites adjacent to wollastonite. Grossular rims with minor quartz occur at contacts of anorthite with wollastonite. In addition, late recrystallisation of parts of the assemblage during high strain has resulted in the development of a fine grained wollastonite-rich assemblage, which also contains variable grossular, anorthite and calcite, as well as trails of fine grained pyrite (Fig. 5.5(h)). 5.5.3 Interpretation of reaction textures The reaction textures from Rumdoodle are broadly similar to those from the rt7 Clnpter 5, Wollastonite - scapolite calc - silic ate s NPCM, in that they generally involve the breakdown of wollastonite and scapolite, and the development of grossular coronas. Rims of grossular between scapolite and wollastonite, and grossular-anorthite symplectites after scapolite and wollastonite can be attributed to the reaction: scapoliæ + wollastonifs => grossular + anorthite + COz+ H2O (6) Grossular rims around wollastonite at contacts with anorthite are found in all three samples from Rumdoodle, and these indicate the progress of the reaction: 2 wollastonite + anorthite => grossular + quartz (7) The reaction of wollastonite to calcite and quartz, and the subsequent formation of wollastonite as rims between calcite and quartz in 4288-87/1, is a result of the progress of, and then the reversal of, reaction (2). The breakdown of scapolite to calcite and anorthite in the same assemblage resulted from cooling through reaction (1), although the growth of wollastonite in these symplectites means that wollastonite must have been stable at this time. The early growth of wollastoniæ at scapolite-qua.rtz contacts, and the growth of wollastonite and anorthite blades in primary scapolite in 4288-91, suggests the early progress of the reaction: scapolite + quartz => wollastonite + anorthite + COz+ H2O (8) Fig. 5.6 shows a petrogenetic grid in the cAsv system at a pressure of 7 kbar, adapted from Fitzsimons and Harley (1994), indicating the locations of the above reactions in T-a(Co2) space. whilst the precise T and a(Coz) values are poorly constrained due to variations in the activities of grossular and meionite and uncertainties in the appropriaæ pressrue, the grid is still a useful tool in deærmining the retrograde P-T-a(co2) history in the Framnes Mountains. In 4288-91, the early development of wollastonite at scapolite-quartz contacts and blades of anorthite and wollastonite after scapolite may well reflect an early stage of decompression. Alternatively, this texture could reflect an influx of H2O at high temperatures (probably >800oc) near the peak of metamorphism. The development of grossular rims between scapolite, wollastonite and anorthite in 118 Chapter 5, Wollas tonite -scøp olite calc -s ilicates (8) (6) (3) 7 kbar o o B 2 çcoll o a áo Wo An 850 (7',) wo Grs Qtz Wo Scp V (4r Grs Cal Qtz \ 800 P--_ / T \ / ..l / / ("c) oQ / 750 \ (2) 6 kbar Scp (EqAn = 78) (1) \ An Cal 700 c,.Q ls Framnes Mountains EqAn = 78 Xgrs = 0.76 P=7kbar 0.2 0.3 0.4 0.5 0.6 CT COe Fig. 5.6. T-aCO2 grid for the CASV system adapted from Fitzsimons & Harley (tgg1) at 7 kbar, showing cooling and decarbonation reactions observed in calc-silicates ftom the Framnes Mountains. T and aCO2 values a¡e only approximate, due to the varying activities of grossular in garnet and meionite in scapolite. The dashed line shows the effect on the grid topology of I kbar of decompression. The horizontal shading indicates the stability field of the wollastonite-scapolite-quarø æsemblages in 4288-91 and 87/1. Numbers refer to equations mentioned in the text. 119 C lnpte r 5, Wollast onite- scapolite calc - s ilicate s 4288-91 may reflect one of two possibilities. It is possible that cooling (+ Hzo influx) occurred through reaction (6) and then through reaction (7). However, no anorthite was produced during the grossular-forming reactions, and grossular rims appear to have formed simultaneously between all three minerals. Therefore, it appears likely that, during cooling, the path in T-a(CO2) space was buffered along reaction (8) until reaching the [cal] invariant point, when wollastonite, scapolite and anorthite would have then all reacted to form grossular and quartz. In 4288-87/1, the first reaction to occur was reaction (6), followed by further cooling through reaction (7). h this case, anorthite produced by reaction (6) subsequently reacted with wollastonite to form grossular rims, and therefore the internal buffering parh along reacrion (8) suggested for 4288-91 appears unlikely in this case. The breakdown of wollastonite and then the reversal of this reaction suggests that following cooling through reaction (2), there was a significant flushing of H2o through the terrain, resulting in very low a(co2). The breakdown of scapolite to calcite and anorthite during cooling may have occurred at a similar time to the secondary wollastonite development, as it occurred when wollastonite was stable. This cooling would have occuned at around 730-i40oC, requiring the a(co2) to have been lower than approximately 0.2 to allow wollastonite stability. A late shea¡ fabric in 4288-8713, apparenrly post-dating all earlier textures, is in places defined by wollastonite with minor calcite, suggesting that this late H2O influx accompanied shearing. Fig. 5.7 shows the rwo fluid independent reacrions (l) and (7) in p-T space, appropriate for 4288-87|L, raking into account the Na content in the scapolite after Goldsmith and Newton (1977) and the andradite content in garnet after Huckenholz et al. (1981), and these clearly show that a cooling-dominated path is favoured for 4288-87/1. In addition, the relative timing of reactions (1) and (7) is important, as, in the case of 4288-87/1, it would constrain the pressure r20 C hapte r 5, W o llast onit e - s c ap o I ite c alc - s ilic ate s at which the cooling took place to being above or below 5-5.5 kbar. However, no convincing timing relationships were observed between these two reactions, although in 4288-87/3, which is compositionally similar, grossular-producing reactions consumed all scapolite before it broke down to anorthite and calcite. It is therefore considered probable that reaction (7) preceded reaction (1), and that cooling took place at pressures above approximately 5.5 kbar. 10 ct) Ir (\ (ú o cr -o o I.JJ l¿ c I (ú E fL ó 6 4 700 800 900 T ('C) Fig. 5.7 P-T constrainß on Rumdoodle calc-silicates using the fluid independent reactions, (1) (Newton & Goldsmith,l9TT) and (7) (Huckenholz et al, 1981), showing cooling through 800-700"C at 6-7 kba¡. 4288-8713 is in most respects similar to 4288-87/1, except that primary scapolite has been almost entirely consumed. However, within the extensive grossular-anorthite symplectites, grossular-rich pseudomorphs of what was probably primary scapolite occur. These textures are consistent with cooling through reactions (6) and (7), which would indicate that cooling occuned from high temperatures (c.800"C). The early development of wollastonite rims between scapolite and q\artz, and wollastonite and anorthite apparently consuming scapolite in A288-91 could possibly indicate early decompression before cooling, as shown in Figure 5.g, T2T Chapter 5, Wollastonite-scapolite calc-silicates although only less than I kbar of decompression would be necessary to allow this reaction to proceed. This could alternatively be explained by an influx of a hydrous fluid near the peak of metamorphism, and it is clear from figure 5.8 that most reactions observed from Rumdoodle would require compression if the conditions were near isothermal. However, early decompression of up to 1 kbar may have occuned before significant cooling. u) óC) o o E { Grs CalQtz 7 I Wo Scp P Wo An (kbar) 6 \ \ 0.2 0.3 0.4 0.5 0.6 aco2 P-a(CO2) grid Ïg: _f.s. in arley (L994), constructed for an X calc- silicates from Rumdoodle. wollastonite-scapolite-quarøuuu peak appr'ximate. onlY In summary, the reaction textures in wollastonite-scapolite calc-silicates from the Framnes Mountains preserve a simila¡ apparent retrograde cooling history to the calc-silicates from the NPCM, along with evidence for significant hydrous fluid flow late during the retrograde history. Peak conditions in the Framnes Mountains appear to have been in excess of g00oc and 5.5 kbar, and cooling of around 100oC took place with little evidence of significant accompanying t22 Clapte r 5, Wollastonite - scapolite c alc-silicate s decompression or decompression preceding cooling. 5.6 The role of the ftuid phase The presence of peak wollastonite in most calc-silicate gneisses from MacRobertson Land precludes pervasive CO2 flushing as a cause of granulite metamorphism, and the spatial association of these wollastonite-bearing calc- silicates with pyroxenites on Else Platform also means that peak metamorphism was either fluid absent or, more probably, that the fluid phase was internally buffered. Clearly, however, fluids have played a role in the evolution of the calc- silicates from both the NPCM and the Framnes Mountains. On Else platform, the evidence for this includes the apparent mobilisation of barium and potassium during cooling, and the formation of late clinozoisite after diopside and anorthite. This clinozoisite development possibly coincided with the hydration event which resulted in the formation of hornblende and biotite observed in late-D2 high strain zones and as coronas in magnesian pyroxenites on Else platform. The calc-silicates from the Framnes Mountains show considerably more evidence for fluid flow during the cooling history than the NPCM calc-silicates. As in the NPCM the stability of wollastonite-bearing assemblages precludes pervasive coz flushing during peak metamorphism. However, the consumption of wollastonite by calciæ and quartz was not purely around the edges of the grains and along cleavage planes, but instead followed irregular pathways through wollastonite grains, often continuing into neighbouring clinopyroxene. This, along with the concentration of pyrite within the calcite and quartz, strongly suggests that a fluid may have been at least partly responsible for the breakdown of wollastonite, and this fluid must therefore have raised the a(coz). A possible source for co2 bearing fluids in the Framnes Mountains may be the Mawson Charnockite, as charnockites have been known to evolve CO2 rich fluids which can raise the a(Coz) of the surrounding country rock (Frost and Frost, 19g7), and 123 Chapter 5, Wo llastonite - scapolite calc- silicate s the Mawson Charnockite is known to have intruded after peak metamorphism (Clarke et al., 1989). A late influx of hydrous fluids is suggested by late secondary wollastonite development between calcite and quartz, and it seems likely that when the rocks cooled through c.730oC that wollastonite was stable and the a(coz) was therefore below -0.2. This late influx of H2o appears to be synchronous with the mylonitic fabric in 4288-8713, and is therefore tikely to be related to the D4 mylonites described by Trail (1970) and Clarke (1988), which they described as accompanying hydrous regression. In summary, it seems clear that granulite metamorphism in MacRobertson Land cannot be attributed to pervasive flushing of Co2, and that in the NpcM, there is only evidence for very late hydrous fluids, possibly at the end of Dz. In the Framnes Mountains, however, a more complex fluid evolution, occurred after peak conditions, with evidence for significant variations in a(coz), which complicate the interpretation of the pressure-temperature evolution of the region. 5.7 Discussion The cooling path described for the calc-silicates of the NPCM is in ma¡ked contrast with the decompression path inferred from wollastonite-scapolite lithologies in the Rauer Group (Harley and Buick, 1992). However, all other calc- silicate occuffences documented from the Proterozoic of East Antarctica preserve textures consistent with neæ-isobaric cooling. For example, Motoyoshi et al. (1991) described similar reaction textures to those from Else Platform from the Bollingen Islands, near the Larsemann Hills, and attributed these to represent a cooling path, possibly from the 1000 Ma event, although rhost et al. (1991) have interpreted mafic and pelitic reaction textures in the region to represent a retrograde history initially dominated by decompression. The breakdown of scapolite to calcite and anorthite in wollastonite-scapolite calc-silicates has also been described in the Reinbolt Hills, 80 km east of Else plaform (Nichols and Berry, 1991). t24 Clnpter 5, Wollastonite - scapolite calc -silicate s The P-T-fluid history recorded by wollastonite-scapolite calc-silicates in the northern Prince Charles mountains is of particular interest when viewed in the general context of the Proterozoic terrain of Kemp and MacRobertson Lands and the Prydz Bay coast. This tenain has long been considered to be a good example of a granulite terrain which has undergone near-isothermal decompression (e.g. Harley and Hensen, 1990), and this interpretation has been made largely on the basis of well-documented reaction textures in pelitic and mafic granulites from the region (Harley, 1988; Nichols and Berry, l99r; Fitzsimons and Harley,l992a). However, in the northern Prince Charles Mountains, calc-silicate reaction textures, both on Else Platform and Mt. McCarthy (this study) and in the Nemesis Glacier region (Fitzsimons and rhost, 1992; Fitzsimons and Harley, L9g4) strongly suggest a near-isobaric cooling path following the metamorphic peak. This is of particular importance when one considers that cordierite-bearing coronas and symplectites within pelites, on which decompressional P-T interpretations have previously been based, are also found in the northern Prince Cha¡les Mountains (Else Platform, Trost Rocks, Fox Ridge (this study)) and, in the case of Else Platform, are found within a kilometre of calc-silicates preserving classic 'cooling textures'. Similarly, in the Reinbolt Hills, 80 km east of Else Platform, Nichols and Berry (1991) inferred a decompressional P-T path on the basis of spinel- cordierite symplectites after garnet and sillimanite, whilst in the Nemesis Glacier region, 50 km west of Else Platform, a cooling path has been infened on the basis of calc-silicate textures (Fitzsimons and Harley, 1994), yet both the Reinbolt Hills and Nemesis Glacier reaction textures are found on Else Platform. In fact, the biotite-poor pelites on Jetty Peninsula excellently preserve an apparent decompression path, whilst the calc-silicates have an equally convincing apparent . cooli1 pattr. onty in pelites of bulk compositions allowing rhe development of syn-D2 biotite-bearing assemblages does the added complexity become app¿¡rent. This clearly has important implications for the interpretations of P-T paths on the basis of a single assemblage without appropriate time constraints. t25 Clapte r 5, Wollastonite - scapolite calc-silicate s A further reason for the preservation of cooling textures with no decompression is the insensitivity to decompression of calc-silicate assemblages which do not contain primary garnet (Harley, pers. comm., 1993). This may explain the somewhat common occuffence of cooling textures in calc-silicates in terrains which otherwise record a decompression-dominated history, such as the Reinbolt Hills (Nichols and 8.rry, 1991) and the Bolingen Islands (Motoyoshi et al., 1991). Calc-silicates from these localities contain no primary garnet and record cooling textures, yet in the Rauer Group, which contains similar decompression textures in pelites to the Reinbolt Hills and Bol ingen Islands, Harley and Buick (1992) document decompression textures in calc-silicates, through reactions involving the consumption of primary garnet. However, in view of the development of grossular coronas on Else platform, it is very unlikely that decompression of more than I kbar could have accompanied cooling, an observation which was also made by Fitzsimons and Harley (Lgg4) for the Nemesis Glacier calc-silicaæs. In previous chapters evidence has been presented that the 'decompression textures' in pelites at Else Platform and Trost Rocks are actually a function of two overprinting thermal events at differing pressures. The high temperatures infened for the peak assemblage (c.800oC), are consistent with peak metamorphic conditions calculated from pelitic lithologies, and therefore the isobaric cooling path is interpreted as cooling from the same peak (syn-Dl) metamorphism observed in the pelites from Jetty Peninsula. It would appear probable that the pelites and calc-silicates from the NPCM preserve differing portions of the same metamorphic history, and that the temperatures of c.700oC inferred for the second thermal event in the metapelites would have been insufficient to reverse the earlier cooling reactions preserved in the calc-silicates. It seems unlikely that the symplectites from the Else Platform calc-silicates could have formed 126 Clnpter 5, Wollastonite- scapo lite calc -s ilic ate s synchronously with the 'decompression' symplectites, as little or no decompression appears to have occurred during cooling from peak conditions at Else Platfoffn, as discussed above. Although no direct evidence exists to confîrm that the cooling textures in the calc-silicates from MacRobertson Land represent a continuous history following peak metamorphic conditions, the fact that a number of cooling reactions occurred in an identifiable sequence, particularly in the Framnes Mountains, supports this interpretation. However, later hydration reactions and recrystallisation in high strain zones at lower temperatures of 650- 700oc may be related to subsequent events (e.g.Dzon Else Plaform, D3 and Da in the Framnes Mountains). Although more precise geochronological constraints are required before the true significance of the various reaction textures becomes clear, these observations highlight the markedly different apparent P-T histories which can be preserved in rocks of differing bulk composition, and the subsequent doubt which must be placed on any interpretation based on observations from a single lithology. 127 Chapter 6, Fox Ridge Chaptet 6 Metamorphic evolution of pelitic shear zones at Fox Ridge, Mcleod Massif 6.1 Introduction In previous chapters, evidence has been presented suggesting that the metamorphic evolution of the granulite terrain of the northern prince Charles Mountains may be significantly more complex than a simple P-T 'loop', but further constraints are needed to confirm the relative timing relationships between assemblages, and their context within the structural evolution. In particulæ, the thermal and baric evolution following the syn-D2 thermal peak remains poorly constrained, as is the timing of late garnet and sillimanite overgrowing cordierite coronas. It is clea¡ that the textural preservation of different portions of the p-T history is highly dependent on the deformation associated with these events. For example, pelitic gneisses on Kamenistaya Platform, which were relatively weakly deformed by D2 in comparison with lithologies on Else Platform, generally preserve less in the way of syn-D2 biotite-sillimanite fabrics and cordierite-spinel coronas, and as a result, relative dating of these assemblages through correlation of fabrics was not possible. In comparison, metamorphic assemblages associated with high strain fabrics are particularly useful in that they allow the correlation of the appropriate assemblage with the regional structural framework. This is particularly the case when a number of overprinting assemblages form along the same fabric, making it reasonable to assume they form a continuum during a single event. In this chapter, petrographic relationships within high grade shear zones on t28 Chapter 6, Fox Ridge Fox Ridge, 50 kilometres SSE of Else Platform on Mcleod Massif, are described and interpreted in the context of the metamorphic evolution established for nearby fragments of the tenain. These shear zones cut pelites and semi-pelites containing a coarse granulite facies assemblage, and the well defined high strain fabric in these rocks provides important constraints as to the timing of the mineral assemblages relative to the structural history of the region, and these can be correlated with assemblages elsewhere in the NPCM, and placed in the context of available absolute age constraints 6.2 Geological Setting and Structural Framework Fox Ridge is a east-west trending ridge on the western side of Mcleod Massif, overlooking Grainger Valley, 50 km south-southeast of Else Platform in the northern Prince Charles Mountains (Fig. 6.1). The geology of the region has not previously been described in detail, although McKelvey and Stephenson (1990) carried out a geological reconnaissance of the Radok Lake region immediately to the south and described a sequence of felsic orthogneisses interlayered with mafic, aluminous and minor calc-silicate gneisses, tightly folded about an east-plunging axis. Due to the lack of in situ outcrop and extensive snow cover on Fox Ridge, only limited structural observations were possible, and hence much of the structural interpretation relies on correlations with neighbouring regions of the NpCM. The high grade shear fabric on Fox Ridge is the last apparent pervasive high grade ductile deformation in the region, and trends east-west, and generally dips steeply to the south, with an elongation lineation dipping moderately to steeply to the east. East- west trending, granulite facies shear zones have been described elsewhere in the NPCM, and it appears likely that a broad correlation can be made between these high strain zones and the Fox Ridge shear zones. On Jetty Peninsula, 50 km NNW of Fox Ridge, a 1-2 km wide east-west trending, high strain zone (the Soviet Shear Zone), dipping steeply to the south, formed late during D2, which was a regional event 729 Chapter 6, Fox Ridge ê/) o Permo-Triassic ü oO b6 OE 68oE ,/- v a bo qJ sediments A hos Range oa t o ú \- Q(l dy:--- ScYlla w = Ð w#" s FOX RIDGE 4 ofu\ s {0'E Napier 60"E Complex Mawson / / / Rayner Complex Framnes Mountains Prydz Bay / Davis 1 ! ,/Anery ) læ J Shelf Northern Prince Larsemann Charles Mountains // Hills Rei nbolt Hiils 200km Fig. 6.1. Map showing the location of Fox Ridge, and the location of the D2 shear zones on Jetty Peninsula and Fox Ridge,separated by the Beaver Lake Graben. 130 Chapter 6, Fox Ridge involving tight folding of the regional gneissic fabric, 51, around an east-west trending axis (see Chapter 2). Fitzsimons and Thost (1992) also described late high grade E-W trending shear zones in the Porthos and Aramis Ranges, which extend to the west and northwest of Fox Ridge, and these were assigned D6 in their structural framework. In addition, McKelvey and Stephenson (1990) described a large scale, east plunging, tight fold in gneisses immediately to the south, contrasting with the consistent high strain fabric on Fox Ridge. The folds described by McKetvey and Stephenson (1990) also had an east plunging lineation, suggesting that, as on Jetty Peninsula, the Fox Ridge shear zone may reflect intensification of strain on the limbs of these folds. It therefore seems clear that the Fox Ridge shear zones have the same characteristics as the late D2 shear zones on Jetty Peninsula and the D6 shear zones of Fitzsimons and Thost (1992). This is supported by the observation of the timing of the high strain deformation relative to the metamorphic history, which is discussed in section 6.6. 6.3 Petrography All samples collected from Fox Ridge contain an early coarse grained assemblage, in which elongate minerals, particularly sillimanite, often define a fabric which is termed 51. This 51 fabric is overprinted by a localised to pervasive 52 sheæ fabric, which varies from being localised dynamic recrystallisation along grain boundaries in some samples, to being pervasively mylonitic and locally ultramylonitic in others. Both the 51 and 52 fabrics are defined by granulite facies assemblages, which shall be termed M1 and M2 respectively. The complex textural relationships are shown in Table 6.1. 6.3.1 Metapelites In all pelitic samples collected from Fox Ridge, the M1 assemblage contains garnet, often with inclusions of sillimaníte, quartz, spinel and less commonly biotite. Quartz, plagioclase and ilmenite are always present, and K-feldspar and rutile are 131 Chapter 6, Fox Ridge common. Occurrences of rounded cordierite within garnet are also interpreted as primary inclusions, and in three samples, large porphyroblastic cordierite also forms part of the peak assemblage. The occurrence of sillimanite as part of the M1 assemblage is variable. In some samples the M1 assemblage is sillimanite-free and in others sillimanite forms fibrous inclusions within garnet. Most commonly, however, coarse blades of sillimanite defining the 51 fabric occur both as inclusions within garnet and wrapping around large garnet porphyroblasts. Spinel often occurs as inclusions within the sillimanite, and corroded remnants of spinel within the matrix a¡e also interpreted to have been part of the primary assemblage. garnet Rock spinel biotite sill rutile illrt cord ksoa¡ corun oDx 48.4C 1o3n 163 4- * {. 1u3n 5A t"3j 16 4go. lbc4mq 163¡4g 2n * I è 5B 1"3jop 1u¿ 164gm l"4mq 16 1u4m la2eJ * 8C 1a 4tll 4m 1 ¡f 2go 104 lu3jn 1u¿"4 4m l"4q 1d* 162.¡3¡ 4.. * 'grJ 1e 18A leE 1¿ 4g lb"4^ 1u¿4m 2g * 18B 1"3jn 1d" 4m 1"4-q 2¡ 2:tv 2g r¡ 19A 1a3jn 16 4m 1bc4mq 16 2¡3y4¡¡, * r* 198 1a3j ta2j 4m 1"4otq 162¡93¡ 2sJ * 208 1u3np 1b" + 4q 1u4e 1¡4g 2g3o * 1u3n 1e 4o 20c è 4q 164gm ta2e * 238 te3j lua 16 1uc4q 1u4¿ 2o * lb* è UA 1a3np 1b" 4n^ lb4-o ? 1¡¡4- l&s * Table 6.1 Assemblages in lithologies from Fox Ridge (all sarples contain quartz and plagioclase). Relative timing: I = associated with the Ml assemblage; 2:secondary co.¿ierite-riôl ãssemblage; 3 = earlY high srain assemblage; 4 = late high strain assemblages; * phase present with no clear textu¡al relationships Textural relatioruhip.: p.itrtary a: l.g9 porphyroblasts; b = inclusions in peak garnet; c = coarse sill wrapping garnets; d = inclusion in sillimanite; e = conoded spinel in mãtrix;- f = inclusion in cordierite; g consuming peak garnet; h early biotite j = = consumed by cordierite; i = consuming biotite; = coronas around sillimanite; k = symplectitic intergrowth with laie garnet; m = associateC wit¡ tate high fabric; srain n = small trails of grains defining Jnear fabric; o = õordierite recrystallised in shea¡ p gamets fabric; = small replacing recrystallised cordierite coronas; g = late fibres on cordierite and sillimaniæ grain boundaries; r = associaæd with late ultramylonite development. The primary assemblage is overgrown by a second assemblage which is dominated by cordierite, plagioclase and quartz. These minerals consume and form 132 Chapter 6, Fox Ridge coronas a¡ound the garnet porphyroblasts, with quartz often forming symplectites with cordierite, and cordierite often forms coronas around sillimanite (Fig. 6.2 (a)). In some samples, ilmenite and spinel also consume sillimanite, apparently synchronous with the cordierite coronas (Fíg.6.2 (c)). Spinel inclusions in garnet are separated from the garnet by a moat of cordierite, with later biotite and fibrous sillimanite forming at the contact between the garnet and the secondary cordierite (Fie. 6.2 (b)). The cordierite coronas are overprinted and recrystallised by a high strain fabric which is locally mylonitic and va¡ies from being parallel to highly discordant to 51. In strongly mylonitised pelites from Fox Ridge, the mylonitic fabric is typically composed of cordierite, garnet, plagioclase , qvaÍfz and ilmenite (Fig. 6.2 (h)). In almost all samples, small garnets form trails which define the mylonitic fabric, overgrow primary garnet and also form clusters within recrystallised cordierite (Fig. (d, 6.2 e, g)). In a number of samples, sillimanite is pulled apart along the mylonitic fabric and these pull-aparts are infilled by ilmenite. Ilmenite also often forms vermicula¡ intergrowths with late garnet, and garnet also forms coronas around trails of ilmenite which define the mylonitic fabric (Fig. 6.2 (g)). The garnet-cordieritetsillimanite early high strain fabric is overgrown by a later biotite-sillimanite-ilmenite assemblage which defines the same high strain fabric (Fig 6.2 (Ð)' Early garnets are often pulled apart along the L2lineation are these fractures are infilled by biotite and ilmenite, whilst the secondary garnet is overgrown by biotite, sillimanite and ilmenite. Biotite and fibrous sillimanite also form along grain boundaries of cordierite, and sillimanite fibres also overgrow earlier sillimanite and spinel (Fig. 6.2 (i)). Late ultramylonites along the same fabric are defined only by biotite and ilmenite, and represent the final stage of mineral growth. In zones within a few millimetres of these ultramylonites, early garnets are substantialy enveloped by late biotite and ilmenite. t33 Chapter 6, Fox Ridge tt¡¡+Ñ..-, 1 a 'Þ\ 1.1i': - ,_ c ã;¿ l-{. _1,- I -.. - Fig. 6.2 Reaction textures in pelites from Fox Ridge. In all cases the width of the field of view is 2 mm. (a) Primary gamet and sillimanite consumed by cordierite. Sample 964-101^. (b) Moat of cordierite separating primary gamet from inclusions of spinel and ilmenite. Biotite and sillimanite occur at the contacts of cordierite with gamet, ilmenite and spinel. Sample 964-244" (c) Spinel occurring an inclusion in primary sillimanite (sill1), and also forming coronas with ilmenite around primary sillimanite. Secondary sillimanite (sillz) forms coronas between sillimanite and cordierite, and the high strain fabric is defined by biotite and sillimanite. Sample 964-10A 134 Chnpter 6, Fox Ridge lir T( J 1 N¡ ! 4 I [,,'.',' 'i,- " ¡ ¿' Fig. 6.2 (continued) Reaction textures in pelites from Fox Ridge. In all cases the width of the field of view is 2 mm. (d,e) Garnet partially consumed by cordierite, which is recrystallised in the high strain fabric. Where cordierite is recrystallised fibres of sillimanite and small garnets occur. A trail of secondary garnets define the high strain (S2) fabric. Sample 964-10A. (d) - plane polarised light, (e) - crossed polars. (f) Primary sillimanite with a corona of secondary gamet elongated in the 52 fabric. Sillimanite and biotite overgrow gamet and also define 52. Sample 964- 54. (g) Primary gamet (gt1) wrappedby the high strain fabric defined by sillimanite, ilmenite and secondary gamet (gt2). Sample 964-23A. 135 9ç,r 'Jþ-Þ96 êId('lleS 'cuquJ ururls q8lq ZS eql SuIuIJêp le(uu8 puu eûexorÁdoquo'1âu¡u3 pêuIuJB eurJ ruroJ ol pãsrilulsÁrcer eSelqulêss? suexorÁdoqilo-leuB8 esJ?oJ (l) 'J8-tg6 etdules 'lcor pâsrllglsfuca¡ ,(Ie¡rtue ur aSlequressE elrrelpJoc-elluutullllS-elllolg (l) 'Vnz-ng6 e¡dueg 'ellrâIprocJo sêIJupunoq urur8 pesl¡¡elsfucer 3uo¡u ¡ncco êtrueuIIIIIS puu elltolg '(eiluuurtps ,{¡eu¡ud ,(q peur¡ep) I5 ot a¡8uu qEq u te elrueturllrs puu 1euu3 ,fuepuoces ,iq peur¡ep cFqeJ zS pado¡elep ,{¡4ee¡¡ (¡) g0Z-þ96 eldruegleruu8 peurur8 aur¡ Jo slrerl puu zgenb puu esu¡cot8u1d 'e¡rralproc ,{q peur¡ep JIJqUJ cllluol,(ru (q) 'tutu Z sI Jo plslJ âtll Jo l{lpl/v\ eID ^\âI^ sesuc IIu u1 'a8pr¿ xod lrro{ selrled uI seJnlxâl uollceo¿ (penuuuoc) Z'9'ãtt ¡ , .t- .\ o É ì I t i ::. t. i ì;i" L ¡. ! \ \ t eì a h.'., ['\- * { I I ( ,t t t )- a a r,l ¡, l' a I ¡. ) !¡; l 1ì. ,T t $ 'i! ¿ t- rl I ,t aSpr¿xo¿'gr?itn Chapter 6, Fox Ridge 6.3 ¡tite-sillimanite rich rocks r a number of mylonitised and relatively magnesian samples (964-8C), the roc now dominantly composed of biotite, sillimanite and cordierite, with less rut nenite, quartz and feldspar (Fig. 6.2(()). Early porphyroblastic cordierite is sul ttially consumed by biotite with less abundant sillimanite, and highly conoded rer ¡nts of primary garnet have been consumed by cordierite and plagioclase and the ;ubsequently by biotite. Although porphyroblastic cordierite is often consumed by oiotite and sillimanite, some cordierite is recrystallised in the mylonitic fabric. Th : latest stable assemblage preserved in these rocks is biotite-sillimanite- pli '¡ioclase-quartz, and often this assemblage only weakly defines the fabric, sul esting that it may have formed during the waning stages of deformation. -Iighly 6.3 aluminous assemblages Sample 964104 is a highly strained cordierite-garnet-sillimanite gneiss sul ;nding a low-strain silica-undersaturated boudin. The boudin contains abundant sil, nanite with less abundant corundum, and the sillimanite often has inclusions of spit- ', ilmenite and rutile. The sillimanite and corundum are enveloped by spinel, ilmer I and cordierite. The timing relationships between the spinel and cordierite are often ar riguous, but in one region, spinel forms symplectites on sillimanite adjacent to cordir, e, and this symplectitic spinel is now enclosed by later sillimanite (Fig. 6.3). Th rreservation of spinel-bearing symplectites can be attributed to the low strain withirr 'e boudin. Spinel, ilmenite and primary sillimanite are often separated from cordierite y narrow coronas or fibres of secondary sillimanite (Frg. 6.3). Outside of 'e boudin, the assemblage is cordierite-rich, contains quartz and does not contain cot tdum, but the reactions are similar, and their relative timing is better constrained b. labric development. Large M1 garnets have occasional inclusions of spinel ar d lmenite, and are wrapped by coarse sillimanite. primary sillimanite often has incluslns of spinel, ilmenite and rutile, and then this sillimanite 137 Chapter 6, Fox Ridge is often rimmed by a second generation of these minerals (Fig. 6.2(c)). Garnet and sillimanite are often significantly consumed by cordierite, particularly where they are proximal to each orher (Fig. 6.2(a)). cd il I s¡ll 2 cor cd cd sp \ \ Fig. 6.3. Reaction textures within the aluminous boudin in sample 964104. Ml sillimanite and corundum a¡e rimmed, and in places consumed, by spinel and ilmenite. Symplectites Gymj óf spinet and cordieriæ consume primary sillimanite, but a¡è overgrown by secondary iilrirna"itè.'nims oi tø2 sillimanite also often separate cordierite from primary spinel and itmenite. Opaque minerals a¡e ilmenite. Width of f,reld of view is 2mm, The high strain 52 fabric is defined by garnet, sillimanite and ilmenite, with rare,late biotite. Ilmenite often occurs as elongate inclusions within secondary garnet and sillimanite, and fibrous sillimanite overgrows primary sillimanite and forms along recrystallised grain boundaries of cordierite. In the lower strain boudin, fibrous sillimanite along cordierite and spinel grain boundaries appears to be synchronous with the sillimanite which defines the shear zone fabric. 6.3.4 Orthopyroxene-bearing assemblages In sample 96448&C, an coarse primary garnet-orthopyroxene-plagioclase- kspar-quartz assemblage is overprinted by a mylonitic shear fabric. The orthopyroxene and garnet are often in contact, and biotite, orthopyroxene and qtJartz form inclusions within garnet. An anastomosing but nearly pervasive shear fabric is 138 Chapter 6, Fox Ridge defined by fine grained traits of recrystallised garnet and orthopyroxene, associated with oriented biotite and dynamically recrystallised quartz and feldspar (Fig. 6.2(k)). Biotite often envelops the M2 garnet and orthopyroxene, suggesting much of it formed relatively late during 52 development. 6.3.5 Mineral Chemistry Mineral analyses were obtained using a JEOL 733 electron microprobe at the Centre for Electron Microscopy and Microstructural Analysis at the University of Adelaide, using the electron-dispersive spectra (EDS) system with an accelerating voltage of 20 keV. Representative analyses of minerals are given in Appendix 3. Garnets belonging to the M1 assemblage on Fox Ridge are invariably almandine rich, with Xps ranging from 0.7 to 0.8. There is a moderate but generally consistent zonation in the garnets, becoming more Fe-rich towards the rim. A more dramatic increase in Xps at the rim is probably due to later re-equilibration, as well as secondary garnet overgrowths. A typical primary garnet zonation for Fox Ridge metapelites is shown in figure 6.4. Secondary garnets associated with the high stain 0.8 Xalm 0.6 0.4 0.2 Xpyr Xgrs 0.0 Xsps 0.20 0,40 0.60 0.80 1.00 r.20 biotite mm cord Fig. 6.4. 7-oning within typical Ml garnet, sample 964_gA, Fox Ridge. fabric garnets have higher almandine and lower pyrope contents than primary garnets, and also show a minor amount of iron enrichment towards the rim. r39 Chapter 6, Fox Ridge Cordierite has variable composition in the Fox Ridge pelites. Large porphyroblastic M1 cordierite in 964-8C has the highest Xps of 0.43, and is slightly zoned with an increase in Mg towards the rim. However cordierite porphyroblasts in 964-84 show the opposite zonation. Within samples, porphyroblastic cordierite is generally more Fe-rich than later recrystallised cordierite, although there is often significant variation in composition. Biotite composition is highly variable between specimens, with Xpg ranging from 0.3 to 0.6, and within samples Xp" can also vary by as much 0.1. Tio2 content in biotite is also variable, between 2 and 6wt%o, with the higher Ti contents generally in the more Fe-rich biotites. Most biotites also contain minor amounts of Ca. Plagioclase composition varies from An23 to Ang3, although generally most plagioclase has compositions between An¿o-zo. Fe-rich pelites have relatively sodic plagioclase, whilst garnet-orthopyroxene gneisses have the most calcic plagioclase. Within specimens, plagioclase associated with cordierite in the consumption of garnet generally has a more calcic composition than plagioclase in the matrix. Spinel is hercynite-rich, with the remainder being dominantly spinel, gahnite and chromite. In Zn- and Cr-poor spinels, which mainly occur as inclusions within sillimanite and garnet, there is between 0.8 and 0.9 Fe2+ per formula unit, with a consistently higher Xps than associated garnet. Conoded spinel in the matrix however, has va¡iable amounts of Cr and Zn, and in an extreme case have l3wt%o ZnO andïwt%o Cr2O3. Orthopyroxene in the garnet-orthopyroxene gneisses have the composition of hypersthene, with an Xpe of 0.4-0.5. Primary orthopyroxenes show zonation, with a rimward decrease in Fe2+ content, and late orthopyroxenes in the shear fabric are more magnesian than the large M1 porphyroblasts. Al content is highest in the cores r40 Chapter 6, Fox Ridge of the primary orthopyroxenes, with 0.16 413+ performula unit, and this decreases towards the rim. Small M2 orthopyroxenes generally have less than 0.1 Al atoms per formula unit. 6.4 Mineral parageneses and a P-T history for Fox Ridge 6.4.1 Interpretation of mineral parageneses The sequence at Fox Ridge is useful in the fact that the dominant high strain fabric always contains temporally related assemblages, and this allows correlation of the relative timing of assemblages between rocks of differing bulk composition. In considering the P-T evolution of the Fox Ridge metapelites, phase relationships in the KFMASH and KFMASHTO systems, as discussed in chapters 3 and 4, will be considered. In most of the samples from Fox Ridge, the peak assemblage is dominated by garnet and sillimanite, occasionally with cordierite and/or spinel. The stabilisation of cordierite at the expense of garnet and sillimanite, which in some cases also involved spinel and ilmenite, implies lower pressure and./or higher temperature conditions through the crossing of the univa¡iant reaction garnet + sillimanite = cordierite + spinel + quartz The earliest assemblage defining the high strain fabric was defined by garnet and cordierite with or without sillimanite and ilmenite, suggesting that isobaric cooling or compression accompanying cooling accompanied early shear development. In a number of samples from Fox Ridge, ilmenite and rutile are abundant phases, and therefore the P-T evolution should also be examined using grids for the KFMASHT system (see Chapters 3 and 4). A KFMASHT pseudosection, for pelitic compositions approximating those of typical pelites from Fox Ridge, and constructed using the KFMASHT grid of clarke et al. (1989) is shown in figure 6.5. This pseudosection shows the apparent decompression required to overprint the Mr t4L Chapter 6, Fox Ridge assemblage with the M2 assemblage. Progressive stabilisation of sillimanite at the expense of cordierite and spinel, and then biotite at the expense of garnet and cordierite, implies that a near-isobaric cooling path during shear development is appropriate (Fig. 6.5). Again, the successive assemblages defining the high strain fabric record a consistently near-isobaric cooling path. KFMASHT+ ilm +qz + ksp + L P gt sill bi ö gt cd sill ll bicd sp s ill T râ' W M1 assemblage M2 cordierite-bearing assemblag€ Fig. 6.5 Qualitative P-T pseudosection for aluminous assemblages at Fox Ridge, based on KFMASH univa¡iant e4uilibria (afær Clarke and Powell, 1991). Arrow shõws apparent cõoüng path during high st¡ain deformation. 6.4.2 Pressure-temperature conditions of metamorphism Calculations of the physical conditions of peak metamorphism at Fox Ridge were made using Thermocalc (Powell and Holland, 1988) using an updated version of the internally-consistent thermodynamic dataset of Holland and Powell (1990), and the Thermocalc output is given in Appendix 4. Average pressure calculations using the inferred M1 assemblage garnet-sillimanite-cordierite-ilmenite-rutile-quartz for sample 96458 for a temperature of 750oC give an average pressure of 5.8 + 0.4 kbars for an aHzo of 0.4, and 6.2 + 0.4 kbars for an alH2o of 0.6. For the M1 garnet- sillimanite-cordierite-ilmenite-quartz-plagioclase-rutile assemblage in sample 8A, average pressure/temperature estimates for low H2O and COz activities are 778 + I37 142 Chapter 6, Fox Ridge oC and 5.7 + 0.9 kbars. Higher water activities give higher pressures and similar temperatures, but with larger standard deviations. For a temperature of 775oC, average pressures of 5.8 + 0.4 kbars were obtained, with average pressures being up to 0.4 kbars higher for higher water activities, similar to the results for sample 58. For a pressure of 6.1 kba¡s, calculated average temperatures were 738 + 49o C. Calculations on the assemblages which define the D2 high strain fabric on Fox Ridge consistently give pressures in the region of 4.5-5 kbars at temperatures of 650- 700oC. Average temperature calculations on the sheared garnet-orthopyroxene assemblage in sample 964-4C give temperatures of 667 + 57oC at 5 kbars, whilst average pressure calculations for the high strain assemblage in sample 964-8C give pressures of 4.6 t 0.51 kbars at a temperature of 700oC. Average pressure- temperature calculations give values of 636 + 57oC and 4.8 t 0.8 kbars for sample 964-58, 674 + 68oC and 4.yt 0.9 for sample 964-84 and 704 t 67"C and 4.4 + 1.0 kbars for sample 964-8C. 10 P M1 assemblage (kbars) / / gt-opx, Sen and Bhattacharya, 1984 8 Average pressure estimates / using Thermocalc 6 Recrystallised assemblage / gt-opx, Sen and Bhattacharya, 1984 gt-opx, Harley and Green, 1982 4 il Average P-T estmates f T ("C) using Thermocalc 2 400 500 600 7w 800 900 1000 Fig. 6.6. Sumnury of P-T estimates for the peak and recrystallised assemblages at Fox Ridge, based on garnet-orthopyroxene geothermobarometry of Sen and Bbattach uya (1984) and Harley and Green (1982) and using the average pressure-temperature approach of Powell and Holland (1985, 1988). Error ba¡s indicaæ standard deviations at95Vo confidence for average pressue-temperature calculations. Garnet-orthopyroxene geothermobarometry on the peak and recrystallised garnet-orthopyroxene assemblages in sample 964-48 is summarised in Fig. 6.6. For t43 Chapter 6, Fox Ridge the coarse assemblage the thermometer and barometer of Sen and Bhattacharya (1984) suggests peak conditions in the region of 750-800oC and 6-6.5 kbars, which is broadly consistent with Thermocalc results. Application of the same geothermometer and geobarometer, as well as the geobarometer of Harley and Green (1932) to the high strain assemblage, in which garnet and orthopyroxene were finely recrystallised, gives temperatures of around 650oc at pressures of around 5.5 kba¡s. 6.4.3 Summary Peak metamorphic conditions in the Fox Ridge region appear to have been between 750 and 800oC, for pressures of 6-6.5 kbars, reflecting the expected slightly lower grade in comparison to Jetty Peninsula and the Porthos Range. This is consistent with the relative proximity of Fox Ridge to the regional amphibolite- granulite transition in the NPCM. The immediate evolution following peak metamorphism is not well constrained, with no evidence existing for isobaric cooling. The development of cordierite coronas between garnet and sillimanite suggest that decompression occurred, but there is no constraint on whether or not this decompression occurred at continuing high temperatures, or whether the cordierite coronas reflect a thermal pulse at lower pressures. Following the development of the cordierite coronas, cooling occurred synchronous with the initiation of high strain along a steeply dipping shear zone. Early within the high strain development, temperatures were in the vicinity of 650oC at around 5 kbars, suggesting that the magnitude of decompression following peak metamorphism was in the vicinity of l-2 kba¡. Cooling continued during the high strain history, with sillimanite and biotite progressively overprinting garnet and cordierite, probably reflecting cooling to amphibolite facies temperatures under near isobaric conditions. The final ultramylonites are defined by biotite, suggesting that temperatures were well below 650oC at the end of high strain deformation. 144 Chapter 6, Fox Ridge 6.5 Correlation of metamorphism \ilith other areas The structural and metamorphic history of the NPCM and surrounding regions still has insufficient and often ambiguous age constraints, and as a result there is still no definite geochronological framework within which to correlate the evolution of the Fox Ridge metapelites. However, the metamorphic evolution of the Fox Ridge metapelites can be related to the history documented for other fragments of the NPCM, such as Jetty Peninsula and Trost Rocks, and can be tentatively extrapolated further to the Mawson and Prydz Bay coasts. The peak assemblage at Fox Ridge is assumed to be the same peak 1000 Ma event found throughout MacRobertson Land and prydz Bay (Tingey,l9g2; Black et al.,1987; Manton et al., 1992). The development of cordierite coronas around garnet is a feature common throughout this region, and has traditionally been attributed to near isothermal decompression from peak conditions. However, as was stated in Chapters 3 and 4, a biotite fabric, usually with sillimanite, developed temporally between the peak assemblage and the cordierite-bearing coronas and symplectites at Jetty Peninsula and Trost Rocks, yet this biotite fabric was generally absent from Fox Ridge. The reasons for the lack of biotite between the peak and cordierite-bearing assemblages at Fox Ridge may be either: 1) a different PT evolution to Jetty Peninsula, with less cooling between the peak assemblage and cordierite coronas; 2) a different bulk composition, particularly with respect to the activity of HzO; or 3) a lack of deformation at lower-T conditions, inhibiting the development of a biotite fabric. The cordierite coronas and cordierite fabric at Fox Ridge can be directly correlated with the secondary cordierite-bearing assemblages on Jetty peninsula and Trost Rocks, although it is less clear whether these cordierite coronas can necessarily also be conelated to the decompression textures described along the Prydz Bay coast from the Reinbolt Hills to the Rauer Islands (e.g. Harley, 1988;Nichols and Berry, 145 Chapter 6, Fox Ridge 1991; Fitzsimons and Harley,l992a; Dirks et al, 1993). on Jetty Peninsula, the major east-west trending shear zone formed late during M2, as it often recrystallised cordierite coronas, but cordierite-spinel symplectites also overgrow the shear fabric. Similarly, on Fox Ridge, the east-west shear zones recrystallise cordierite coronas, but are also defined by cordierite. This supports the assumption that the east-west trending shear zones on Fox Ridge are the same generation as, or are at least temporally related to the major Jetty Peninsula shear zones. As such, the age of the high strain zones and associated cooling is likely to be c.940 Ma (Section2.6). The timing of the cordierite-bearing coronas which are overgrown by the high strain fabric on Fox Ridge is not constrained, other than that it postdates D1 and the coronas are recrystallised by Dz. However, on Jetty Peninsula, the cordierite-spinel symplectites formed within pull aparts along the L2lineation in the hinge of the D2 fold, during early stages of D2 deformation, and in places these coronas were then recrystallised by the D2 high sfain fabric. As the cordierite-coronas on Fox Ridge are likely to have formed as part of the same event as the cordierite coronas on Jetty Peninsula, it is therefore considered tikely that the cordierite-bearing coronas at Fox Ridge are closely related in time to the subsequent high suain fabric. 6.6 Discussion The petrographic relationships in the metapelitic shear zones from Fox Ridge suggest that region cooled from garnet-sillimanite-cordierite granulite facies to amphibolite facies conditions (defined by the development of hornblende- cummingtonite assemblages in mafic rocks) within the timescale of a single deformational event, and therefore this cooling was likely to have been rapid. The steep orientation and apparent reverse shear sense of the east-west trending shear zones on Fox Ridge and Jetty Peninsula, along with the upright nature of associated fold structures, are consistent with having formed during crustal thickening as opposed to extension, and therefore a mechanism must be found to account for the t46 Chapter 6, Fox Ridge isoba¡ic cooling of the terrain synchronous with convergent deformation. If the major source of heat for granulite metamorphism was the advection of heat in the form of magmas, overprinting a broader thermal anomaly relating to an increased conductive geothenn, then clearly rapid cooling may be possible. Significantly, the isobaric cooling appears to have followed at least some exhumation of the terrain, as it occurred subsequent to the development of cordierite coronas around garnet. It is conceivable that the high strains associated with D2 are associated with a renewed thermal perturbation due to magmatic intrusion, with a corresponding thermal weakening of the crust. With no evidence of extensional deformation following the upright folds and shear zones, it appears reasonable to assume that all subsequent uplift was primarily erosion-driven. The geodynamic implications of the metamorphic and structural evolution of the region is discussed in more detail in following chapters. 147 Chapter 7, Geodynamic and therm^al considerations Chapter 7 Some geodynamic and thermal considerations in granulite terrains 7.1 Introduction Large portions of the continental crust exposed at the Earth's surface are dominated by regional Precambrian granulite terrains, and yet many aspects of the processes which led to the formation of these tenains remain unclear. This chapter is a discussion of two important questions relating to the evolution of granulite terrains, which will be addressed in light of the metamorphic and structural evolution of the northern Prince Charles Mountains presented in previous chapters.. The first question to be addressed concerns the duration of thermal perturbations which result in granulite metamorphism. A convenient relative me¿¡sure is provided by the timescale of lithospheric deformation during orogenesis, and the question is whether the timescale of metamorphism is shortlived (transient) or on a comparable timescale to lithospheric deformation. It is clea¡ that the answer to this question has significant implications for the interpretation of metamorphic P-T-t paths, and for understanding the relationship between metamorphism and deformation. In particula¡, it is important for understanding the significance of the near-isothermal decompression paths which have been documented for many granulite terrains. The second question concerns our interpretation of the structures preserved in granulite terrains, which is often impeded by a lack of understanding of the way the mid to lower crust deforms under such thermally perturbed conditions. Such 148 Chapter 7, Geodynamic and thermal considerations structural interpretation is fundamental for the development of geodynamic models for the formation of ancient granulite terrains. 7 .2 The duration of granulite metamorphism Many regional granulite facies metamorphic terrains are interpreted to have formed in an environment of crustal thickening (e.g. Park, 1981, Harley, 1989; Bohlen, 1991), associated with convergent plate boundaries. The thermal evolution of tenains in these settings has in general been understood in terms of two simplified end-member models (Mezger et al., 1990). The first involves only the effects of crustal thickening, which serves to increase the heat production within the thickened crustal column, whilst the second involves the emplacement of mantle-derived magmas at various levels in the crust before and during thickening. It has been shown by a number of authors (Loosveld and Etheridge, 1990, Sandiford and Powell, l99l) that the thermal effects of crustal thickening are largely dependent on the response of the mantle part of the lithosphere to lithospheric deformation (Fig. 7.1(a)). However, the potential temperature calculations by Sandiford and Powell (1991) show that even with significant thinning of the lithospheric mantle during crustal thickening, mid crustal temperatures are unlikely to significantly exceed 600oC (Fig. 7.1 (c)). In addition, Sandiford and Powell (1991) show that to get temperatures of 700oC in the middle of thickened crust due to thinning of the lithospheric mantle, unrealistic Moho temperatures of 1200oC need to be attained (Fig.7.1(b)). It therefore appears necessary to invoke advection as a dominant process in heating the mid crust to granulite facies conditions, and this implies that the thermal perturbations resulting in mid-crustal granulite metamorphism may well be transient (Loosveld and Etheridge, 1990; Sandiford and Powell, L991; Scrimgeour and Sandiford, 1993). If the thermal perturbations are short on the timescale of orogenesis, then this raises the question of the significance of the large scale changes in pressure under near-isothermal conditions, which have been infened in many granulite terrains (e.g. Harley and Hensen, 1990; Clarke and Powell, 1991). t49 Chapter 7, Geodynamic and thermal considerations (a) A crust mantle lithosphere B -> high heat lo*ô f"=fr'1 heat I flux (b) (c) 400" 600" 200" 300" 400" 2 2 800" f 500" 1 000" 1 1 600" 1 200" 700" P=1 P ='ll2 1 21 2 f f c c Fig.7.1 (a) Schematic illustration of the^ lilhospheric-scale deformation geometries discussed in this paper- ' (after Sandiford & Powell, 1991), where /c and fl åre rneåstues of the vertical strain on the crust and lithosphere, respectively A. The initial lithosphere prior to deformation, which is assumed to be in thermal ùuitl¡rium. B. Ahomogenousthickeningdeformation(onthescaleofthelithosphere),suchthát/c=fI>1). C. A deformarion in which crustal thickening is accompanied by thinning of the subcir¡stal lithoÁphere such that /c > 1, > fl. '"fl*,," 250"C. ry during crustal thickening shown in (a), where deformation geometries capable of generating (Q /c-fl. diagram similar to @), except showing the potential temperaturcs at half crustal depth for the same lithospheric deformation geometries. 150 Chapter 7, Geodynnmic and themutl considerations A recognition of the necessity of transient advective processes in granulite metamorphism has important implications for the way we view P-T paths. Over the past ten years, it has been popular to construct P-T paths as smooth clockwise or anticlockwise 'loops' (e.g. England and Thompson, 1984), which implies that the thermal evolution of the terrain is intimately related with, and occurs on a similar timescale to, the changes in pressure induced by the lithospheric deformation and exhumation. However, recognising that the growth and destruction of mountain belts (which is the most viable means of bringing about changes in pressure) generally occurs on a timescale of >40 Ma (England and Thompson, 1984; De yoreo et al., 1991), then a single P-T 'loop' of this duration is inconsistent with heat sources that are primarily advective in nature. If the perturbation was due to advection of magmas, then the timescale of the thermal event would be likely to only 0.1 - 5 Ma (Huppert and Sparks, 1988; De Yoreo er al., 1991; Sandiford et al., 1992). A feature of the evolution of advectively heated terrains is ttrat cooling to the background geotherm would be expected to initially be nearly isobaric (Sandiford et al., lgg2). Some support for the notion that a number of granulite facies terrains have undergone prolonged heating and cooling cycles is provided by an assessment of radiogenic isotope blocking temperatures which have been used as evidence that some granulite events occurred on the timescale of 100 Ma or more (Harley and Black, 1987; Humphries and cliff, 1990; oxburgh, 1990). However, precise u-pb geochronology of an apparently long-lived granulite event in the Archean pikwitonei Domain of North America by Mezger et al. (1990) has revealed a metamorphic evolution which appears to represent a combination of a number of distinct thermal events during a single long-lived tectonic event, with only short intervals of garnet and zircon growth. Significantly, Mezger et al. (1991) also apptied precise geochronology to the Mid Proterozoic granulites of the Adirondacks, and showed that the timing of the metamorphic peak varied across the terrain, and attributed this to a 151 Chapter 7, Geodynamic and themtal considerations series of episodes of magmatism, which resulted in different parts of the terrain experiencing high-grade metamorphism episodically over a period of 150 Ma. An example of a transient thermal event overprinting a smooth P-T 'loop' is seen in the P-T-t model proposed for the Prieska Copper Mines in the 1100 Ma Namaqua-Naral Belt of South Africa by Cornell et al. (1992) (Fig. 7.2). In rhis model, which is relatively well constrained by geochronological data, a typical long- lived collision-related clockwise P-T path as proposed by England and Thompson (1984) is interrupted by thermal excursions related to intrusion of magmas, resulting in a relatively short-lived granulite event at c.1100 Ma. As magmatic intrusion is likely to be required to reach the temperatures recorded for most low to medium pressure granulite events (Bohlen, 1991; Sandiford and Powell, 1991), then this may 7 1215 P 6 M (kbar) 5 M2 4 M3 1080 1104 3 2 1 12 0 0 200 4OO 600 800 T I.t9.7.2. Pressure-temperature path, including age consEaints in Ma for the Prieska Copper Mines, Namaqua-Natal Belt, South Africa showing a thermal excu¡sion at c.1100 Ma overprinting a typical clockwise P-T paths related to continental collision. (Adapted ftom Cornell et. al. (1992)) represent a typical evolution for such terrains. In higher pressure terrains, particularly those above l0 kbars, the argument that transient processes must necessarily be involved to account for the temperatures necessary for granulite metamorphism is not so convincing, due to the ability of rocks at these depths to be heated to temperatures above 700oC by heat production within thickened crust and conductive heat flow ts2 Chapter 7, Geodynarnic and thermal considerations from the mantle (e.g. England and Thompson, 1984), particularly if there is appreciable thinning of the mantle part of the lithosphere (Fig. 7.1(b); Sandiford and Powell, 1991). If, as suggested by the discussion presented above, advection of heat is a necessary requirement in low to medium pressure granulite terrains, then the significance of pressure changes at elevated temperatures needs to be reassessed in light of a potentially transient thermal evolution. In the next section, the implications of near-isothermal decompression paths in granulites are discussed in the light of these considerations. 7.3 Geodynamic implications of decompressional P-T paths in granulite terrains As discussed in the previous section, the likely transience of thermal perturbations resulting in granulite metamorphism raises questions about the significance of P-T paths which imply large scale changes in pressure at high temperatures. Of particular importance is the significance of near-isothermal decompression paths following the peak of metamorphism, which are widely considered to be characteristic of many granulite terrains (e.g. Harley, 1989). Examples of such terrains include the Musgrave Block (Clarke and Powell, 1991), the Limpopo Belt (Windtey et ø1, 1988), the Aldan Shield (Perchuk et aI, 1985), the Grenville Province (Nantel and Martignole, 1978) and East Anta¡ctica (Harley and Hensen, 1990). Evidence for such decompression is often in the form of coronas and symplectites which appear to have formed under essentially static conditions, or associated with upright structures, with no evidence of accompanying extensional deformation. The interpretation that these textures record significant denudation without contemporaneous extensional deformation implies that either, (1) that partitioning of crustal strain during convergence allowed local, rapid exhumation of rocks coincident with crustal thickening, (2) following the waning of the forces 153 Chapter 7, Geodynamic and thetmnl considerations driving compression, the cnrst was sufficiently strong to support elevated topography, and that erosion occurred at a rate significantly faster than the rate of thermal relaxation, (3) that the bulk crustal extensional deformation responsible for denudation was partitioned in such a way that large areas experienced no penetrative strain, or (4) that the P-T history has been incorrectly interpreted. V/ith regard to possibility (1), evidence from modern mountain belts such as the Himalayas, in which granulite facies rocks are currently exposed at the surface of the same mountain belt in which they formed, provide indisputable evidence that granulites can be rapidly exhumed during active crustal thickening. For example, in the Nanga Parbat region of the Pakistan Himalaya, Zeitler et al. (1993) document rocks which were metamorphosed to cordierite - sillimanite - K-feldspar grade at 6 kbars aÍ.4.5 - 3.3 Ma (U-Pb monazite), before undergoing decompression melting during rapid exhumation to the surface where they are currently exposed in the hangingwall of a major thrust. However, such examples are confined to relatively n¡urow orogenic wedges, which generally only have a width of only 50-100 km, and in which granulites are only exposed as thrust bounded windows. In broad regional terrains such as East Antarctica, in which granulite exposures are of the order of several hundreds of kilometres in width, evidence for such rapid exhumation during convergence is lacking. In order to assess possibility (2) it is necessary to explore some aspects of the mechanics of lithospheric deformation. Most granulite facies metamorphic terrains have been interpreted to have been produced in thickened, thermally perturbed crust following continental collision (e.g. Bohlen, 1991). In such a collisional event, the crustal thickening is limited by the increase in gravitational potential energy stored in the belt which must eventually balance the forces driving compression (Molnar and Lyon-Caen, 1988, Zhou and Sandiford, 1992). When this critical thickness is reached, there can be no further increase in crustal thickness, and further convergence r54 Chapter 7, Geodynarnic and thermal considerations will be accommodated by the range building outwa¡ds to form an elevated plateau (Molnar and Lyon-Caen, 1988). With the waning of the forces driving convergence, the mountain belt may either undergo extensional collapse if the crust is not strong enough to resist the horizontal buoyancy forces (England, 1987; Zhou and Sandiford, 1992) or it will be denuded by erosion. The regional metamorphism of the middle crust to granulite conditions (c.700- 800"C) across a broad orogen implies high Moho temperatures and hence a thermally weakened lithosphere, which will be susceptible to extension in response to horizontal buoyancy forces (England, 1987). Indeed, England (19s2) and Zhou and Sandiford (1992) demonstrate the relationship between lithospheric strength and Moho temperature, and show that extensional strain rates will be achieved in response to gravitational buoyancy forces following the relaxation of the convergent driving forces providing that Moho temperatures are above 650-700"C (Fig. 7.3(a)). There are considerable uncertainties attached to such calculations, but calculations by Zhou and Sandiford (1992) have shown that even if the activation energy of power law creep for the strength controlling minerals in the lithosphere is increasedby l\Vo to create an implausibly strong crust, thickened crust with mid-crustal temperatures of above 500oC is likely to deform at significant strain rates in response to horizontal buoyancy forces (Fig. 7.1(c), 7.3 (b)). This would rherefore appear ro imply rhar scenario (2) is also inappropriate, particularly as temperatures in the mid-crust during granulite metamorphism are likely to be significantly higher than 500oc. Therefore, it seems that two possibilities remain for the interpretation of textures that imply near-isothermal decompression of a regional low to medium pressure granulite terrain in the absence of evidence for extensional deformation. (1) The interpretation of the reaction textures in terms of their pressure-temperature- time evolution is incorrect, and the lower pressure assemblage may reflect a later 155 Chapter 7, Geodynamic and thermal considerations (a) (b) 2 2 1.5 1.5 +1 0 1 f | fl1 -1 0.5 0.5 ln e (s't ) è = 1o' lr¡rt 0 0 1 1.5 2.0 1 1.5 2.0 fc fc Fig. 7.3 (a) Tbe fc-fl plane contou¡ed for the instantaneous strain rate of a mountain belt undergoing extension in reslþnse to the relaxation of the forces driving convergence. Contours a¡e for ln (e) in units s-I. The inslantaneous shain rate is calculated numerically by equating tbe buoyancy force witb the vertically integrated str,ength assuming the same potential tbermal sfucture as that assumed for Fig. 7.1. The shaded area indicates the region in which the exþnsional stain rates arc greater than l0-2M8 I. (adapted from Zhou and Sandiford, Le92). (b) The fc-fl plane contoured for va¡iations in the the vertically integrated strength of the lithospbere, where the cbange in strength is Biven by a reduction (-lÙVo) and increase (+LÙVo) in the activation energy of power law creep for quartz and olivine, assuming the sa.rre potential therrral structu¡e as that assumed for Fig. 7.1, and an extensional shain rate of 10-zM¿Í'. This shows that for thermal shuctures which allow high temperatures in the mid-c¡ust and appreciable lower crustal melting (Path C, see Fig. 7.1), then even an implausibly Jtong lithosphere will defonn at at least 10-'Ma-' if the conve gent driving forces are no longer acting on thè orogen (adapted f¡om Zhou and Sandiford, 1992) 156 Chapter 7, Geodynamic and thermal considerations thermal pulse or an entirely separate event, ie. the decompression is not near- isothermal. (2) Bulk crustal extensional strains are strongly partitioned, which is perhaps surprising in such thermally perturbed mid crust. This raises the question of where in the crust the strain is occurring. One possibility is that decoupling of ductile and brittle zones in the crust could occur, so that vast regions of the crust may not record extension that only occurred through processes such as ductile flow of lower crust (see section 7.4) In summary, isothermal decompression seems unlikely to be driven primarily by erosion in regional low to medium pressure granulite terrains. Therefore, in cases where evidence for rapid tectonic exhumation is absent, overprinting thermal events at progressively lower pressures may be a more realistic interpretation. The alternative possibility is based on the notion of partitioning or decoupling of strains through the crust, and a possible mechanism of strain partitioning is discussed in the following section. 7.4Lower crustal channel flow parameters in granulite terrains In our attempts to understand the evolution of granulite terrains, a long- standing problem has concerned the significance of the horizontal structures which dominate many terrains. In general, these structures have been interpreted as a record of the processes that produced the orogenic-scale architecture, and in particular, debate has focussed on the question as to whether they reflect compressional or extensional deformation (e.9. Park, 1981; Sandiford, 1989), with the assumption that such horizontal stn¡ctures must in some ways reflect the displacements imposed at the boundary of the orogenic zone. An alternative viewpoint relating to horizontal structures in granulites is that they may record evidence of ductile flow of thermally perturbed mid to lower crust as a response to the evolving lithospheric deformation geometry, and that this may not 157 Chapter 7, Geodynamic and thermal considerations be related to the overall kinematics of the orogenic zone. It has long been recognised that regions of high topography with deep isostatic roots will tend to spread laterally and place adjacent lithosphere under horizontal compression (England, 1987; Sandiford, 1989), but the processes by which Moho relief is eliminated have remained unclear. Observations relating the movements of the upper crust relative to the lower crust imply that brittle upper crustal deformation may be decoupled from more uniform deformation in the low viscosity middle to lower crust (Mooney and Meissner, 1992), and that therefore the lower crust represents a weak zone during orogenesis, at least at elevated temperatures. Indeed, Bird (1991) proposed a model of lateral extrusion of thermally weakened lower crust by a process of Poiseuille (planar channel) flow, which removes crust from beneath mountains and in doing so, smooths and levels the topography. Bird applied this principle to modern orogenic belts such as Tibet and the Basin and Range Province, and suggested that such channel flow should provide a geologically significant process in smoothing the Moho and hence flatæning topography in orogens where the Moho temperature is elevated (>500oC) for Moho relief with a wavelength of 100-200km. If this is the case, then terrains undergoing regional granulite facies metamorphism, which must have Moho temperatures approaching 900-1000oC, should record the effects of planar channel flow. This section examines the possible role of lower crustal channel flow during the formation of granulite terrains, and whether its effects can be recognised in ancient granulites such as the Proterozoic terrain of East Antarctica. 7.4.1 Evidence for channel flow in modern terrains Studies of the effect of lateral flow in smoothing Moho relief have so far only concentrated on modern orogenic belts, where an explanation is required for unexpectedly smooth Moho relief or surface topography. Clearly, such tenains in which observed structures can be related to well constrained tectonic settings, provide the best test for whether the lower crust flows when the geotherm is sufficiently perturbed. 158 Chapter 7, Geodyrwmic and thernal considerations Basin and Range Province, USA. In the Basin and Range Province of western North America, deep seismic profiles have shown that the present crustal thickness is very uniform with a flat Moho and low surface topography (Wernicke, 1985; Thompson and McCarthy, 1986). This observation is surprising as the magnitude of Tertiary upper crustal extension varies significantly from l5Vo to up to 3007o across the terrain (Gans, 1987). In order to explain the present crustal geometry, Gans (1987) proposed a model in which the brittle upper crust was decoupled from the ductile lower crust during extensional deformation, allowing material from beneath less extended upper crustal domains to flow laterally beneath more highly extended domains. However, Gans (1987) regarded channel flow alone as insufficient to account for the present geometry, and he proposed that channel flow was accompanied by considerable magmatic underplating of the crust. More recently, Block and Royden (1990) and Kruse et al. (1991) proposed that lower crustal flow alone could account for the uniform crustal thickness of the Basin and Range. Similarly, Bird (1991) proposed that extension in the Basin and Range Province was initiated by substantial heating of the Moho, and that this enabled lateral extrusion to occur simultaneously with, and following, upper crustal strain. This negates the need to invoke voluminous additions of magma from the mantle to produce the current flat Moho beneath the Basin and Range. Europe Deep seismic reflections from beneath central and western Europe provide further constraints on the potential role of lower crustal flow in flattening Moho relief. Meissner et al. (1987) observed that the crust beneath the Variscan and Caledonian regions of Europe has a remarkably uniform thickness with no significant Moho relief, yet evidence exists that they formed during crustal thickening events which necessarily would have required deep crustal roots. They proposed a timescale of 50- 159 Chapter 7, Geodynamic and thermal considerations 300 Ma for flattening of the Moho, constrained by the continuing existence of roots beneath the Alps and Pyrenees. However, this lower limit is perhaps unrealistic as a global constraint in view of the fact that the Alps have undergone considerable deformation in the last 50 Ma, and in addition the thermal and rheological structure of the Alpine crust is not necessarily typical of all mountain belts. On the basis of observations from European crustal seismic profiles, Meissner et al. (1991) suggested that the base of the crust must be continuously re-equilibrating after orogenic collapse, and coupled the observation of the flat Moho with the existence of subhorizontal seismic reflectors in the lower crust. Ratschbacher et al. (1991) proposed that the brittle upper crust was decoupled from the ductile lower crust during lateral extrusion late in the Alpine Orogeny, and that the lateral extrusion was in a direction orthogonal to the direction of compression. Tibet The Tibetan Plateau is characterised by having unexpectedty flat topography over anomalously thickened crust, and Bird (1991) proposed that this flat topography is a direct result of lateral extrusion of lower crust beneath Tibet to flatten the Moho. He also suggested that lateral extrusion could provide a suitable mechanism for the northward expansion of the plateau. Similarly, Zhao and Morgan (1937) suggested that a very fluid lower crust acted as a hydraulic reservoir which regulated the height of Tibet during the injection of Indian crust. Applying the compression / delamination model proposed for the orogen by England and Houseman (1989), Bird (1991) suggested that Moho temperatures beneath Tibet probably approach 1000oC, and therefore planar channel flow would seem a likely process in the thermally weakened lower crust. 7.4.2Honzontal structures in the mid to lower crust Horizontal structures in exhumed granulites 160 Chapter 7, Geodyrumic and thermal considerations A characteristic feature of granulite terrains exhumed from the mid to lower crust is the abundance of recumbent structures and horizontal layer parallel fabrics and, in general, these horizontal structures have been interpreted to reflect the kinematics operating across the deforming zone. For example, Park (1981) attributed these structures primarily to thinned-crust collision or mantle decoupling during collision, whilst Sandiford (1989) proposed extensional collapse of thickened crust as a more likely mechanism. However, their broad similarity throughout granulite terrains worldwide suggests that they may in reality reflect a general response of ductile lower crustal rocks to lithospheric deformation, and that caution should be used in inferring tectonic models on the basis of these structures. If the deformation in the lower crust is decoupled from that in the more rigid upper crust and lithospheric mantle as suggested by Bird (1991), Meissner et al., (1991) and Mooney and Meissner (1992), then these horizontal structures may record the lateral response of relatively inviscid middle to lower crust to relative thickening or thinning of the lithosphere. Horizontal seísmic reflectors in the lower crust In addition to the horizontal structures observed in many exhumed lower crustal terrains, deep seismic crustal profiles have often revealed subhorizontal reflectors through the lower crust, especially beneath regions which have undergone lithospheric deformation during the Phanerozoic (Meissner et al., 1987; Brown et a1., 1987; Meissner et al., 1991; Singh and McKenzie,1993\. Various origins have been proposed for these subhorizontal reflectors, with the most popular being: (1) the presence of free fluids in the lower crust (e.g. Newton, 1991), (2) mafic underplating, layered intrusions and sill-like bodies (e.g. Warner, 1990; Singh and McKenzie, 1993) and (3) horizontal ductile deformation and shearing (e.g. Reston, 1988). In reality, it is likely that all three play some role in explaining the existence of horizontal seismic reflectors in the lower crust. Significantly, however, Mooney and Meissner (1992), note that, for relatively young crust, the minima in crustal viscosity- 161 Chapter 7, Geodynamic and themnl considerations depth curyes coincide with the location of seismic lamellae within the lower crust, and cite this correlation between viscous crust and horizontal reflectors as evidence for regional scale lateral ductile flow of material within the lower crust beneath rigid mid-crust. In addition, Mooney and Brocher (1987) showed that the onset of lamination in the lower crust does not globally conelate with any intracrustal velocity boundaries. This implies that the lower crustal layering is more dependent on crustal viscosity rather than compositional contrasts as would be expected with mafic intrusions, and, coupled with the observation of horizontal fabrics in exhumed granulites, this suggests that the response of the middle to lower crust to lithospheric deformation is dominated by horizontal ductile flow. 7.4.3Importance of the rheological structure of the crust The rheological structure of continental crust will significantly affect the magnitude and nature of any planar channel flow in the mid to lower crust. Many geodynamic models assume a simplified rheological structure of the lithosphere with a quartz-dominated crust and olivine-dominated mantle (e.g. England, 1987), and this would suggest that for a steady-state geotherm, the lowermost crust would be the weakest part of the lithosphere. However, it appears likely that much of the lower crust is mafic in composition (Hall, 1986), and therefore may be stronger than the quartz-rich mid-crust. This has led Wernicke (1991) to propose a'fluid crustal layer' in the mid-crust, which behaves as a relatively inviscid fluid at geological timescales (>0.01 Ma) and subcontinental lengthscales (100-1000km). This fluid crustal layer would experience intense thinning beneath stable upper crust during extension and beneath shortened upper crust during compression, and the base of this layer would experience large-scale, uniform sense simple shear (Wernicke, 1991). Clearly, it is important to consider the rheological structure of the crust before making generalised assumptions about the thickness or crustal level of any channel for ductile crustal flow. Quite possibly, the continental crust will be rheologically stratified in such a way that there may be more than one level at which crust will be undergoing channel t62 Chapter 7, Geodynarnic and thermal considerations flow; for example, flow in the quartz rich middle crust, and then a second layer of flow at the base of the feldspar rich lower crust. 7.4.4 Channel flow in granulite terrains: some expectations Although the potential for lower crustal channel flow driven by pressure gradients arising from orogenic-scale lithospheric deformation geometries is now becoming recognised in modern active tectonic settings, there has been no serious attempt to assess whether models of lateral flow of lower crust can be applied to the evolution of ancient orogenic belts. If this process is relevant to high temperature metamorphic terrains, then this decoupling of upper and lower crustal deformation could have significant implications for our interpretations of structures and reaction textures preserved in the middle to lower crust. Before attempting to apply models of planar channel flow to documented granulite terrains, it is important to first consider the thermal and baric evolution and structural history that would be expected to be recorded in these terrains if they had been affected by channel flow. N ec e s s ary conditions for channel flow Fig.7.4 shows a crustal deformation geometry which has resulted in lateral pressure gradients arising from topographic and Moho relief. For the purpose of this study, it is not important whether this geometry is a consequence of compressional or extensional deformation. What is important is that the lower crust must be thermally perturbed by such processes as thinning of the mantle lithosphere accompanied by the advection of melts (Sandiford and Powell, 1991; Zhou and Sandiford, 1992). Further important parameters to consider when assessing the likely effects of lower crustal channel flow are the wavelength and amplitude of the Moho and surface relief, under the assumption of an initially sinusoidal geometry. Bird (1991) argued that wavelengths in the order of 100 km will most effectively allow lateral extrusion of lower crust. At higher wavelengths, the process becomes more highly dependent on the rheology of the lower crustal lithologies, but if there is a significant perturbation 163 Chapter 7, Geodynnmic and thermal considerations of the geotherm it is still likely to occur at wavelengths as high as 500km. For example, Kruse et at. (1991) suggest that crustal flow occurred over a length scale of 700 km in the Basin and Range, and they suggested that this was possible for lower crustal temperatures of 700-1000oC. At wavelengths significantly lower than 100 km, Moho relief can be supported by the finite strength and flexural rigidity of the mantle lithosphere, and similarly, low wavelength topographic relief can be supported by the flexural rigidity of the upper crust, and this prevents lower crustal flow in these circumstances (Bird, 1991). Relative timing of channelflow An important consideration when looking at the geodynamic and thermal consequences of channel flow is the timing of crustal flow relative to lithospheric deformation. This is intimately related to the thermal structure of the crust during deformation, and the forces driving the lithospheric deformation. In extensional terrains, in which the lithosphere is substantially thinned with subsequent high heat flow from the asthenosphere, the thermal structure of the crust would potentially allow channel flow that initiates during, but could outlast, active upper crustal extension. This is supported by observations from the Basin and Range (Block and Royden, 1990; Bird, 1991), in which there is evidence for elevated temperatures in the lower crust during deformation allowing compensation of ongoing upper crustal thinning by regional scale flow in the lower crust. During convergent deformation, however, it is more difficult to invoke substantial flow coincident with crustal thickening, particularly if there is homogeneous thickening of the entire lithosphere, due to the thermal blanketing effect of a thickened lithospheric mantle. It is likely that no significant flow of the crust would occur during homogeneous lithospheric deformation, as the timescale of the conductive thermal response of the lithosphere is long in comparison with the timescale of active convergence. However a more immediate response may be t64 Chapter 7, Geodynamic and themral considerations achieved if crustal thickening is accompanied by thinning of the mantle lithosphere or by mantle derived magmatism (Fig.7.1; Section 7.2). In a general mechanical model for doubly vergent orogens, Willett et al. (1993) propose that flow of thermally activated viscous lower crust may be largely responsible for the creation of a broad plateau, following crustal thickening. In the case of Tibet, where delamination of the mantle part of the lithosphere is proposed to account for the present elevation by England and Houseman (1989) then it may be that channel flow, as invoked by Zhao and Morgan (1987) and Bird (1991) has only recently been activated as a consequence of increasing Moho temperatures following delamination. Whilst the potential timing of flow remains poorly constrained, an important aspect of the models of Bird (1991) and Willett (1993) is that they suggest that flow can occur whilst the convergence is still occurring across the orogen. Geodynamic and baric consequences of channelflow For the crustal geometry shown in Fig. 7.4, it is clear that the variation in crustal thickness would result in lateral pressure gradients in the mid to lower crust, and as a consequence could potentially drive channel flow. Fig. 7 .4 also shows the instantaneous velocity vectors for points within the crust for such a crustal deformation geometry, showing that bulk flow will be convergent within the initially thicker crust, and divergent within the adjacent thinner crust. The schematic strain ellipses in Fig. 7.4 show that in regions where bulk flow is convergent, the rocks are undergoing extension, whilst in regions where flow is divergent the rocks undergo horizontal shoræning. Consequently, regions within thermally perturbed thicker crust undergoing channel flow may be expected to have relatively planar, horizontal ductile fabrics. The expression of the shortening in the relatively thinner crust may well be the development of upright folds and fabrics, particularly in the region where the crust ceases to flow pervasively. During channel flow, the lower parts of the relatively thicker crust will tend to decompress, whilst the lower part of the thinner crusq under horizontal compression, will tend to move to greater depths. As an isostatic response 165 Chapter 7, Geodynamic and thermal considerations to the lower crustal flow, the surface topography of the thicker crust would decrease, whilst the thinner crust would be elevated, hence flattening the large scale topography. However, the topography cannot be completely annihilated by flow, allowing erosion to act to remove the residual topography * 'l v A,f^ Mo ,1. mantle lithosphere Moho mantle lithosphere fig.7,4 Crusøl strain geometry which would potentially allow channel flow if the therrral structure was appropriate (see text). Darker stippled region indicates a poæntial channel of flow of themrally perturbed mid to lower cmst. Whether this channel occurs within or at the base of the crust is dependent on the rheological structure of the crust. Arrows indicate the instantaneous velocity vectors as a consequence of channel flow. Ellipses represent st¡ains in the regions where flow is divergent (within thicker crust) and convergent (within thinner crust), The lower diagram represent s the crust following flow, with low residual topography Implications for the extent of thickening during orogenesis The recognition of channel flow as a potentially important process may help resolve an apparent paradox presented by many low pressure, high temperature 166 Chapter 7, Geodynamic and thermal considerations terrains. In many of these terrains, strains recorded in rocks exposed at the surface would appeü to be evidence of greater bulk crustal thickening than is revealed by the sum of the metamorphic depth of rocks at the present erosion surface (typically around 15 km) and the present crustal thickness (typically around 35 km). An example of such a terrain is the Anmatjira Range in central Australia, which is a Mid Proterozoic low-pressure (700oC, 4.5 kbar) granulite tenain in which the structural geometries are consistent with intense convergent deformation associated with granulite metamorphism (Collins et al., 1991; Hand, 1993). The absence of evidence for significant post-peak extension suggests a maximum syn-metamorphic crustal thickness of 50 km, and the fact that the region was probably thickened subsequently during the Palaeozoic Alice Springs Orogeny (Collins and Teyssier, 1989), suggests that crustal thickness during granulite metamorphism was significantly less than 50 km. A feasible explanation for the apparent con[adiction of the association of intense convergence with the exposure of low pressure rocks is that the upper crustal strain intense upright structures fabrics HEAT Figure 7.5 Sketch diagram indicating a possible explanation for ùe exposure at the surface of low pressure terrains which preserve evidence for intense convergent deformation but minimal crustal thickening. In this scenario the thermally weakened lower crust is flowing laterally in response to thickening of the upper cmst. Ellipses show the variation in srain between thè upper and lowei cmst. t67 Chapter 7, Geodynamic and thermal conriderations was decoupled from lower crustal strain during compression, and that compressional strain was partitioned into the stronger upper cmst, whilst the lower crust flowed from beneath the thickened upper cnrst (Fig. 7,5). This notion is supported by the extremely perturbed thermal structure of such low pressure-high temperature terrû¡ns (Loosveld and Etheridge, 1990), which implies a thermally weakenetl cn¡st. Implications for metamorpltic P-T paths The flow of lower crust during and following lithospheric deformation in orogenic belts may have significant implicltions for the interpretation of the P-T paths preserved in high grade terrains. Quantification of the pressure changes resulting from channel flow is dependent on a number of factors. The first important factor is the geometry of the region of the crust undergoing flow. In general, if the flow is in response to pressure gradients due to a relatively narrow zone of thickened crust, then the thicker crust will undergo significant decompression, with relatively minor associated compression in adjacent thinner crust. If however, the flow is annihilating Moho relief across a relatively small area of anomalously thin crust, then the baric effects of channel flow will be significant in the thinner crust, and minor in the thicker cÍust. If we take, for example, the case of a region of thickened crust, then simple isostatic calculations would suggest that, if the planar channel extends to the Moho, then flow of lower crust would result in a maximum decompression of rocks nearthe Moho of c.1.5 kbars foreach km of topography which is annihilated by flow. In reality, this value is probably significantly smaller, as this requires the assumption that the Moho in the adjacent thinner crust does not increase in depth as a result of flow. Pressure changes would also be lower at higher levels in the channel of ductile flow. As channel flow is a thermally activated þrocess, it is not appropriate to invoke significant temperature changes during flow. Although quantihcation of the baric effects of lower crustal flow is inhibited by our lack of understanding of the processes involved, it seems likely that if channel flow is a relevant process in granulite metamorphism then it could potentially account for pressures changes at 168 Chapter 7, Geodynamic and thcrmnl considerations high temperatures, particularly when accompanied by relatively horizontal planar stluctures. In overlying more brittle cnìst, the annihilation of much of the topography would be likely to lead to less exhumation, and therefore less decompression during cooling. Whilst there remains uncertainty as to the role of channel flow in granulite terrains, the evidence that it appears to occur in modern terrains, coupled with the existence of planar horizontal fabrics in thermally perturbed crust, support the notion that the lower crust may indeed undergo flow during and following orogenesis. If this is the case, then it provides an alternative explnnation to the existence of horizontal structures in granulites, and should be taken into account when interpreting the metamorphic and structural record in granulite terrains. 7.5 Summary Clearly an understanding of the mechanism of granulite formation and the evolution of granulites within convergent orogens is largely dependent on an understanding of the thermal evolution of these terrains. Any investigations of the metamorphic and structural evolution of low to medium pressure granulites should take into account the likely transience of thermal perturbations relative to the timescale of orogenesis and the difficulties in inferring erosion as a mechanism for near-isothermal decompression. In addition, the potential thermal, structural and isostatic consequences of planar channel flow of thermally perturbed lower crust are a much neglected, but perhaps vitally important consideration when making geodynamic interpretations of granulite terrains. '169 Chapter 8, Discussion Chapter 8 Discussion 8.1 Introduction The Mid Proterozoic granulite terrain of East Antarctica is one of the most extensive regional granulite terrains on Earth, extending for over 1000 km along the Antarctic coastline, with a width which appears to exceed 300 km, although there is considerable uncertainty regarding the original dimensions of the orogen. It is therefore important to understand how such an extensive high grade terrain could have formed, and to account for variations in the metamorphic and structural evolution across the terrain, in order to better develop the concepts relating to the thermal and geodynamic evolution of such terrains discussed in Chapter 7. The metamorphic and tectonic evolution of the granulite terrain of MacRobertson Land has generally been poorly understood in comparison with the terrains of the Rayner Complex of Enderby and Kemp Lands to the west, and the Prydz Bay coast to the east. This has primarily been due to poor outcrop along the Mawson Coast, and the remote and inaccessible nature of the northern Prince Charles Mountains. The tenains of both the Rayner Complex and Prydz Bay have relatively well documented metamorphic histories in which the metamorphic evolution has been interpreted as being dominated by near-isothermal decompression from peak conditions. The relative lack of knowledge of the tectonic evolution of MacRobertson Land has been a major obstacle in tying together the terrains into a coherent model for the evolution of the Proterozoic terrain of East Antarctica, particularly as the northern prince Charles Mountains represents the only exposure which extends significantly inland from the coast to the southern margin of the terrain. 170 Chapter 8, Discussion 8.2 Summary of the metamorphic evolution of the NPCM and surrounding regions Evidence presented in this thesis suggests that the eastern part of the northern Prince Charles Mountains has undergone two periods of heating which were both followed by near-isobaric cooling, and which were separated by a period of decompression. Peak metamorphic conditions within the region are relatively uniform, varying from 750-830"C at around 6-7.5 kbars. A period of cooling followed peak metamorphism, and reaction textures within calc-silicates suggest that this cooling was nearly isobaric. A second heating event at lower pressure occuned synchronous with the development of upright east-west trending fabrics, with temperatures reaching 700oC at 5-6 kbars, and was also followed by near-isobaric cooling during the development of steeply dipping high strain zones. A significant implication of this evolution is that the preserved portion of the P-T history involves uplift during the development of structures consistent with having formed during convergent deformation. Before considering the implications of the metamorphic and structural evolution of the regions examined in this thesis, the metamorphic evolution of surrounding regions, and in particular those elsewhere in the NPCM, need to be considered. Most previous metamorphic interpretations of the NPCM have been based on reaction textures within calc-silicates, mainly from the Nemesis Glacier region, which have suggested near-isobaric cooling from peak conditions of 800- 830oC at 6-Tkbars (Fitzsimons and Thost, L9921, Fiøsimons and Harley, 1994). An anti-clockwise P-T path has been infened for gneisses from the Athos and Porthos Ranges by Thost and Hensen (1992), based mainly on the development of coronic garnet within calc-silicates and two-pyroxene mafîc gneisses, which is suggestive of isobaric cooling. The anti-clockwise nature of the P-T path of Thost and Hensen (1992) was inferred solely on the basis of cordierite inclusions within garnet, and therefore must be regarded as tentative. The notion of an anti-clockwise P-T path for T7I Chapter 8, Discussion the Athos and Porthos Ranges is not consistent with observations from pelitic gneisses collected in this study from Hunt Nunataks within the Athos Range, in which reaction textures are more consistent with an evolution similar to that described for Fox Ridge and Trost Rocks. In most of the Porthos, Athos and Aramis Ranges, however, pelitic units often appear essentially unreactive compared with those in the Beaver Lake region, whilst reaction textures in felsic units, such as the growth of late garnet in garnet-orthopyroxene felsic gneisses at Mt. Bechervaise, are consistent with a cooling-dominated history for the western NPCM. As discussed in Chapters 2 and 3, geochronological evidence (Tingey, L982; Manton et al., 1992; this study) suggests that the entire metamorphic evolution of the NPCM occurred during a single orogenic event between c.1000-930 Ma, with no high-grade Pan-African overprint, as suggested elsewhere in the terrain (e.g. Dirks et al., 1993). Therefore, the metamorphic evolution of the NPCM can be discussed within the context of a single tectonic event, which is assumed to be the same c.1000 Ma event described in the Rayner Complex to the west (Black et a1., 1987), on the Mawson Coast (Clarke, 1988) and along the Prydz Bay coast (Sheraton et al, 1984; Kinny et al., 1993). A consistent east-plunging lineation throughout most of this terrain (Mawson coast, clarke, 1988; Prydz Bay, Dirks and Hand, 1994; NpcM, Fitzsimons and Thost, t992, this study) supports the notion that they were all deformed in part of the same kinematic event. In addition, a number of 1000 Ma U- Pb zircon ages have been derived from gneisses in the Eastern Ghats region of India (Grew and Manton, 1986), which was adjacent to the MacRobertson and Kemp Land coasts in the Mid Proterozoic (Unrug, 1993), suggesting that this region may also have been affected by the same event. Although it would seem that the extensive Proterozoic terrain extending from Enderby Land to the Vestfold Hills may have all been involved in the same orogenic event in the Mid Proterozoíc, there are significant variations in the metamorphic t72 Chapter E, Discussion evolution across the terrain. In general, the terrain can be divided into a region that preserves evidence for near-isobaric cooling, in the NPCM and the Mawson Coast (Clarke et al., 1989; Thost and Hensen,1992; this study), whilst to the east and west, closer to the margins of the Napier and Vestfold Blocks, reaction textures implying near-isothermal decompression have been preserved (Harley, 1988; Nichols and Berry, I99l; Fitzsimons and Harley, I992a). In addition, the regions close to the Vestfold and Napier Blocks generally record peak conditions at relatively higher pressures, in the region of 9-10 kbars (Ellis, 1983; Harley, 1988). In Enderby, Kemp and MacRobertson Lands, there is no evidence for high grade metamorphism following the Mid Proterozoic, with the exception of greenschist to amphibolite facies shear zones at c.500 Ma (Tingey,1982; Black et al., 1987; Manton et al., L992). However, as discussed in Chapter 2, evidence for significant high grade metamorphism and deformation at c.500 Ma exists in the Prydz Bay region (Ren et al.,1992,Zhao etal.,1992; Dirks et al., 1993), and this introduces considerable complexity into the interpretation of the Mid Proterozoic evolution of the region. Indeed, some authors have recently proposed that the entire metamorphic evolution of the post-Archean sequence in Prydz Bay occurred during Pan-African times (Dirks, 1994). The polycyclic evolution of parts (or att) of the Prydz Bay coastline makes identification of the retrograde P-T path in the region difficult to interpret. Perhaps the most reliable indication of which textures were formed at c.1000 Ma is the association of 1000 Ma metamorphism with the development of the east-plunging elongation lineation described throughout the terrain, including the NPCM (this study) and the Mawson Coast (Clarke, 1988). Recenr studies in Prydz Bay have documented significant decompression associated with the development of the 1000 Ma lineation in Prydz Bay, including decompression of -6 kbars from peak conditions of -10 kbars in the Rauer Islands (M. Hand, 1993, pers. comm.). The notion that significant decompression occuned in the Prydz Bay region is supported by the observation of large scale (>10 km wide) high strain normal shear zones t73 Chapter 8, Discussion throughout thePrydz Bay region (M. Hand, 1993, pers. comm.). Therefore the most likely interpretation of the Prydz bay coastline on the basis of existing knowledge is that a relatively high pressure (8-10 kbars) Mid Proterozoic event was locally followed by significant decompression, and that this Proterozoic event was overprinted by lower pressure Pan African granulite metamorphism, which was particularly intense in the region of the Larsemann Hills. If this is the case, then in the Larsemann Hills, an early relatively high pressure event (M. Hand, 1993, pers. comm.) is related to Mid Proterozoic metamorphism, and the later low pressure evolution documented by StUwe and Powell (1989b) is related to subsequent Pan African metamorphism (Dirks et al., 1993). Early high pressure assemblages have been recorded from other localities nearby such as Sostrene Island (9-11 kbar,900- 950oC, Thost et al, 1991) and, before the identification of Pan African granulite metamorphism in the region, these were interpreted by some authors (e.g. Motoyoshi et al, 1991) to be Archean. 8.3 Evidence for the tectonic setting of the 1000 Ma event A number of lines of evidence exist that the Mid Proterozoic event in East Antarctica involved significant crustal thickening. In Enderby Land, the significant uplift of the Archean Napier Complex has been attributed by Sandiford (1985) and Harley (1991) to underplating by the Rayner Complex resulting in crustal thickening. In addition, the significant crustal depths during metamorphism (-30-35 km) of rocks now exposed in the Rayner Complex and Prydz Bay (Ellis, 1983; Black et al., l9g7; Harley, 1988, Thost et al., 1991) suggests that the crust was significantty thickened at the time. In MacRobertson Land, depletions in heavy rare ea¡th elements in the low- Ti Mawson Charnockite suggest that garnet was a residual phase in partial melting, requiring high pressures and a thickened crust (Young and Ellis, 1991), and Ellis (1987) proposed that such I-type HREE-depleted granitoids were a 'fingerprinr' of thickened crust during magmatism. The charnockites of the Porthos Range in the NPCM have a similar geochemistry to the Mawson Charnockite (Munksgaard et al., t74 Chapter 8, Discussion 1992), and as such probably represent a similar origin. Munksgaard et al. (1992) also proposed that the normalised trace element patterns of granulite orthogneisses in the Porthos Range are similar to those in modern day arc settings, and tentatively suggested such a setting for the orthogneisses of the NPCM. The palaeogeography of East Antarctica during the Mid Proterozoic appears to be consistent with a model in which the orogenic event in East Antarctica involved crustal thickening related to continental collision, under the assumption that tectonic processes acting in the Mid Proterozoic were broadly similar to the plate tectonic process active today. Although it is now generally accepted that Gondwana did not completely assemble until during the Cambrian (Li and Powell,1993), it is also likely that Eastern Gondwana, consisting of southwestern Australia, East Antarctica, India, Sri Lanka, Madagascar and southern and eastern Africa had assembled by or during the Mid Proterozoic (Yoshida et al, L992: Unrug, 1993). In fact ir is likely that the major Mid Proterozoic orogenies which are found throughout Eastern Gondwana (the Natal, Lurio, Eastern Ghats, Albany-Fraser and the East Antarctic (Rayner) orogen) are related to the welding together of these cratons during the period 1300-1000 Ma (Unrug, 1993), and most models for the evolution of these mobile belts have invoked continental collision (e.g. Waters, 1990; Cornell et a1., Ig92). As such it seems reasonable to suggest that the Mid Proterozoic orogen in East Antarctica may be reiated to collision between A¡chean blocks during the assembly of East Gondwana. 8.4 Thermal evolution of the NPCM The pressure-temperature-time evolution recorded in granulites from the NPCM appears to be too complex to reflect a simple p-T 'loop', although all available geochronological evidence from the region suggests that all granulite metamorphism occurred during a single major tectonothermal event between 1000 and 920 Ma (hereafter termed the 1000 Ma event). This therefore lends support to the notion that the terrain underwent at least two thermal pulses at varying t75 Chapter 8, Discussion pressures, and that these pulses were followed by near-isobaric cooling. As discussed in section 8.3, existing evidence suggests that the 1000 Ma event involved crustal thickening, and an evolution involving crustal thickening has been proposed for the 1000 Ma event in East Antarctica by a number of authors (Harley and Hensen, 1990; Nichols and Berry, I99l; Fitzsimons and Harley,l992b). If the 1000 Ma event was related to crustal thickening then it is difficult to invoke a process to maintain high temperatures in the mid-crust, even if there is substantial thinning of the mantle part of the lithosphere (Sandiford and Powell, 1991). As discussed in Chapter 7, it therefore becomes necessary to invoke advection of heat in the form of melts, which would presumably result in transient thermal pulses. In the case of MacRobertson Land, this is supported by the existence of voluminous intrusions during both D1 and D2. It should be stressed that the fact that these events are followed by isobaric cooling does not imply that they were not associated with deformation which resulted in significant changes in pressure. In fact, the association of cooling with vertical shear fabrics suggests that the cooling occurred during continuing high strain. Instead, it appears that the rate of cooling was significantly faster than any pressure changes associated with crustal deformation. The regions near the western margin of the Amery Ice Shelf (Trost Rocks, Jetty Peninsula and Fox Ridge) are significant in that they preserve evidence for decompression as well as for isobaric cooling, and thus may form a transitional region between the areas to the east in Prydz Bay, which are dominated by decompression (Harley and Hensen, 1990; Nichols and Berry, l99L; Fitzsimons and Harley, L992a), and regions to the west in the porthos, Aramis and Athos Ranges, where reaction textures are indicative of cooling (Fitzsimons and Thost, L992; Thost and Hensen, 1992; Fitzsimons and Harley, 1994). Interestingly, the regions dominated by isobaric cooling are dominated by voluminous syn-tectonic intrusions, whilst such large scale magmatism is generally absent in prydz Bay. 176 Chapter 8, Discussion However, this also means that a mechanism must be found to account for the lack of an obvious advective magmatic heat source for granulite metamorphism in Prydz Bay, and also to account for the associated decompression-dominated history. The upright nature of D2 structures in the NPCM suggests that cooling from the two thermal events in the NPCM occurred during ongoing convergent deformation. In attempting to explain the widespread observation of isobaric cooling in granulites, Bohlen (1987) proposed two possible scenarios. The first involved crustal thickening due to continent-continent collision, resulting in a clockwise P-T path with peak temperatures being attained during unloading, in which case isobaric cooling may follow decompression. The second scenario, and the one preferred by Bohlen (1987) involved thickening of the crust through underplating of magmas at the base of the crust, intrusion of magmas into the crust and/or passage of magmas through the crust, resulting in heating of the crust before and during loading, with isoba¡ic cooling occurring during the initial stages of retrogression. In the NPCM, it is proposed that neither of these scenarios is appropriate on its own, but that a combination of both of these processes occurred. on the basis of upright structures accompanying metamorphism, geochemical evidence (Young and Ellis, 1991; Munksgaard et al., rg92) and the broadly decompressional (clockwise) P-T path during the thermal peaks, the evolution of the terrain appears most likely to be attributable to a crustal thickening event. However, the apparently transient thermal pulses during this broadly decompressional history, and the abundance of voluminous syn-tectonic intrusions, suggests that magmatism played a crucial role in the thermal history of the terrain. 8.5 Constraints on timing and duration of the 1000 Ma event Geochronological data from the Proterozoic terrain of East Antarctica reveals a spread of interpreted metamorphic ages between 1030-920 Ma. Kinny et al. (1993) proposed that Proterozoic deformation and metamorphism in the Rauer Group all 177 Chapter 8, Discussion occurred within a timeframe of 30-40 Ma, between 1030 and 1000 Ma. In most other regions, however, peak granulite metamorphism is inferred to have occurred between 1000 and 960 Ma (Grew, 1978; Tingey,1982; Black et a1., 1987, Young and Black, 1991;Manton et al., 1992) with evidence in MacRobertson Land for a second thermal event at950-920 Ma (Young and Black, 1991, Manton et ú,.,1992; this study). Such a longJived evolution of up to 100 Ma is not unexpected for a major tectonic event, and reflects the timescale necessary for the growth and destruction of an orogenic belt (England and Thompson, 1984; England, 1987). However, this in no way suggests that granulite metamorphism was continuous over such a timescale, and instead it probably reflects a number of episodic, relatively transient thermal events during a prolonged tectonic evolution, probably occurring at different times in different parts of the terrain, as described in the Adirondacks and Pikwitonei Domain of North America by Mezger et at. (1990, 1991). 8.6 Geodynamic implications On the basis of the evidence presented above, it seems probable that the 1000 Ma event involved crustal thickening associated with convergent or transpressional deformation, which formed a broad orogen. Existing geochronological data suggests that the timescale of the orogenic event was at least in the order of 50-80 Ma, but the thermal perturbations existed on a significantly shorter timescale. The east- to southeast-plunging lineation widely documented for the 1000 Ma event often exhibits kinematic indicators suggesting a south-east up sense of movement (Dirks and Hand, t994; this study). As the general trend of the 1000 Ma orogen appears to be east-west, the fact that kinematic indicators suggest movement which is not entirely perpendicular to the orogenic margins may well reflect a transpressional component to 1000 Ma deformation. In general though, the sense of convergence is generally consistent with the thrusting of the Proterozoic terrain over the A¡chean Vestfold Block, resulting in the observed burial of the Vestfold Block at 178 Chapter 8, Discussion the time (Hoek et al., L992). This is in contrast to the Napier Complex in Enderby Land, which underwent significant uplift at 1000 Ma (Ellis, 1983; Sandiford and Wilson, 1984), suggesting that the margin of the Proterozoic terrain with the Napier Complex may be significantly different to the Vestfold Block margin. An important implication of the metamorphic evolution of the NPCM is that the transient thermal events occurred at progressively lower pressures, although evidence for significant decompression decreases from east to west across the region. The fact that this decompression formed during the development of upright structures consistent with convergent deformation suggests that decompression occurred during crustal thickening. Such an interpretation is inconsistent with models of homogeneous crustal thickening, in which burial accompanies thickening (e.g. England and Thompson, 1984). In the absence of evidence for extension, the question arises as to whether this decompression was driven primarily by rapid erosion or also in part to partitioning of strain during convergent deformation. It has been shown that rapid uplift of granulites occurs in regions of ongoing crustal thickening in regions such as the Himalayas, in which strain in the upper crust is largely accommodated by thrusting, but such a model is probably only applicable to n¿urow orogenic wedges, rather than broad regional terrains (Chapter 7). Therefore, it is perhaps unreasonable to account for the decompression of a regional terrain which now preserves relatively uniform peak pressures in terms of thrust-related uplift, as this would presumably lead to more obvious variations in metamorphic crustal depth, bounded by discrete high strain zones. However, if the crustal strains during convergence were partitioned between the upper and lower crust, then this could account for inhomogeneous thickening of crust. It seems likely, due to the higher pressures on the Prydz Bay coast, that the crust in the Prydz Bay region was more thickened than that in the NPCM. Such an 179 Chapter 8, Discussion interpretation is also consistent with the notion that the Prydz Bay region was the first region to undergo crustal thickening, with expansion of a plateau towa¡ds the south or southwest, ultimately incorporating the NPCM, as was also proposed by Fitzsimons and Harley (I992b). If the PrydzBay crust was thickened relative to the NPCM, then the lower crust in the region, which clearly was at high temperatures of up to 1000"C (Thost et al., 1991), would potentially have flowed laterally in attempt to reduce the horizontal buoyancy forces induced by the topography. In doing so, channel flow could have provided a mechanism for southward expansion of the plateau, as suggested for Tibet by Zhao and Morgan (1987) and Bird (1991). If this was the case, then such flow could have resulted in late convergence in the NPCM, and possible uplift of the mid to upper crust as the lower crust w¿rs thickened. Therefore, the late compressive deformation associated with uplift in the NPCM may well be related to the flow of thermally perturbed lower crust from beneath the more thickened crust of Prydz Bay. Clearly the lack of precise geochronological constraints to relate the evolution of Prydz Bay to the NPCM, and the lack of knowledge of the original geometry of the orogen, make such models necessarily speculative. rWhat does seem clear, however, is that the NPCM was affected by transient thermal perturbations during a longlived orogenic event, and shows no evidence for extensional deformation. 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Journal of Geophysical Research, 97,14207 -1422L 196 Appendix I, Field obsernations APPENDIX 1 Geological observations from other regions visited in the northern Prince Charles Mountains 41.1 Mt. Lanyon At Mt. Lanyon, approximately 100 km SSW of Jetty Peninsula, three days were spent in the field to examine the exposure on the southern cliff face. The sequence on this face consisted of amphibolites, calc-silicates and quartz-rich metasediments, all dipping steeply to the north. The predominant rock type is a banded mafic gneiss consisting primarily of hornblende, with actinolite, feldspar and possibly relict pyroxene. There are distinct mafic and felsic layers within this gneiss, and it has been intruded by early pegmatites containing hornblende which is possibly replacing clinopyroxene. This gneiss is interlayered with a massive quartz-rich metæediment containing variable biotite. Also within the sequence are abundant calc-silicates comprising calcite-hornblende- sphene + diopside * actinolite * retrograde epidote. The largest body of calc-silicates cross cuts the layering within the mafic gneiss, and appears to have been mobilised afær the development of the gneissic foliation, possibly into a shear zone. It contains numerous rafts of mafic gneiss which were often folded, boudinaged or sheared prior to the emplacement of the calc-silicate, and appe¿¡rs to be bounded by high strain zones with late pegmatites which may have intuded at the same time as the calc-silicate mobilisation. These pegmatites contain hornblende and sphene, and are a later generation than the hornblende pegmatites which parallel the foliation. The structure at Mt. Lanyon is similar to that at Else Platform, with a steeply dipping, east-west trending gneissic foliation (SÐ, with abundant evidence of an earlier isoclinally folded 51 foliation. The more mafic layers within the sequence are often boudinaged, with apparent amphibole growth between boudins of px-bearing layers. I-1 Appendix I, Field obserr)ations Within the calc-silicates and other less competent layers the folding is often chaotic. Vergence suggests that the entire cliff face is part of the overturned southern timb of an east plunging ?F3 anticline. There are at least two generations of shear zones on Mt. Lanyon - early boudinaged and folded shear zones, and more common, Nw dipping mylonitic shear zones, 5-50 cm in width, which seem to postdate all other ductile deformation. In summary, it appears likely that there is evidence for three stages of metamorphism at Mt. Lanyon - a granulite facies event early in the deformation, which was overprinted by retrograde metamorphism to amphibolite facies probably coincident with the development of 52, and then a late retrogression of hornblende to epidote. Mt. Collins ^1.2 Three days were spent in the field at Mt. collins, which is a 8 km x I km mountain approximately 25 km southwest of Mt. Lanyon. It comprises a series of igneous rocks with interesting and spectacular contact relationships. The ea¡liest rock type is a mafic country rock consisting primarily of biotite and hornblende, which has been intruded by a biotite-rich intermediate phenocrystic rock to form a banded gneiss, often with a migmatitic appearance. This was followed by the intrusion of a suite of granites and charnockites, which all appear to be part of a single series of intrusions. They range from charnockites, with opx partially altered to hornblende and biotite, to pink Kspar granites with vinually no mafic minerals. Often the charnockites have large feldspars and a yellow-brown greasy appearance similar to the McKinnon Glacier and Mawson charnockites. There is often a very gradational contact between the granites and the charnockites, although in places there are diffuse but reasonably well defined contacts between different types of granite suggesting that r-2 Appendix I, Field obseruations magma mixing may have occured. Interestingly it seemed that where there were gradational contacts between the pink granite and the charnockite, the pink granite was generally found in the vicinity of contacts with the country rock, suggesting that magma - wall rock interaction may have also been significant. In addition, the granitic suite at Mt. Collins has been cut by numerous late doleritic dykes up to 20 metres in width, which postdate the development of the dominant foliation. The final intrusion was of biotite pegmatites, with shearing along their margins. As at Mt. Lanyon, the foliation at Mt. Collins strikes at 0700, dipping steeply to the north. 41.3 Mt. Bechervaise / Hunt Nunataks, Athos Range Three days of fieldwork were done in the Mt. Bechervaise region of the Athos Range, 150 km west-northwest of Else Platform, and 50 km south of Stinear Nunataks. There were only two main rock types in this aÍea - banded pelitic and semi-pelitic gneisses of probable metasedimentary origin containing bodies of leucogneiss, and deformed bt-gt + opx granite which is also partially melæd. The gneissic fabric is folded by at least two generations of folds which are coaxial, both plunging moderately to steeply to the east. The first genration is almost isoclinal, whilst the second generation produced tight, partly recumbent large scale folds. The pelitic and semi-pelitic gneisses are generally migmatised, with abundant partial melts and locally derived leucogneisses, and contain biotite and garnet, with variable sillimanite. There are rt least two generations of garnets - one which predates the growth of biotite in 51, and one which postdates and overprints the early foliation. Within the leucogneisses there are gt + bt rich zones within which garnet defines a relict, often folded foliation, possibly after biotite, whilst in the more felsic leucogneisses the garnets do not define a foliation. There are also at least two generations of partial melts, one pre-dating the major foliation- producing event (Dr?) and one which is coincident with it. Both appe¿u to have had later growth of metamorphic garnets. r-3 Appendix I, Field obsen)ations The pelitic metasediments are similar in appearance to those on Else Platform, except with more apparent partial melting, with felsic segregations defining the gneissic foliation. On Hunt Nunataks a melanosomal layer within a pelitic gneiss was apparently quartz-free and appeared very similar to the ?corundum-bearing pelites collected at Trost Rocks in the Amery Ice Shelf. The second major lithology in the area is a gt-bt * opx porphyritic granodiorite which is also pa¡tially melted and seems to predate all significant deformation. The opx appeared to have partially altered to biotite, and the abundant garnets seem to be of metamorphic origin. The granodiorite and pelitic gneisses have been intruded initially by locally derived leucogneisses, then by a biotite aplite, followed by basaltic dykes and biotite pegmatites. The sequence on ML Bechervaise and Hunt Nunataks generally dips steeply to the south, having the same general east-west strike seen at all areas visited in the PCMs. There were two recognisable phases of deformation, beginning with the development of a gneissic foliation defined by biotite, felsic segregations and garnet (after bt?) within the felsic segregations. This was followed by the development of E-W trending, small scale, semi-recumbent tight to isoclinal folds, which fold most of the felsic segregations, and are assumed to be parasitic on the major E-W structure. Within folded leucogneisses a weak axial planar 52 foliation defined by garnet and biotite could be seen. Al.4 Stinear Nunataks The Stinea¡ Nunataks, an isolated group of peaks 50 km north of Mt. Bechervaise and approximately 80 km south of Depot Peak which has received little or no detailed study by geologists, was visited on a one day geological reconnaissance from Dovers Base. Fieldwork was restricted to the eastern cliff face of Summers Peak, which is the largest in the group, and it was found to be primarily pelitic gneisses intruded by a fine r-4 Appendix l, Field observations grained charnockite. The pelites have a gneissic foliation defined by biotite, which resulted from an isoclinal folding event which folded existing felsic partial melts and an early biotite foliation. There are rare bodies of leucogneiss, which predate the gneissic foliation and are isoclinally folded, and consist of feldspar and quartz with lesser biotite and variable garnet. These leucogneisses are probably locally derived partial melts. The main phase of partial melting seems to predate the current gneissic foliation as the felsic segregations are boudinaged and isoclinally folded, and the biotite foliation wraps a¡ound the boudins. The northern half of the eastern cliff of Summers Peak is a homogeneous medium to fine grained charnockitic opx-bt-qz-feldspar rock containing small garnets. Its contact with the pelitic gneisses was obscured by snow, but it could be seen high on the cliff and appeared to be sharp but conformable with the gneissic foliation, suggesting an early intrusive origin. In addition, a small nunatak between Summers Peak and the nearby Zebra Peak was examined and found to consist of a layered felsic bt-gt + opx gneiss of probable metasedimentary origin. 41.5 Mt. McCarthy, Porthos Range An evening trip from Dovers Base was made to the western end of Mt. Mccarthy in the Porthos Range. The rocks along the western ridge of Mt. McCarthy are predominantly mafic opx-rich gneisses with less common felsic gt-opx gneisses, and a small pod of calc-silicate. Although the general regional tend of the gneissic foliation in this region seems to be E-W, the sequence along the ridge tends at 1300 dipping 30oNE, and has an elongation lineation plunging almost directly down dip. The mafic gneisses often contain small garnets, with the assemblage identifiable in hand specimen being opx- cpx-gt-fspar, with little or no quartz. These garnets form rims around pyroxene at contacts with quartz (D. Thost, 1992, pers. comm.). The mafic gneisses often have I-5 Appendix l, Field obsen)ations isoclinally folded segregations in them which a¡e extremely pyroxene-rich, and in places it there were two pyroxenes within these segregations. The calc-silicate lens contained the assemblage scapolite-wollastonite-diopside, with the scapolite breaking down to fine symplectites of anorthite, calcite and hyalophane, and the wollastonite breaking down to calcite and quartz. The rest of Mt. McCarthy (to the east) is a garnet-rich charnockite cut by garnet leucogneiss, and with an E-W trending shear zone along its northern margin. 41.6 Crohn Massif, Porthos Range Four hours were spent at Crohn Massif examining the charnockite - counÇ rock relationships at the southeastern end of the massif. The lithologies here were monotonous opx-bearing felsic gneisses with occasional xenoliths of more mafic bt-opx gneiss, cut by orthopyroxene- and garnet-bearing pegmatites and shallowly dipping pseudo-tachylites. A large body of charnockite has a sharp intrusive contact with these gneisses. Farley Massif and Mt. Jacklyn, Athos Range ^1.7 These mountains are both within a few kilometres of Dovers Base and were visited on day trips from the base. They both are composed of a homogeneous garnet bearing granitic gneiss of probable intrusive origin, with occasional large feldspar phenocrysts. There is a strong gneissic foliation defined primarily by biotite, dipping at around 40o to the NW. The garnets often have abundant inclusions of quartz and in places a¡e rimmed by feldspar and then biotite. I-6 APPENDIX 2 Data for Sm-Nd dating of Else Platform t granite Garnet sranite -EP85 Garnet 0.0956g Spike weight Nd/Sm = 0.03002 Whole rock 0.160409 Nd/Sm = 0.03312 Ndgr = 11.3847 ppm, Sftg¡ = 9.55016 ppm I47ll44 = 0.507M Ndwn = 23.1829 ppm, SmwR = 4.52833 ppm l47lI44 = 0.11816 GarnetNd l43ll44 = 0.s14306 Std Dev. 0.0000s2 Rel. Dev. (7o) 0.0101 No. of blocks 25 Whole rocl< Nd l43fl$ = 0.s11908 Std Dev. 0.000113 Rel. Dev. (7o) 0.0220 No. of blocks 25 Age - 939 Ma Initial 143/14 ratio = 0.5111801 Ndou whole rock model age = 1.E7 Ga eNdlo¡ = -I4.2 eNd11¡ = -4.8 II-1 APPEI\IDIX 3 Representative electron microprobe analyses Mineral analyses were obtained using a JEOL i -?3:n:,croprobe at the Cent¡e for Electron Microscopy and Microstructural Analysis at the University of Adelaide, using the energy-dispersive spectra @DS) system with an accelerating voltage of 20 keV. Ba-silicates (armenite and hyalophane) were analysed using the wavelength-disperive spectra (WDS) system. Thin sections used for these analyses are held in the Department of Geology and Geophysics, University of Adelaide, with the accession number 964-. J e n)¡ P eninsula p elites 964-83, Assemblage 1 964-606E,, Assemblage 2 964-655, Assemblage 4 Trost Rocks 964-TR1, Semi-pelite 964-TR4, Corundum-bearing pelite 964-TR8, TRl1, Pelites Calc-siltcates 964-656C, 656D, MH3, MH64, 628 - Else Platform Grossular-andradite compositions, 964-MH64, MH3 Armenite compositions, MHl Hyalophane compositions, MHl, 656D, MM2 Framnes Mountains - 491, 87lLr 8713 Fox Ridge 4C, sheared garnet orthopyroxene gneiss 54, 58, 20C Pelites jotite-r¡¡1i¡np 8C. B nit': 3neiss Abbreviations at top of columns refer to the textural setting of the analysis, e.g. sym.cd - in symplectite with cordierite SZfab - within shearfabric cons.gt - consuming garnet sc.cc-rim between scapolite and calcite rim.wol - rirn adjacent to wollastonite Ml - peak a.r.remblage 964-83 central Else Platform spinel biotite cordierite Parnet ilm sym.cd adi.sil frac.gt sym.cd adi.cd sym.sp in sill sym.so adi.st larse adi.sil st.sil adi.st core I/2 wav nm.cd small )2 o.20 0.22 0.17 0.13 38.90 38.96 49.9U 49;t2 49.9U 5U.14 49.66 4E.9ó 49.1J3 49.11 39.16 39.4L 39.Ió 39.43 Tio2 o.26 1.98 3.14 0.09 o.l2 0.10 0.06 53.52 AJ2o3 62.8s 62.67 6t.& 62.47 16.63 15.89 34.25 34.5t 34.41 34.O5 34.03 33.35 33.78 33.56 22.89 22.87 22.88 23.t9 0.04 18.37 FeO 26.64 29.74 30.01 30.89 10.51 11.00 5.08 5.25 5.62 5.37 6.71 5.32 5.55 5.32 30.03 3L.29 30.18 32.08 4s.02 MnO 0. l9 0.10 0.24 0.17 0.02 0.07 0.24 0.22 0.24 0.04 0.19 0.90 0.83 0.95 0.51 0.31 Mgo 7.32 5.86 6.U 5.45 18.35 18.04 10.25 10.0s 9.74 9.80 9.10 9.84 9.49 10.2t 7.82 7.05 7.30 6.53 0.16 CaO 0.02 0.34 0.23 0.26 0.23 0.21 0.21 0.24 0.19 0.05 1.r9 1.38 1.40 t.t4 0.0r Na2O 2.27 1.87 1.44 1.15 1.69 K20 9.01 8.74 13.69 Cr2O3 0.30 0.78 1.10 0.41 all one cordierite 0.11 )tal | 99.8 101.3 100.5 100.7 95.ó 9ó.1 99.8 lu).r y)-9 99.E 100.1 97.E 9E.l 9E.4 to2.6 102.8 101.9 L02.9 v).1 Si 0.005 0.006 0.005 0.003 2.81t 2.805 4.985 4.959 4.987 5.014 4.986 4.995 4990 4.982 3.005 2.99L 2.989 2.993 3.025 Ti 0.005 0.108 0.170 0.007 0.009 0.007 0.004 0.000 1.102 0.985 Æ6c.) 2.0t2 2.0t3 2.N6 2.018 t.417 1.348 4.033 4.058 4.054 4.014 4.O28 4.01l 4.052 4.0t4 2.O39 2.046 2.059 2.075 0.001 Fe2+ 0.605 0.6'17 0.693 0.708 0.635 0.663 0.424 0.438 0.469 0.449 0.563 0.454 0.472 0.452 1.898 1.986 L.927 2.037 0.950 Mn 0.004 0.002 0.006 0.010 0.001 0.00ó 0.020 0.019 0.020 0.003 0.000 0.016 0.058 0.053 0.06r 0.033 0.007 Mg 0.296 0.238 0.248 0.222 1.976 1.936 1.526 1.495 1.451 1.461 t.362 1.497 t.+44 1.543 0.880 0.797 0.830 0.738 0.006 Ca 0.002 0.026 o.024 o.027 0.025 0.023 o.023 o.026 0.021 0.005 0.096 o.tl2 0.115 0.093 Na 0.119 0.099 0.077 0.061 0.r49 K 0.830 0.803 0.795 )tal | 3.U43 3.031 3.031 3.018 't;tEE 1.752 10.998 1r.005 10.9E6 10.979 10.991 10.992 10.979 11.011 7.976 7.986 7.981 7.969 2.066 XFe 0.6713 0;14 0.7362 0.761 0.2433 0.2549 0.2t74 0.2267 0.2444 0.2349 0.2927 0.2326 0.2469 0.22U 0.6832 0.7135 0.ó988 0.7339 964-606E: central Pladorm spinel biotite cordierite olap fetr. nm.cd core si02 0.16 0.19 0.14 0.,10 38.04 37.38 36.1 4't.89 49.26 48.75 49.20 49.18 49.10 48.89 59 38:17 -t.27 38.76 39.16 39.28 37.88 Tio2 5.08 6.42 0.15 o.t2 AJ2o3 û.82 46.71 60.07 61.46 16.01 15.73 l5.l 33.16 33.81 y.27 34.03 33.98 33.63 33.95 23.|t 23.46 23.49 2t.82 22.58 FeO 29.97 25.17 29.39 29.87 1 1.53 t2.ll 15.32 4.90 4.18 4.85 5.53 5.58 5.54 5.t't 30.85 30.80 30.60 31.87 32.30 MnO 0.45 0.1I 0.06 0.20 o.24 0.03 0.n 0.16 0.20 0.79 0.87 0.79 0.80 0.90 Mgo 6. t5 3.46 5.27 5.85 15.80 14.75 10.08 9.69 10.14 10.14 9.98 9.62 9.54 9.76 7.45 7.67 8.16 5.51 6.03 CaO 0.25 0.09 0.23 0.31 0.21 0.25 o.22 o.t7 0.14 0.25 r.30 1.18 r.25 1.34 t.2l Na2O 0.91 5.21 1.86 1.57 7.77 K20 9.27 9.32 0.84 14.39 1.l8 0.74 9ð.9 9).ó 9ð.U 99.9 9ó.1 95.E 93.1 96.2 97.9 98.3 99.2 98.8 98.1 98.0 99.01 1O2:, Si 0.004 0.007 0.004 0.01I 2.763 2.734 2.758 4.970 5.000 4.942 4.958 4.975 4.997 4.973 2.677 2.9s6 2.955 2.919 3.004 2.992 Ti 0.278 0.353 0.4t7 0.011 0.009 0.000 0.000 Æ(*ct) 2.m6 1.922 2.Ort 2.003 1.370 1.356 1.366 4.057 4.U5 4.096 4.M3 4.053 4.036 4.071 1.33t 2.078 2.088 2.079 2.039 2.054 Fe2+ 0-701 0.735 0.ó98 0.691 0.700 o.741 o.977 0.425 0.355 0.411 0.466 0.472 0.472 0.440 1.968 1.944 1.921 2.114 2.O85 Mn 0.013 0.003 0.004 0.0r8 0.021 0.002 0.009 0.013 0.018 0.000 0.01I 0.051 0.055 0.050 0.054 0.059 Mg 0.256 0.180 0.223 0.241 1.710 r.608 1.146 1.499 1.534 1.533 1.498 1.450 1.448 1.480 0.846 0.8ó3 0.913 0.652 0.693 Ca 0.020 0.007 0.019 0.034 0.023 0.027 0.0?i+ 0.018 0.015 o.027 0.299 0.106 0.095 0.101 0.114 0.r00 Na 0.049 0.352 0.t02 0.084 0.677 K 0.859 0.870 0.919 't.U)2 3.011 3.209 3.041 3.030 t.AA t.669 u.uJ2 r0.978 ll.0l0 11.009 10.990 10.985 10.991 4.9961 E.ü XFe 0.7322 0.8032 0.7s8 0.74t4 0.2m5 0.3154 04602 0.2209 0.1879 0.2115 0.237t 0.2455 0.?r';58 0.2291 0.6993 0.6927 0.6779 0.7ø4 0.7505 964-655 northeastern Else Platform biotite cordierite ßamet olas rnc.opx adl.cd px.sym adi.cd sympl. matrix DorDh adi.bt adi.st small core rim core rim.cd oDx. cd svm¡ io2 38.75 38.34 38.19 37.52 48.87 49.37 48.81 49.12 49.36 38.83 39.67 38.72 39.68 38.50 45.94 55.69 51.47 49. 51.54 50.92 Tio2 2.48 2.38 3.05 2.49 0.13 0.12 o.20 0.r4 0.01 0.19 0.16 AT203 16.07 t6.37 t5.57 16.35 33.83 34.r4 33.85 33.62 33.76 22.76 23.36 22.73 23.Q 22.89 35.21 27.39 3.33 4.89 3.75 3.06 2.58 3.34 FeO 9.29 8.78 9.63 10.51 3.99 3.98 4.51 4.24 3.83 31.23 26.17 26.80 26.32 27.93 o.75 0.31 22.49 2I.35 22.28 23.06 22.36 23.28 MnO 0.00 0.20 0.04 0.00 0.11 0.22 0.1I 0.10 1.90 t.4t 1.48 t.2t 1.30 o.25 0.14 o.36 0.30 0.76 0.47 o.20 0.49 Mgo t9.32 19.09 18.54 18. 14 10.54 10.55 10.55 to.29 10.6? 5.55 9.85 8.39 9.84 8.02 22.45 21.95 77 )? 2t.86 21.97 21.83 CaO 0.18 0.26 0.05 0.17 0.19 0.15 0.29 0.23 0.22 1.88 1.48 1.77 1.48 1.76 17.64 9.46 0.39 o.26 0.30 0.36 0.1ó 0.45 Na2O 0.00 0.00 0.00 0.00 I.UI 5.98 r.44 K20 9.51 8.96 9.12 8.69 0.10 )tal I 95.ó 94.4 94.2 93.9 et .5 98.4 9E.0 97 .7 98.1 to2.2 101.9 99.9 101.9 100.4 100.9 99.0 Si 2.796 2.790 2.799 2.765 4.973 4.978 4.9s6 4.993 4.99t 2.988 2.975 2.984 2.97s 2.966 2.tol 2.529 1.905 1.856 r.897 1.916 1.936 1.899 Ti 0.135 0.130 0.168 0- 138 0.009 0.005 0.004 0.005 0.004 AI r.367 1.404 t.345 1.421 4.058 4.058 4.052 4.029 4.O24 2.06s 2.066 2.065 2.069 2.079 1.898 1.467 0.145 o.215 o.164 0.135 0.114 0.t47 Fe2+ 0.561 0.534 0.590 0.648 0.339 0.336 0.383 0.360 0.324 2.0t0 1.642 1.727 1.650 1.799 o.029 0.012 0.696 0.665 0.692 0.720 0,703 0.726 Mn 0.012 0.002 0.010 0.019 0.009 0.009 0.t24 0.090 o.w7 0.ut7 0.085 0.010 0.005 0.011 0.009 0.024 0.015 0.006 0.0r6 Mg 2.0'17 2.0't0 2.025 1.993 1.598 1.585 1.596 1.558 r.608 0.637 1.101 0.964 1.100 0.921 t.238 1.2t9 t.231 1.2t6 1.230 1.213 Ca 0.014 0.020 0.004 0.013 0.020 0.0r7 0.031 0.025 o.o24 0.155 0.1 t9 0.146 0.119 0.145 0.864 0.461 0.015 0.010 o.0t2 0.014 0.00ó 0.018 Na 0.095 0.527 0.104 K 0.875 0.832 0.853 0.817 0.005 )tal I 7.824 7.794 7.787 7.795 ru.99E 10.993 ll.0lE 10.974 10.9E8 7.979 7.992 7.983 7.990 7.995 4.997 5.001 XFe 0.2t2s 0.205t 0.22s7 0.2454 0.1752 0.1748 0.1934 0.1878 0.1675 0.7595 0.5985 0.6419 0.6001 0.6615 0.3598 0.353 0.36 0.3719 0.3635 0.3744 964-TR1: Trost Rocks biotite plasiocl.ase gamet in st st rim b€tw.st betw.st st.matr. retr.sD plas Dl.st.bt trost I trost I core rim small la matrix sio2 35. l5 34.94 35.04 35.42 35.17 36.51 55.32 54.76 37.66 38.14 38.Is 37.42 37.tO 37.08 0.t7 o.l2 0.04 T1o2 6.33 s.92 4.57 4.54 4.75 2.99 0.20 0.06 0.04 0.04 A12o3 14.94 15.23 15.61 15.81 15.46 16.04 26.56 27.92 22.1t 22.30 22.41 22.25 21.81 2r.7r 57.08 57.33 s9.M 58.69 58.48 59.26 FeO t3.o7 12.92 13.53 13.59 13.63 9.78 0.14 1.02 MnO 0.08 0.07 0.07 0.16 0.16 29.96 30.10 29.96 32.20 32.s8 32.s4 34.29 34.60 34.87 32.75 34.65 32.49 Mgo 14.40 13.45 13.79 13.81 t3.28 18.15 o.t2 0.71 0.98 0.85 0.83 0.74 0.96 0.07 0.10 0.19 0.07 CaO 0.17 0.36 0.28 0.16 o.24 0.,10 8.74 9.40 6.73 6.38 6.55 5.06 4.72 4.19 5.62 5.59 5.55 6.76 5.37 6.80 Na2O 0.12 0.08 0.18 0.09 0.ll 6.36 5.66 1.43 1.68 1.76 1.36 t.29 t.52 K20 9.41 9.r9 9.25 9.47 9.16 9.06 0.00 0.39 0.39 o.57 0.18 0.00 Cr2O3 0.24 0.27 0.42 0.56 0.38 otal I 93.6 92.1 923 93.0 9I.E 93.0 91.5 99.0 9E.ó 99.6 y).7 y).t 9E.2 9E.0 Si 2.710 2;t2t 2.73r 2.739 2.749 2.832 2.486 2.428 2.980 2.991 2.985 2.977 2.987 2.997 0.m0 0.005 0.003 0.001 0.000 0.000 Ti 0.367 0.347 0.268 0.264 0.279 0.1't4 0.007 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.001 0.001 AI 1.358 1.399 t.434 t.Mt t.424 1.467 Lq1 1.459 2.063 2.062 2.M7 2.087 2.070 2.069 1.950 t.937 t.952 1.950 1.961 1.966 Fe2+ 0.7s8 o;157 0.793 0.791 0.802 0.571 0.005 0.034 0.000 0.000 0.000 0.000 0.000 0.000 Mn 0.000 0.m0 0.005 0.005 0.005 0.000 0.006 0.006 1.983 t.974 1.961 2.143 2.194 L2m 0.831 0.829 0.818 o.772 o.824 o.765 Mg 0.940 0.887 0.910 0.905 0.879 1.193 0.00s 0.047 0.065 0.056 0.05ó 0.050 0.066 0.002 0.002 0.005 0.002 Ca 0.020 0.042 0.033 0.018 0.028 0.046 0.585 0.621 0.794 0.745 0.764 0.600 0.566 0.505 0.243 0.239 o.232 o.284 0.228 0.285 Na 0.010 0.007 0.015 0.007 0.000 0.009 0.306 o.269 0.1213 0.t4t2 0.1476 0.1161 0.1113 0.1317 K 1.407 1.388 1.398 1.420 1.388 1.363 0.022 0.021 0.031 0.010 't '/.ó5ó Otal | 7.569 1 .548 7 .5ðl .59 t 7.553 4.8U2 4.822 7.988 7 .91ö 7.9ö l 1 .919 7 .gtE 7.96E X-Fe 0.M& 0.4605 0.4658 0.4665 0.477 0.3238 0.6565 0.6979 0.0412 0.0483 0.0504 0.0398 0.0381 0.0454 0.774 0.7765 0.779 0.73tt 0.7836 0.7284 964-TR4: Trost Rocks biotite feldspar iln garnet spinel iìm.cd in.cdksp con.ilm cons.sp adi.gt cons.byst plag kso kso adi.sym adi.cd core adl sym with.cd ea¡ly sym.cd with-bi si02 36.52 37.15 36.12 37.27 35.78 35.95 47.52 63. l5 63.78 0.r2 39.M 39.72 39.82 38.50 0.26 o.23 0.33 0.36 49.79 49.64 49.70 Tio2 7.3r 3.14 6.91 4.92 6.43 6.86 0.11 0.31 o.26 52.7t 0.12 0.10 AJ2o.3 14.99 16.20 15.71 15.31 15.67 15.97 33.17 18.37 18.51 0.06 23.06 23.39 23.32 22.96 62.65 62.39 63.59 6t.14 34.4L 34.15 34.31 FeO 13.54 t2.64 r4.98 13.90 13.27 t4.@ 0.45 0.02 44.35 31.85 29.92 30.16 31.63 30.17 30.34 29.02 29.23 MnO 0.04 0.15 0.22 o.25 0.06 0.14 o.2t 0.30 0.87 0.09 0.58 0.89 0.13 0.05 0.06 5.40 5.24 5.61 Mgo 12.95 t5.52 t2.90 13.90 14.15 12.73 1.13 6.4 7.51 8.52 7.53 7.64 6.23 8.34 7.24 0.01 0.14 CaO o.27 0.26 0.07 0.09 0.22 0.16 t6.33 0.44 0.43 1.51 1.50 1.17 t.s4 10.08 10.37 10.05 N¿O 2.03 0.66 0.96 0.ll 0.05 0.16 0.28 0.26 K20 9.52 9.57 9.36 9.85 9.5s 9.7t 14.82 14.63 0.03 o 0.63 0.71 0.78 0.42 95.1 94.6 96.1 95.5 95.3 95.5 99.7 91.9 9E.E 9E.8 102.8 tO2.2 L0/'.2 103.0 101.5 99.9 102.1 98.6 Si 2.743 2.821 2.695 2.802 2.70t 2.694 2.1 88 2.973 2.974 0.007 2.975 2.942 2.961 2.930 0.007 0.006 0.009 0.010 4.975 4.966 4.970 Ti 0.413 0.179 0.388 0.278 0.365 0.386 0.004 0.011 0.009 1.001 0.002 AI 1.328 1.450 1.382 1.356 1.395 l.4ll 1.800 1.o20 1.017 0.002 2.071 2.043 2.033 2.0ffi 1.996 2.022 2.W 1.998 4.053 4.028 4.444 Fe2+ o.765 0.722 0.841 0.786 0.754 o.794 0.017 0.001 o.937 2.028 2.074 2.028 2.013 o.682 0.697 0.648 0.678 Mn 0.002 0.010 0.014 0.015 0.002 0.006 0.008 0.006 0.056 0.055 0.038 0.057 0.003 0.001 0.001 0.451 0.439 0.469 Mg 0.824 0.998 0.815 0.885 0.905 0.808 0.043 0.73L 0.79 0.E43 0.854 0.308 0.255 0.332 0.299 0.0t2 Ca 0.030 0.o29 0.007 0.010 0.02s 0.018 0.806 0.022 0.021 0.125 0.t237 0.t2M 0.1253 1.501 1.546 1.498 Na 0.181 0.060 0.086 0.005 0.003 0.017 0.030 0.027 K 1.386 1.409 1.355 t.435 1.399 t.4rt 0.890 0.870 0.00r 7.4et 7.ólE 1.483 7.566 7.559 7.522 4.998 4.9ðl 4.981 2.U)2 7.98s 8.036 8.023 8.040 2.995 2.984 2.991 2.989 X-Fe 0.4815 0.4199 0.508 0.4705 0.4545 0.4958 0.7352 0.7219 0.706/. 0.7022 0.6891 0.7321 0.6615 0.6939 o.23rl O.22tl 0.2385 Sam 9ó4-TR8: Trost Rocks biotite cordierite feldspar gamet M2 adi.cd Mz adi.st cons.bi adi.sp 2/3sD.st adi.sym adi.st adr.st matnx l(sD latß, rim.ad c(re rim lsill.cor svm.cd edee.stearl si02 37.08 36.26 35.95 50.31 49.76 46.07 50.72 s0.64 49.74 54.35 55.91 63.43 39.10 38.27 37.82 37.97 0.08 0.16 o.25 0.31 0.2r 0.38 Tio2 4.76 7.12 6.40 0.14 0.01 41203 t6.43 15.59 t6.42 34.27 34.71 37.95 34.78 34.67 34.06 28.45 28.42 18.75 23.00 22.60 21.96 1'' ?) 62.22 61.45 62.86 61.16 6t.44 63.6t FeO t4.w 15.82 t4.52 5.71 5.67 6.99 5.2t 5.47 5.28 0.49 0.15 32.21 31.23 28.76 33. l6 MnO 0.0ó 0.l l 0.31 0.06 o.27 o.20 0.14 o.79 0.97 0.67 1.04 29.77 29.4t 31.35 31.66 3t.45 3r.69 Mgo t4.23 12.22 t2.61 9.93 10.0ó 9.88 10.20 t0.23 10.04 6.35 6.53 7.75 5.64 0.08 0.15 0.15 0.16 0.22 CaO 0.10 0.06 0.24 0.27 0.30 0.16 0.15 0.13 0.21 10.58 10.2s 0.68 t.N r.46 1.49 l.4t 7.20 7.59 7.25 6.21 6.84 7.4L Na2O 5.s7 5.57 L.67 2.M K20 9.83 9.50 9.74 0.07 13.63 0.16 1.95 1.25 1.96 1.59 1.13 1.08 0.37 9ó.ó 96.7 96.2 100.5 100.7 10r.0 l0l.l l0l.l 99.3 99.8 100.5 98.3 102.r lol.l 100.5 101.51 100.6 102.7 103.4 lr si 2.755 2.688 2.677 5.001 4.943 4.608 4.999 4.994 4.993 2.464 2.502 2.964 2.977 2.968 2.942 2.958 0.002 0.004 0.007 0.008 0.006 0.010 Ti 0.266 0.397 0.358 0.010 AI r.439 1.363 t.442 4.016 4.065 4.474 4.040 4.030 4.031 1.521 1.500 1.033 2.065 2.066 2.013 2.050 2.011 1.981 1.992 1.994 1.988 t.987 Fe2+ 0.788 0.883 0.814 0.475 0.471 0.585 0.429 0.45r 0.443 0.0r8 0.006 2.051 2.025 1.870 2.r60 0.683 0.673 0.705 o.732 0.722 0.702 Mn 0.003 0.007 0.019 0.005 0.0r0 0.007 0.006 0.051 0.0úf o.u4 0.069 0.002 0.003 0.003 0.004 0.00s Mg 0.896 o.767 0.795 1.471 1.489 1.472 1.498 1.503 t.502 0.720 0.755 0.897 0.654 0.294 0.309 0.290 o.256 0.280 o.293 Ca 0.011 0.007 o.027 o.029 0.031 0.017 0.016 0.014 0.o22 0.514 o.492 0.034 0.114 0.121 o.t24 0.118 Na 0.490 0.483 0.151 0.307 0.008 K 1.415 r.365 t.Q6 0.004 0.812 0.0ó8 1.513 7.4-t6 7.539 10.991 lr.0l4 ll.l55 10.981 10.991 10.991 5.U22 4.991J 5.000 7.e1E 7.y)E E.197 8.008t 2.992 3.04'. X-Fe 0.468 0.535 0.50s8 0.244 0.2N3 0.2842 0.2227 0.2309 0.2279 0.7402 0.7284 0.6758 0.7675 0.6988 0.685 0.7082 0.7Æ 0.7207 0.70s8 TRll: Trost biotite cordierite kspar RArnet withsil cons.sil rim.sil cons-sil cons.sil cons.st svm.sD svrn.sD ksD rim.svm rim.cd coreMl late', core.sm rim.sm st. sio2 35.13 3s.39 36.30 35;t8 35.87 48;18 49.32 48.94 48.08 63.01 37.76 38.45 38.62 38.40 38.n 38.39 0.06 0.28 0.15 T1o2 5.01 7.30 8.39 6.94 0.19 0.06 0.16 o.46 0.21 0.45 AJ2o3 15.51 15.61 15.73 15.51 l6.l 33.4t 33.41 33.84 32.91 18.26 22.14 22.5t 22.55 22.32 22.38 22.20 60.67 60.,1O 61.62 6t.69 FeO 13.80 15.29 15.64 14.55 14. 5.62 5.9 5.79 5.21 0.09 31.84 32.06 30.42 32.84 31.77 31.82 32.t8 32.39 31.t6 31.68 MnO 0.15 0.19 0.0ó 0.10 0.18 1.07 0.59 o.67 0.7t 0.88 0.84 0.04 0.15 0.02 0.06 Mgo 13.66 It.46 t2.10 12.99 10.11 10.00 10.22 9.96 6.00 7.L2 7.88 6.10 6.82 6.51 6.9s 6.32 6.84 7.38 CaO 0.21 o.l'l 0.10 o.25 0.31 0.19 0.19 0.29 0.28 0.43 1.68 1.45 1.63 L.7t t.43 1.50 Na2O 0.97 o.tz 0.2t o.32 K20 9.84 9.85 9.6s 9.95 15.10 Cr2O3 0.59 1.09 0.4s 1.05 9ö.5 98.7 !).2 9õ.J 98.5 100.5 to2.2 101.E loz.t l0l.ó 101.3 Si 2.715 2.664 2.654 2.679 2. 4.959 4.997 4.939 4.975 2.96r 2.963 2.954 2.959 2.968 2960 2.979 0.002 0.008 0.004 Ti o.291 0.413 0.461 0.391 0.3 0.015 0.005 0.012 0.016 0.004 0.009 N GG) t.4r3 r.385 1.356 1.370 4.004 3.990 4.027 4.015 l.0ll 2.M8 2.039 2.037 2.034 2.041 2.030 1.977 1.977 1.980 1.980 Fe2+ 0.802 0.866 0.861 0.820 0.478 0.478 0.489 0.451 0.004 2.089 2.060 1.949 2.123 2.055 2.065 0.744 0.752 0.710 0.721 Mn 0.016 0.005 0.009 0.007 0.071 0.039 0.044 0.046 0.057 0.055 0.001 0.004 0.001 Mg 0.894 0.731 o.749 0.824 0.860 1.531 1.509 t.537 1.536 0.702 0.815 0.899 0.703 0.786 0.753 0.286 0.26t 0.278 0.299 Ca 0.024 0.019 0.011 0.028 0.034 o.021 0.020 0.031 0.031 0.o22 0.141 0.1195 0.134 0.1415 0.1187 0.1248 Na 0.088 0.006 0.01I 0.017 K t.474 1.438 1.369 1.445 1.459 0.905 t1.024 1I.004 ll.u35 Il.0l7 5.UI4 8.013 8.026 8.022 8.015 8.019 8.006 X-Fe 0.473 0.5423 0.5346 0.4987 O.49ll 0.2377 0.24M 0.24t3 0.2268 0.0469 0.0394 0.0443 0.047 0.0393 0.0416 0.722 0.742t 0.7188 0.7067 Else Platform scapolite diopside anorthite rim.sym adi.an adì.wol wol rim cc rim oroqradewol rim prog. an an an an sio2 42.72 41.73 41.57 42.01 53.26 52.90 52.85 52.9L 51.77 44.52 43.86 43.73 44.28 50.08 50.20 Tio2 0.37 0.22 o.l4 o.l2 o.2l 0.04 0.06 0.05 0.32 0.14 41203 29.89 29.47 29.69 29.78 0.91 0.85 1.57 0.65 1.42 36.43 36.14 35.68 36.41 0.48 0.47 Fe2O3 FeO 0.10 0.10 0.01 6.12 6.25 5.50 6.55 4.98 0.07 0.16 0.11 0.16 0.33 MnO 0.19 0.43 o.23 o.26 0.42 0.58 o.43 0.ó4 0.43 0.14 0.03 o.t2 0.19 0.63 0.31 Mgo 15.00 14.38 14.49 14.45 t4.78 CaO 20.72 20.87 20.84 20.60 24.77 25.1 8 24.81 25.08 25.M t9.57 19.80 19.38 20.o2 47.07 47.56 Na2O 1.38 1.48 2.42 1.63 0.19 0.35 0.33 0.24 K20 0.11 0.09 0.10 0.02 0.08 0.05 95.5 94.3 95.1 94.4 100.5 100.1 99.9 100.3 98.4 101.0 100.3 99.4 tot.2 Si 6.564 6.515 6.460 6.533 1.966 1.966 1.958 t.967 1.948 2.O37 2.024 2.036 2.026 0.984 0.985 Ti 0.043 0.025 0.016 0.014 0.006 0.001 0.002 0.002 0.005 0.002 AI 5.415 5.425 5.440 5.4ó0 0.040 0.037 0.068 0.o29 0.063 1.965 t.967 1.958 1.964 0.011 0.011 Fe2+ 0.012 0.013 0.001 0.1 89 0.t94 0.170 0.203 0.157 0.003 0.006 0.004 0.000 0.003 0.005 Mn 0.024 0.056 0.030 0.035 0.013 0.018 0.014 0.020 0.014 0.005 0.001 0.005 0.008 0.010 0.005 Mg 0.826 0.797 0.800 0.800 0.829 Ca 3.411 3.492 3.470 3.433 0.980 1.003 0.985 0.999 1.009 0.959 0.979 o.967 0.981 0.991 0.999 Na 0.410 o.M9 0.729 0.491 0.017 o.032 0.030 0.021 K o.022 0.019 0.020 0.0(X 0.002 0.001 15.901 15.981 16.t79 15.970 4.0t4 4.015 4.N2 4.019 4.020 4.987 5.008 5.000 5.001 80.506 80.837 81.328 81.995 0.8137 0.804 0.8243 0.7973 0.8411 0.9826 0.9687 0.9698 0.9787 (EaAn) (XFe) (Xan) s 964-656D: northeastern Else Platform scapolite diopside anorthite wollasønite adi.svm adt wol nm wol nm woi insvm adisc insvm sio2 4r.52 4l.u 41.62 4t.99 41.15 41.81 51.06 51.60 51.69 43.39 43.06 44.24 49.3r Tio2 0.17 0.21 0.01 0.04 0.14 0.03 0.08 0.19 AJ2o3 29.75 29.47 29.96 30.18 28.81 28.85 0.86 0.91 0.76 35.95 3s.69 36.13 0.52 Fe203 FeO 0.14 0.25 0.o2 0.05 0.14 0.06 8.8s 10.98 8.93 0.04 0.18 MnO 0. l8 0.08 0.14 0.17 0.72 0.7r 0.56 0.28 0.0? o.49 Mgo t2.39 11.03 12.22 0.04 CaO 20.88 21.36 21.32 2t.14 20.77 20.23 24.41 24.35 24.n 19.58 19.37 19.74 47.30 Na2O 1.61 1.30 1.25 r.28 1.58 1.90 0.28 0.17 0.01 0.48 K20 0.18 0.03 0.05 0.05 o.2L 0.21 94.4 94.3 94.4 94.9 92.1 93.1 9E.4 v).6 99.1 v).5 9E.3 100.3 Si 6.478 6.500 6.486 6.s01 6.539 6.601 1.958 1.970 1.968 2.020 2.O24 2.O37 o.976 Ti 0.020 0.o25 0.001 0.005 0.004 0.00r 0.003 0.003 AI 5.472 5.424 5.503 5.509 5.396 5.370 0.039 0.041 0.034 1.973 1.978 1.962 0.012 Fe2+ 0.018 0.033 0.002 0.007 0.019 0.m8 0.284 0.351 0.284 0.001 0.003 Mn o.024 0.0r l 0.018 0.023 0.024 0.023 0.018 0.01l 0.003 0.008 Mg 0.708 0.628 o.694 0.001 Ca 3.491 3.572 3.560 3.507 3.536 3.423 1.003 0.997 1.016 0.977 0.976 0.974 1.002 Na 0.487 0.393 o.376 0.385 0.487 0.s83 0.025 0.015 0.001 0.018 K 0.036 0.005 0.009 0.010 0.M2 0.M2 16.027 t5.963 15.955 15.942 t6.O22 16.026 4.019 4.ü)9 4.015 5.007 4.994 4.980 82.413 80.803 83.42s 83.633 79.871 78.989 0.7t39 0.64t6 0.7W4 0.9747 0.9847 0.9991 0.9824 (EqAn) üdi) (Xan) Else Platform Else Platform scapolite dion 0n woll dio aÍ7 core core core sio2 43.6t 43.98 43.96 50.54 44.57 49.85 43.59 43.38 43.83 42.78 42.79 42.65 50.94 44.46 49.31 T1o2 0.06 0.15 0.19 0. o.l2 0.19 41203 28.72 28.79 28.74 o.97 36.57 0.63 27.68 27.9 28.67 27.84 28.W 27.56 r.00 36.39 0.52 Fe2O3 FeO 0.08 0.10 0.21 13.75 0.34 0.28 0.045 0.14 0.12 0.2t 0.13 0.13 0. 16.15 0.18 MnO 0. 19 0.18 0.17 o.42 0.1ó 0.M o.263 o.37 o.l2 0.14 0.32 0.07 0. 0.55 o.22 o.49 Mgo 9.00 0.ll 0.09 o.2L 8.16 0.04 CaO 19. l0 t9.46 t9.4'l 24.Ot 19.56 46.s7 L9.567 19.s8 l9-48 19.52 19.90 19.22 23.8L 19.88 47.30 Na2O 2.O5 2.t7 2.M o.47 0.ll 1.895 2.Ut 1.93 1.77 1.90 2.28 2. 0.48 K20 0.44 0.48 0.46 0.05 0.467 0.36 0.42 o.47 0.5E 0.45 0.08 )lal I 94.2 95.2 95.0 98.7 101.7 94.0 94. 929 93.7 92.5 93.6 100.ó 101.1 Si 6.774 6.772 6.776 1.972 2.O30 0.987 6.829 6.7',t6 6.768 6.759 6.720 6.767 6.791 1.960 2.O33 o.976 Ti 0.007 0.018 o.o22 0. 0.004 0.003 AI 5.258 5.226 5.223 0.045 1.964 0.015 5.1 13 5.154 5.221 5.185 5.202 5.155 5. 0.046 t.962 0.012 Fe3+ Fe2+ 0.010 0.0r3 0.02't 0.448 0.013 0.005 0.006 0.019 0.016 0.028 0.017 0.017 0.s25 0.003 Mn 0.025 0.023 0.023 0.014 0.006 0.008 0.035 0.049 0.016 0.018 0.043 0.009 0.018 0.008 0.008 Mg 0.523 0.004 0.021 0.048 0.473 0.001 Ca 3.178 3.21O 3.215 1.004 o.954 0.984 3.284 3.278 3.223 3.3M 3.349 3.267 3. 0.991 o.974 t.o02 Na 0.616 o.647 0.609 o.042 0.030 0.576 0.626 0.579 0.543 0.578 0.700 0. 0.018 K 0.087 0.094 0.091 0.000 0.093 0.07r 0.083 0.094 0.117 0.w2 0.009 15.949 15.986 15.963 4.006 5.ü)9 4.013 4.982 75.269 74.215 74.086 0.5384 0.9583 70.42t 71.812 74.018 72.843 73.445 71.829 73. 0.4737 I (EqAn) (xdÐ (Xan) (ESAn) (xdi) (Xan) (grossular compositions are shown on a separate table) 9il-C8z northeast Else Platform øørthite dbpside solenc clinozoisite t¡ an an cenEe nm sio2 4.57 M.6t 44.6 50.74 51.04 30.18 38.59 39.33 Ti()2 020 0.t2 0.13 0.20 33.73 0.03 o.24 A12C)3 36-57 36A2 36.39 1.00 t.26 4.65 29-99 3t.19 Fe2O3 FeO 0.34 0.17 t6.67 15.03 o.u 4.32 2.80 Mrro 0.1ó 0.ll 0.n 1.89 1.31 0.10 o-72 053 Mso 726 r0.B 030 CaO 19.56 1938 19.88 23.6t 2t-22 n.83 23.68 23.93 Na2O o.47 o.4 o.0l K2O lot./ tolJ lut.t lol.3 l00J 97.U si 2.OtO 2.ß5 2-033 1.963 1.960 1.00? 3.139 3.LÙ| Ti 0.(II7 0.m4 0.004 0.00ó 0.855 0.0æ 0.014 AI t96/. 1.959 1.962 0.04ó 0.057 0.Iil 2.875 2.977 Fe2+ 0.013 0.m7 0539 0.483 0.00ó o.293 O.IEó Mn 0.mó 0.m4 0.008 0.062 0.043 0.002 0.05{) 0.036 Mg 0.418 0586 0.014 Ca 0.954 0.947 o.974 0.979 0.t73 0.995 2.M3 2.037 Na 0.u2 0.ß9 0.001 K 5.Ut1' 4.v)A 4-982 4.010 4.urú J.()õ{J 0.95E3 0.9ó07 I 0.43ó9 0.5482 (Xan) rxdil Grossular of with recalculations for andradite and wol inc.sc rim.sc rim sc.cc wol.sc sc.wol sc.wol .sc .sc .sc Tioz 0.41 0. l8 0.13 0. l0 0.29 0. l9 0.16 0.35 0.47 0.24 0.13 0.08 0.05 41203 16.75 16.93 18.04 16.43 16.36 15.76 15.49 16.s6 14.50 16.26 22.40 22.53 22.91 Fe2O3 FeO 9.68 10.15 8.31 9.42 8.83 9.77 11.25 9.97 tt.44 9.78 0.46 0.37 0.l0 MnO 0.39 0.ó8 0.58 0.49 0.54 0.41 0.58 0.73 0.48 0.57 0.35 0.41 0.20 Mgo 0.20 0.06 0.18 0.21 CaO 34.00 33.ó0 34.29 34.07 33.13 34.34 34.07 34.32 33.09 34.06 36.84 37.64 37.41 Na2O o.t2 0.38 1.08 0.25 0.01 0.19 0.23 K20 0.09 si 3.038 3.065 3.025 3.O37 3.036 2.993 3.O28 3.003 3.028 3.009 2.935 2.938 2948 Ti 0.o24 0.011 0.008 0.006 0.018 0.01I 0.010 0.021 o.029 0.014 0.008 0.004 0.003 AI 1.556 t.557 1.659 1.551 1.565 r.492 1.464 r.549 1.406 t.537 2.020 2.004 2.037 Fe3+ Fe2+ 0.638 o.662 0.542 0.631 0.599 0.656 0.754 0.662 0.787 0.656 0.030 0.023 0.006 Mn 0.026 0.045 0.038 0.033 0.037 0.028 0.039 0.049 0.033 0.039 0.023 0.026 0.013 Mg 0.o24 0.007 0.022 0.025 Ca 2.870 2.807 2.866 2.923 2.880 2.956 2.927 2.9t8 2.915 2.927 3.019 3.044 3.024 Na 0.018 0.060 0.1ó8 0.040 0.002 o.o27 0.034 0.000 K 0.009 Si 2.9744 3.0099 2.9739 2.9636 2.8722 2.9438 2.9289 2.9329 2.9328 Ti 0.0236 0.0104 0.0075 0.0172 0.0107 0.0093 0.0205 0.0284 0.0138 AI r.5235 t.5287 1.6305 1.51 1.5278 t.432 t.4232 t.5tt2 r.3614 1.4982 Fe3+ Fe2+ 0.6248 0.6503 0.5329 0.61 0.5848 0.6298 0.7333 0.6455 0.7619 0.ó39 Mn 0.0252 0.@'38 0.0378 0.0365 0.0266 0.038 0.0478 0.0322 0.0375 Mg 0.0229 0.0072 0.021 0.0242 Ca 2.8105 2.7568 2.8174 2 2.8119 2.8359 2.8452 2.8462 2.8237 2.8525 Na 0.0181 0.0582 0.1611 0.0386 0.002 K 0.0088 Fe2+ o.t26t 0.2197 0.1263 0.0851 0.016 -0.342 0.0626 0.0555 0.0072 0.0285 Fe2O3 8.5881 7.4723 7.M66 9.5424 t6.76 1t.433 10.129 12.593 10.381 FeO 1.955 3.43 L.9692 t.2993 0.2423 -5.31 0.961 0.8568 0.1086 0.4357 Xalm o.M26 0.0727 0.M24 0.0056 -0.135 0.0212 0.0188 0.0025 0.0097 Xsps 0.0085 0.0145 0.ot27 0.011 0.0127 0.0105 0.0129 0.0162 0.01l2 0.0128 Xprp 0.009 0.0024 0.0073 0.0082 Xgrs 0.7149 0.7t21 0.7563 0.7111 0.7154 0.6642 0.6549 0.694 0.6299 0.6887 Xadr 0.234 0.2006 0.1886 0.2663 0.4509 0.3086 0.271 0.3492 0.2807 Armenite comoositions. Sample MHl. Else Platform sio2 51.53 53.99 51.86 54.77 47.22 53.32 50.96 53 59 5t.97 s0.50 49.18 53.70 49.37 A1203 25.73 2s.04 26.20 25.21 27.82 26.09 27.36 25 70 26.52 27.92 27.86 24;t6 27.4s CaO 8.80 8.37 8.98 8.43 9.60 8.96 9.45 8.9s 9.11 9.77 9.76 8.47 9.81 Na2O 0.60 0.93 0.39 0.80 0.19 0.57 0.25 0.ó0 0.5ó 0.14 0.20 0.64 0. r0 KzQ 0.79 r.49 0.77 1.50 0.13 1.05 0.45 l.18 0.80 0.13 0.14 1.51 0.09 BaO 10.53 8.23 10.50 8.13 12.08 t0.74 1 1.16 9.03 10.57 12.65 12.34 8.51 t2.37 s04 0.08 Si 9.474 9.738 9.44r 9.776 8.913 9.s28 9.232 9.622 9.389 9.r25 8.992 9.752 9.160 AI 5.s73 5.323 5.620 5302 ó.188 5.494 5.841 5.437 5.647 5.945 6.002 5.299 5.952 Ca t.734 1.6t7 1.750 t.612 1.940 1.715 1.834 1.722 1.763 1.890 1.910 1.647 1.93r Na 0.2r2 0.326 0.13ó o.275 0.069 0.t97 0.087 0.2û9 0.194 0.050 0.070 0.225 0.037 K 0.1 8ó 0.343 0.178 0.342 0.032 0.239 0.105 0.271 0.184 0.029 0.032 0.351 0.021 Ba 0.7s8 0.581 0.749 0.5ó8 0.892 0.751 0.792 0.635 0.747 0.895 0.883 0.605 0.891 S 0.01I SiAI 15.047 15.061 15.061 15.078 15.101 15.022 t5.073 15.059 15.036 L5.07 14.994 15.051 15.l12 CNKB 2.890 2.867 2.8t3 2.797 2.933 2.902 2.818 2.837 2.888 2.864 2.895 2.828 2.880 exSi 0.475 0.74t 0.443 0.780 -0.088 0.s29 0.233 0.624 0.390 0.126 -0.008 0.755 0.161 Na+K 0.398 0.669 0.314 o.6Lt 0.101 0.436 0.t92 0.480 0.378 0.079 0.102 0.576 0.058 Else Platform and Mt MH1 656n It svmDl hval Ba-rich Ba-Door Ba-rich Ba-rich large larse symDl rim.wol rim sâme large rim symDl sio2 64.49 62.t| 52.22 6r.33 58.ó4 52.45 54.54 54.6t 55.69 55.00 48.4t 58.97 54.96 54.59 55.44 52.63 48.45 52.91 46.22 51.03 A1203 20;1',7 20.41 24.38 19.93 21.60 23.26 21.77 21.85 21.34 21.76 22.70 20.79 22.06 21.98 21.76 22.93 21.82 22.26 28.10 23.34 CaO 0.29 o.22 o.67 0.45 2.W 1.30 0.o2 0.o2 o.44 0.41 2.45 0.60 0.05 o.t2 0.2s o.54 3.36 1.10 8.28 0.64 Na2O 0.31 0.33 r.34 o.49 o.43 2.88 0.49 0.48 0.46 0.45 0.31 0.22 0.43 0.50 o.43 0.,1O 0.39 0.53 o.26 0.29 K20 7.28 13.68 s.88 12.81 11.30 3.68 10.20 10.57 11.06 10.36 6.40 12.47 10.35 10.08 10.64 8.41 7.68 8.86 4.36 8.26 BaO 6.99 7.77 16.23 6.83 8.99 t6.41 14.59 14.t5 12.54 13.77 22.19 9.76 14.70 14.72 12.98 t8.29 17.25 16.10 13.78 19.38 s04 0.02 0.07 rtal 1100.1 104.5 100.7 101.9 103.0 100.0 101.7 101.7 lol.5 101.7 102.5 102.E roz.5 102.0 l0r.5l ll Si 2.968 2.877 2.6t6 2.89r 2.040 2.639 2.719 2.719 2.748 2.726 2.536 2.8t7 2.717 2.713 2.736 2.644 2.563 2.660 2.320 2.598 AI t.t26 l.lr4 1.439 1.107 1.943 r.379 r.279 1.282 L2N t.271 1.401 t.t7l 1.285 1.287 t.265 1.357 1.360 1.319 1.662 1.400 Ca 0.014 0.011 0.036 o.023 0.977 0.070 0.001 0.001 0.023 o.022 0.138 0.031 0.003 0.006 0.013 0.o29 0.190 0.059 o.445 0.035 Na 0.027 o.029 0.130 0.045 0.035 0.281 0.048 0.M7 o.M4 0.043 0.031 0.020 0.041 0.048 0.041 0.039 0.040 0.052 0.02s 0.029 K o.42't 0.808 0.375 0.770 0.004 o.236 0.648 0.67r 0.696 0.655 0.427 0.760 o.652 0.639 0.ó69 0.539 0.518 0.568 o.279 0.536 Ba o.126 0.141 0.318 0.126 0.007 0.323 0.285 0.276 o.242 0.267 0.455 0.183 o.284 0.286 o.251 0.360 0.357 o.317 0.271 0.386 S Dral I 4-688 4.980 4.914 4.962 5.m6 4.92E 4.9E0 4.996 4.9)3 4.9E4 4.98ð 4.982 4.982 4.919 4.975 SiAI 4.094 3.991 4.055 3.998 3.983 4.018 3.998 4.001 3.988 3.997 3.937 3.988 4.O02 4 4.001 4.001 3.923 3.979 3982 3.998 CNKB 0.594 0.989 0.8s9 o.964 1.023 0.91 0982 0.995 1.005 0.987 1.051 0.994 0.98 o.979 0.974 0.967 1.105 0.996 l.o2 0.986 aCn o.2l2l 0.1426 o.3702 0.1307 0.0068 0.3549 o.2902 0.2774 0.2408 0.2705 0.4329 0.1841 0.2898 0.2921 0.2577 0.3723 03231 0.3183 0.2657 0.3915 aOr 0.7189 0.817 0.4366 0.7988 0.0039 0.2s93 0.6599 0.6744 o.6925 0.ó636 0.2t063 0.7646 0.6653 0.6527 0.6869 o.5574 0.4688 0.5703 0.273s 0.s436 aAn o.0236 0.01I I 0.0419 0.0239 0.955 0.0769 0.001 0.001 0.0229 o.0223 0.1313 0.0312 0.0031 0.00ól 0.0133 0.03 0.t719 o.0592 o.4363 0.0355 aAb 0.0455 0.0293 0. t5 l3 0.0467 o.0342 0.3088 0.0489 0.u72 0.043E 0.0436 0.0295 0.0201 0.0418 0.049 0.0ø'21 0.0403 0.0362 0.0522 0.0245 o.0294 , Calc-silicates from Framnes Mountains ctn tu, woll SCAD dioo afi woll nm.¿tn rim.wol with.wol svm.qfs svm-lfs ím.otz wilh.wol 39.05 4:5.26 4:J.41) 49.IJ 39.31 50.34 43.4t) 42.37 4E.89 T1o2 0.04 0.0r 0.16 0.16 o.L7 0.03 0.09 o.02 0.10 o.24 0.08 0.03 41203 27.32 2't.36 27 2t.67 35.57 22.26 35.31 34.96 1.26 19.63 22.23 30.96 2.45 34.96 35.59 24.27 Fe2O3 FeO 0.25 o.o2 l.6l 0.14 L.32 o.32 0.,10 10.53 3.51 1.7 0.18 12.39 0.32 0.05 0.66 0.18 MnO 0.22 0.r8 0.05 0.30 0.12 0.05 0.15 0.19 0.46 0.59 0.15 0.08 0.r3 Mgo 0.07 0.23 0. 0.16 t1.46 0.11 0.06 9.56 o.2l CaO 18.30 18.43 17.98 37.22 t9.67 37.M 37. 18.98 19.73 47.69 24.90 3s.64 35. 22.76 23.ffi 19.90 19.75 47.26 26.97 Na2O 2.t2 1.97 1.98 0.01 0.55 0.06 o.47 0.rm 0.46 o.46 o.34 0.,1O o.52 0.lr K20 o.74 0.73 0.73 0.48 0.13 0.05 Cr2O3 0.40 0.38 0.72 0.59 0.06 o.27 0.42 0.90 99.9 9ó.2 99.U 98.7 94.: y).9 v).2 9E.4 9E.2 Si 6.9t2 6.966 6.968 2.965 2.0ú 2.936 2.9s8 2.033 2.O33 0.97t L.934 3.013 2. 6.146 1.931 2.O33 2.011 o.974 2.944 Ti 0.005 0.006 0.009 0.006 0.001 0.001 0.005 0.01 0.002 0.001 AI 5.07'1 5.0s2 r.94 1.959 t.970 1.982 1.955 t.929 0.057 t.784 5.705 0.111 1.929 t.974 2.008 Fe3+ Fe2+ 0.033 0.003 0.102 0.006 0.082 0.043 0.013 0.067 o.377 0.226 0. 0.o24 0.397 0.013 0.002 0.011 0.01I Mn 0.o29 0.023 0.007 0.019 0.007 0.015 0.002 0.005 0.00ó 0.030 0.019 0.006 0.003 0.002 Mg 0.008 0.026 0.012 0.653 0.012 0.013 0.546 0.006 Ca 3.090 3.093 3.028 0.985 2.979 0.955 0.990 1.009 1.020 2.944 2.E91 3.811 0.970 0.990 0.982 1.008 2.028 Na 0.647 0.597 0.606 0.002 0.050 0.002 0.043 0.036 0.016 0.I,m 0.025 0.036 0.u7 0.004 K 0.148 0.146 0.147 0.096 0.006 0.003 Cr 0.024 0.023 0.177 0.037 0.023 0.033 0.012 0.014 8.0ó5 4.y)5 5.ü)6 2.064 15.935 4lIn 5.007 5.022 2.005 0.7953 0.8063 0.801 0.9ó48 o.634 o.94t7 o.s79 0.9648 0.9543 (Xme) ffdi) CXan) I 964-4C from shear Fox biotite felsoar RArnet Szlab SZ.st cons.gt with.qz pl.sZ Ml SZfab rim.oox Mlcore Mlrim SZ.oox SZ.oox sio2 36.86 37.r3 48.01 47.53 48.66 38.46 39.20 39.00 3s.t9 38.61 38.58 38.47 50.66 51.83 50.00 51.4s 52.40 T1o2 4.84 4.63 0.04 AJ2o3 15. l6 15.69 33.34 33.95 33.47 23.31 22.84 22.70 21.O2 22.83 22.66 22.s2 2.77 1.84 3.13 1.72 2.38 FeO r5.59 16.61 0.09 o.29 32.23 32.t8 32.63 32.81 32.59 33.23 33.82 30.27 30.01 29.78 28.61 28.72 MnO 0.04 o.t7 o.32 0.10 1.39 1.10 t.43 1.55 t.45 1.46 1.40 o.42 o.34 o.32 0.35 0.17 Mgo 14.30 t3.45 0.15 0.33 5.93 5.68 5.39 4.98 5.93 5.O2 4.55 17.93 18.98 18.02 18.94 19.21 CaO 0.22 0.29 16. t8 16.94 16.55 2.41 2.69 2.78 2.12 2.45 2.66 2.ffi o.34 0.42 0.48 0.4 0.33 Na2O 2.20 l.9l 2.78 o.l2 0.10 0.13 K20 9.77 9.66 )tâl 196;1 9'1.5 99-9 1U0.9 trJ2.2 103.7 103.7 103.9 9E.3 104.0 103.7 103.4 LU2-4 103.5 I Si 2.728 2.734 2.2ffi 2.165 2.189 2.927 2.977 2.960 2.918 2.939 2.955 2.963 t.912 1.930 1.898 1.943 1.939 Ti 0.269 o.256 0.001 0.011 AI 1.323 t.362 1.801 t.823 1.775 2.Wt 2.045 2.M0 2.020 2.M9 2.M6 2.M5 o.t23 0.081 0.1210 o.o77 0.104 Fe2+ 0.965 t.023 0.003 0.011 2.051 2.U4 2.071 2.237 2.075 2.128 2.178 0.955 o.934 0.945 0.904 0.889 Mn 0.002 0.007 0.012 0.004 0.æ0 0.071 o.w2 0.107 0.093 0.095 0.091 0.013 0.011 0.010 0.011 0.005 Mg t.577 1.4't6 0.000 0.010 o.022 0.672 0.643 0.609 0.605 0.673 0.573 0.523 1.009 1.053 1.019 1.066 1.059 Ca 0.018 0.023 0.795 o.82',1 0.798 o.197 o.219 0.226 0.18s 0.200 0.218 o.214 0.014 0.017 0.019 0.018 0.013 Na 0.195 0.169 0.243 0.017 0.015 0.009 K 0.922 0.907 otal I 7.802 7.7E3 4.997 5.009 5.043 8.028 8.000 8.009 8.072 8.045 8.030 8.015 XFe 0.38 0.41 0.80 0.83 0.77 0.6816 0.6866 0.6909 0.7t36 0.6824 0.706 0.724s 0.9472 0.8873 0.9276 0.8474 0.839 964-54: Petite shear Fox biotite feldsoar Pamet cordierite soinel ksp.SZ rim.gt gtpull- cons.sil SZfab Szfab inc.st elonsatÉ sio2 37.08 35.79 35.61 36.98 36.66 63.60 55.36 61.87 39.02 38.47 38.55 38.87 38.30 38.57 37.99 49.UI 49.t7 49.26 0.05 37.34 Tlo2 4.41 2.03 1.34 2.33 3.30 0.31 0.11 0.03 0.08 0.33 o.24 ¡J2o3 t7.63 19;10 19.99 r8.85 18.45 r8.67 28.86 23.72 22.69 22.63 22.66 22.95 22.76 22.54 22.19 33.55 33.48 33.s 58.44 56.72 62.tt Fe2O3 FeO 16.1 1 r8.85 17.90 r4.t3 t6.57 2.t4 0.69 35.54 35.51 35.62 35.t7 35.84 36.2L 35.22 7.26 7.73 6.52 34.83 35.80 0.60 MnO 0.09 0.02 0.13 0.11 0.83 1.03 0.88 0.89 1.07 0.98 0.59 0.08 0.18 0.04 0.14 Mgo 13.33 11.53 12.01 15.24 Lt.94 0.ll 3.28 4.91 4.55 4.51 5.04 4.50 3.87 5.07 9.25 8.94 9.46 5.46 3.s3 0.13 CaO 0.36 0.44 0.21 0.42 0.34 0.56 5.75 4.58 t.32 1.46 1.48 1.44 1.36 1.39 1.16 0.2s 0.21 0.26 0.10 0.r8 Na2O Lt4 5.03 8.57 0.01 0.20 1.32 1.84 K20 9.67 9.63 9.55 9.ó8 9.62 15.35 0.05 Cr2O3 o.76 1.45 )ral I 98.7 98.0 96.6 91 .6 96.9 99.1 1U0.5 v).6 104.3 103.6 103.7 t0ø..4 103.9 103.5 rO2.2 99.5 99.E v).7 100.9 v)3 si 2.685 2.642 2.651 2.677 2.705 2.952 2.466 2:t56 2.976 2.962 2.965 2.96t 2.945 2.979 2.9ffi 4.972 4.977 4.966 0.001 1.003 Ti 0.240 0.1 13 0.075 0.127 0.183 0.011 0.006 0.002 0.006 0.02s 0.005 AI (-G 1.504 1.7 t4 1.754 1.609 1.605 1.022 1.516 t.246 2.W 2.054 2.055 2.Mt 2.063 2.052 2.038 4.O07 3.995 3.997 1.9,!0 1.948 1.967 Fe3+ Fe2+ 0.975 r.164 1.114 0.856 t.022 0.080 o.026 2.268 2.286 2.291 2.2N 2.305 2.339 2.295 0.6r5 0.655 0.550 0.820 o.872 0.013 Mn 0.00ó 0.001 0.005 0.004 0.054 0.M7 0.058 0.057 0.070 0.064 0.039 0.007 0.016 0.003 0.012 Mg I.438 1.268 t.332 r.644 1.3 l3 0.007 0.218 0.558 0.522 0.517 o.572 0.516 0.445 0.589 1.397 1.348 t.421 0.229 0.153 0.018 Ca 0.028 0.034 0.017 0.032 0.o27 0.028 o.274 0.218 0.10E 0.120 o.t22 0.117 0.rtz 0.115 0.w't 0.027 0.022 0.028 0.003 Na 0.103 0.434 0.740 0.001 0.039 o.o72 0.104 K 0.894 0.907 0.907 0.894 0.905 0.909 0.003 )tal l't-770 7.842 7.850 7.838 7.762 5.032 4.993 4.v)3 E.()o4 E.Oll 8.008 E.UD 8.017 1.995 8.019 rI.025 II.020 ll.03u 3.0ó5 3.07E XFe 0.40 0.48 0.46 0.34 0.44 0.03 0.39 0.23 o.7s9t 0.7632 0.7669 0.7499 0.7676 0.7893 0.7601 0.3059 0.3268 0.279 0.7818 0.8504 964-58: Fox biotite cordierite kspar Parnet rim.cd small.sz cor.sil v/ith.sil cons.cd late.gt cons.sil coreMl rimMl with.bi inc.sill sz.fab coreMl sz 37.66 37.34 37.11 37.56 37.24 o.t7 o.2l sio2 33.s0 33.46 33.68 34.47 47.46 47.20 46.45 46.99 47.31 47.38 46.88 46.75 47.34 ó1.90 o.o2 0.11 0.28 0.15 Tlo2 3.42 3.79 2.49 2.9t 0.04 0.04 o.02 0.08 0.18 19.20 2L.99 22.27 22.t7 22.r8 22.42 56.56 55.94 At203 17.86 17.66 20.95 18.25 32.50 32.61 33.45 32.49 32.60 32.76 32.2t 32.27 3t.73 34.M 33.94 34.82 36.03 34.t6 32.28 32.86 FeO t8.'15 19.53 t7.70 t7.93 7.40 7.92 7.71 8.51 8.01 8.53 7.02 8.44 0.08 t.r7 0.93 0.95 o.92 o.62 0.09 0.04 MnO 0.10 0.01 0.05 0.04 0.ll 0.20 0.16 0.20 0.04 0.1ó 0.10 0.23 4.39 5.O2 4.14 4.22 4.73 3.69 3.84 Mgo 8.73 9.03 10.86 t0.32 8.67 8.71 8.21 7.94 8.04 8.27 8.34 8.52 7.77 0.18 1.36 1.30 1.36 0.60 1.34 0.11 CaO 0.31 o.23 0.08 0.03 0.17 0.11 0.03 0.05 0.93 3.10 3.4t Na2O 0.08 0.16 0.03 0.01 0.04 K20 9.33 9.23 9.09 9.38 15.30 1.31 0.91 92.0 92.7 95.1 e3.4 97.E 100.6 100.8 100.6 l0I.ó 100.ð 4.938 5.O22 2.899 2.980 2.945 2.950 2.959 2.938 0.005 0.006 Si 2.65s 2.638 2.55t 2.669 4.980 4.944 4.886 4.954 4.972 4.951 4.984 0.006 0.017 0.003 Ti 0.204 o.225 0.142 0.169 0.000 0.m0 0.003 0.003 0.000 0.001 0.006 0.006 0.001 2.U11 2.U17 2.060 2.085 1.960 1.938 AI 1.669 1.641 1.870 1.666 4.021 4.02'1 4.t47 4.038 4.O39 4.036 4.036 4.Ot7 3.968 1.018 2.O5t 0.003 2.252 2.239 2.3t5 2.373 2.254 0.794 0.808 Fe2+ t.243 1.288 l.r2l l.l6l 0.649 0.694 0.678 0.750 0.7M o.746 0.624 0.746 0.727 0.078 0.062 0.0ú1 0.061 o.M2 0.002 0.001 Mn 0.007 0.003 0.002 0.000 0.009 0.017 0.015 0.018 0.003 0.014 0.009 0.020 0.518 0.590 0.490 0.495 0.556 0.162 0.168 Mg 1.031 1.061 1.226 l.t9r 1.356 1.359 t.286 t.247 t.259 1.287 t.32t 1.3,10 1.229 0.009 0.085 0.1 15 0.110 0.116 0.051 0.113 0.003 Ca 0.026 0.019 0.007 0.003 0.0r9 0.013 0.004 0.006 0.004 0.918 o.t77 0.194 Na 0.017 0.o32 0.007 0.002 0.007 o.M'l K 0.943 0.928 0.879 0.927 7.778 '1.780 7.811 7.792 0.7461 o.775s 0.7963 0.7603 0.8259 0.8269 XFe 0.55 0.55 0.48 0.49 0.3237 0.3381 0.3452 0.3?57 0.3s87 0.3669 0.321 0.3575 0.3719 0.760t 964-8C: Biotite-sillimanite Fox shear zone biotite cordierite felúpar wraD.cd adi.cd cons.cd con.Ml.i oomh w.bi.si corc.lse rim.lse l¿¡ce con.st wo-bi-sil conscd conMlsl M1 sio2 33.51 33.14 32.96 33.ó1 46.26 46.8 46.55 46.16 46.54 45.67 57.41 57.14 57.72 36.63 Tioz 2.64 2.47 2.56 2.60 0.04 0.01 0.09 o.t2 0.13 A1203 18.92 18.87 19.00 19.13 3t.92 30.98 31.95 3t.54 31.45 32.U 23.68 23.76 23.66 21.78 FeO 19.33 19.66 19.56 20.31 8.97 8.29 8.80 8.85 9.32 9.21 0.03 0.09 33.93 MnO 0.11 0.17 o.23 0.r2 o.26 0.57 o.44 0.46 0.47 o.42 0.06 0.14 2.20 Mgo 8.18 8. l6 7.87 7.58 6.77 6.92 7.05 6.93 7.Os 6.76 3.73 CaO 0.14 0.15 0.18 o.23 0.12 o.l2 o.l2 0.04 0.10 0.07 5.81 s.94 5.88 1.05 Na2O 1.23 0.3ó o.27 0.18 o.M 0.58 7.33 7.70 7.56 K20 9.18 9.40 9.09 9.31 0.03 0.07 )tal | 92.0 92.0 91.4 92.9 95.5 93.3 95.2 94.2 9s.4 94.7 94.3 949 95.0 Si 2.653s 2.6366 2.6345 2.650 4.953 5.023 4.979 4.992 4.988 4.92r 2.700 2.682 2.698 2.952 Ti 0.1572 0.1477 o.1537 0.154 0.003 0.001 0.003 0.0ûl 0.008 AI 1.7663 r.7699 r.7899 1.778 4.O29 3.980 4.029 4.021 3.974 4.069 1.313 1.315 l.3u 2.070 Fe2+ 1.2805 1.3077 1.3074 1.339 0.803 0.756 0.787 0.800 0.836 0.829 0.001 0.003 2.287 Mn 0.0074 0.01l3 0.0152 0.008 0.o24 0.053 0.040 0.042 0.043 0.038 0.003 0.006 0.150 Mg 0.965s 0.9679 0.9374 0.891 1.081 t.124 1.124 t.Ll7 t.126 1.084 0.,148 Ca 0.0121 0.0t27 0.0155 0.019 0.014 0.014 0.013 0.005 0.011 0.008 o.293 0.299 o.294 0.091 Na 0.255 0.076 0.055 0.038 0.82 0.t21 0.668 0.701 0.ó85 K 09271 0.9537 0.9264 0.936 0.002 0.004 râl I 7.770 7.808 7.780 7.775 11.160 11.025 lI.03l ll.0ló ll.u70 ll.07u 4.978 5.009 4.99{) XFe 0.57 0.57 0.58 0.60 0.43 O.,to 0.41 0.42 0.43 0.43 0.30 0.30 0.30 o.77 964-2OCz Fox shear biotite cordierite feldspar crack.st with.cd with.cd with.Dl cons.st late tE near st sz fab sz fab w. bi.ilm la¡se adi st cons qt core.sz core Ml si02 35.31 35.60 35.90 35.58 35.70 46.68 48.69 47.99 47.ffi 47.87 55.67 54.22 55.48 48.00 37.42 36.78 36.80 37.50 37.4 37.89 Tioz ó.11 4.74 5.10 5.05 5.12 0.15 o.25 0.15 0.08 o.24 0.04 0.15 o.24 o.22 0.,10 o.2l 0.18 0.05 0.08 AJ2o3 14.65 15.2t 15.09 15.07 14.80 33.09 33.36 33.06 32.51 32.39 26.77 26.78 25.80 31.39 22.3r 22.26 21.65 22.05 2t.99 22.23 FeO 12.94 13.57 13.88 14.07 14.03 5.64 5.59 6.69 6.09 8.03 0.16 0.69 0.1I 32.4t 32.93 34.02 32.59 32.97 33.06 MnO 0.10 0.05 0.20 0.05 0.08 0.10 0.14 1.03 0.89 l.ll 0.82 0.88 0.82 Mgo t2.19 13.00 t3.62 13.38 13.54 9.56 9.83 9.29 9.2t 7.94 0.04 0.13 5.24 5.23 4.05 5.53 5.32 5.67 CaO 0.30 0.13 0.07 0.31 0.15 o.L2 0.07 0.13 0.14 0.05 8.92 9.69 8.04 14.65 1_58 1.62 t.54 1.51 1.62 t.62 N¿O 0.09 6.25 5.75 6.68 2.79 K20 9.43 9.34 9.6t 9.58 9.38 0.19 0.02 o.o2 0.03 )ral I 90.9 91.7 93.3 93.0 92.8 95.4 97.E 97.4 95.7 96.5 97.9 96.9 97.O 97.5 si 2.754 2.75s 2.737 2;126 2.739 4.903 4.975 4.956 4.985 5.010 2.552 2.5t9 2.570 2.250 2.947 2.922 2.959 2.963 2.gffi 2.957 Ti 0.358 0.276 0.293 0.291 0.295 0.012 0.019 0.012 0.006 0.019 0.001 0.005 0.008 0.008 0.o23 0.013 0.011 0.003 0.005 AI 1.347 1.388 1.356 1.362 1.339 4.097 4.019 4.024 4.O14 3.996 1.447 1.467 l.Æ 1.735 z.UIL 2.085 2.O52 2.O53 2.O50 2.046 Fe3+ Fe2+ o.844 0.878 0.88s 0.902 0.900 0.495 0.478 0.578 0.534 0.703 0.006 0.027 0.004 2.135 2.187 2.288 2.t53 2.180 2.158 Mn 0.006 0.003 0.0r8 0.000 0.005 0.000 0.000 0.003 0.m4 0.005 0.069 0.060 0.076 0.055 0.059 0.054 Mg t.4t7 1.499 1.547 1.528 1.548 1.497 1.497 1.430 1.438 1.239 0.003 0.009 0.615 o.620 0.485 o.652 0.627 0.660 Ca 0.025 0.01I 0.006 0.026 0.013 0.013 0.008 0.014 0.016 0.005 0.438 o.482 0.399 o.736 0.134 0.138 0.133 0.128 0.138 0.135 Na 0.018 0.555 0.518 0.600 0.254 K 0.938 0.922 0.935 0.93ó 0.918 0.011 0.001 0.001 0.002 Iral I 7.683 7.136 1.159 l.'t'|tJ 't;t55 XFe 0.37 0.37 0.36 0.37 0.37 0.2485 0.242 0.2877 0.27Ut 0.362 O.44 0.48 0.,10 o.74 o.723t 0.7278 0.7673 0.7205 0.7258 0.7177 (Xan) APPENDIX 4 Representative P-T estimates us ing Thermocalc Pressure-temperature estimates for pelitic assemblages made using the average pressure-temperature approach of Holland & Powell (1985, 1988) and using the internally consistent dataset of Holland & Powell (1990). Fox Ridge gt-opx gneiss...... ry-1 Fox Ridge pelites...... IV-2 - 8 Trost Rocks pelites...... IV-g - 12 Jetty Pemimsula pelites .IV-l3-15 Appendix 4, Thermocalc outPut P-T ESTIMA , FOX E AVERAGE TEMPERATURE CALCULATIONS Fox Ridge, Garnet-Opx Gneiss, Sheared Assemblage, Sample 964-4C an inconplete independenE set of reactions has been calculated for sd(T) limit = 60oC : Rock name : 4C (suggested P = 6.0 kbar) (for a(IÐO) = 0'2) ksp gr PY alm q iIm ru II2O L. 00 0.200 a 1.oO 4.15e-4 0.0081-5 0.319 1.00 L.00 0 sd(1n a) o 0.77074 0.5'742]- 0 ' 09028 0 0 an en fs mgls Ph1 ann easc 0 0121 4 0.76'7 0.255 o .L77 o. oLL4 o.0954 0.0197 ' 0'47526 sd(ln a) o.01-629 0.16651 0.20734 0.8?719 0.28526 0-4795]- react'Íons 1) 6q + 2Ph1 = 2ksp+21120+3.sL 2\ 2alm + 6q + 2plr1 = 2ksp + 2p.Y + 2H2O + 3fs 3) alm + 3ru + ph1=ksp+PY+3Í1m+H2O 4) alm + 3ru + ptrL = ¡ Diagrnos Eic lnf or¡naEion av, sd, fiE are results of doubling (first tsable) / haivíng (second table) uncertaÍnEy-on ln a : ã itt suspecE if any are v differenE from 1sq values' à* "-i"In a rèsiduals nõnnallsed Eo 1n a uncerEainÈles : i"rge-.b."fuEe"t" values, say >2.5,- n-oin! to suspecE lnfo' hat are Ehe diagonal elsnents of Ehe haE naErjx : large values, sáy >0.38, pol1t-tso influential data' Foi'gSc confider¡ce, fit (= sd(fit) ) should be less Etta¡t 1'54 av sd fits lsq 692 54 t.67 T sd fir e* hats ksp 692 54 L.67 0 0 gr 686 55 1.64 0.7 0.11 py 720 45 L.30 2.t 0.17 aIm 696 55 1.65 -0.3 0. 03 q 692 54 t.67 0 0 i1m 692 54 7.67 0 0 ru 692 54 r.67 0 0 tno 692 54 t.67 0 0 an 692 54 1.67 -0. 0 0. 00 en 666 51 1.44 -]-.4 0.15 fs 689 49 r.a¿ -1.5 0.01 rrìgts 698 53 1.61 L.0 0.05 phl 6'74 74 1.64 0.4 o .46 8.0 P ¿.0 ¿.5 5.0 5.5 6.0 5.5 7.0 7.5 rvlf 6t2 65{ 667 679 692 706 7L9 733 7¿7 sd 65 6T 57 55 54 55 5'l 6T 66 f 2.r 1.9 1.8 1.7 7.'7 t.'l 1.8 10 z.v rv-t þpendit 4, Thermocalc outPut AVERAGE PRES SURE CALCI.JLATIONS Fox Ridge, M1 Assemblage, Pelite Sample 964-58 an ÍndependenE set of reactions has bee¡r calculated Rock name : 58tr"f1 (suggesEed T = 750oC) (for x(CO2) = 0.L and x(I{2o) = 0.5) py alm S i1n ru sp a 0.0 072! 0.394 1. 00 1.00 1.00 0. 170 sd(a)/a 0.5 8585 0.15000 0 0 o 0.20667 silI crd fcrd a 1.00 0.374 0.101 sd (a) /a 0 0.11756 0.27923 reactions 1) 5q + 2sP = s¡d 2\ 3crd = 2py + 5q + 4si11 3) 3fcrd = 2alm + 5q + 4si11 4\ 2q + 3ilrn + sil1 = alm + 3ru calculaEions P (T) sd dP/dT InK 1 7.0 1 03 0.0049 2.560 2 5.2 0 69 0.0026 -6 9L4 3 6.4 0 50 0. 0093 5 015 4 7.0 0 89 0. 0065 -0 931 Rock 5BM1 : averaçte pressures (for x(co2) = 0.1 and x(H2o) = 0.5) Le¿st Median of squares - rmr1tsiple end-msriber diagnosEics Assunes tÌÞt, 9 activities, chosen randornly, are OK, tÌ¡en solves for the rsnainder plus P and/or T' BesE resulEs wlren median of activiEy residuals is leasE (ÈfS). Gives proEecElon againsE ¡næc of 1 bad acEivities. SuggesEed outliers only sérious if associated wit]1 differenE condltions or lf LlfS is large (>2?) sqrt Lltf,S = 0.83 giving P = 6.8 kbars noE solvable = 7t rþsÈ featured oullier is gr of the best l4S resulEs (11[5<1 .5 Lì4S), average P = 6.5ldcar, with scatEer 0.8 )dcars (2 signa on disErÍbutÍon) Single end-meriber dlagnostic informatslon av, sd, flt are result of doubling Lhe u¡reertainty on Ln a : a ln a is suspect if any are v different from lsq values. e* are 1n a residuals nornalised to 1n a uncerbaÍnEies : large absoluEe values, say >2.5, point to suspect info. hat a¡e tJre diagonal elqnenEs of the hat, nratrjx : large values, say >0.44, poinE to influentslal data- For 959 confidence, ftts (- sd(fÍts)) should be less than 1.51; however a larger value nray be OK - look aE tlre diagrnosticsl av sd fiL lsq 6.\9 0 42 1.13 P sd fít e* hat pv 6 44 0 32 0 77 1 4 0.23 aIm 6 05 0 42 I 02 0 7 0.17 q 6 19 0 42 1 13 00 ilm 6 19 0 42 1 13 00 ru 6 L9 0 42 1 13 00 sp 6 10 0 40 1 02 -0 9 0.08 si11 6 79 0 42 L 13 00 crd 6 23 0 44 1 12 -0 2 0.05 fcrd 6 1,6 0 51 L 13 0 1 0.35 Toe 600 650 700 750 800 850 900 lv P 5.6 5.8 6.0 6.2 6.4 6.7 6.9 sd 0.43 0.39 0.39 0.42 0.48 0.56 0.65 f 1.3 t.2 1.1 1,.1 r.2 7.4 1.5 T\I-2 Appendix 4, Themocalc out7ut AVERAGE PRES SURE CALCULATIONS Fox Ridge, M1 Assemblage - Pelite, Sample 9&-58 an independenE set of reactions has been caLculated Rock name : 5Bù11 (suggested T = ?50oC) (for x(CO2) = 0.1 and x(lI2O) - 0.4) py alm q ilm ru sp a 0.0 072! 0.394 1. 00 1. 00 1.00 0. 170 sd(a) /a 0.5 8585 0. L5000 0 0 0 0.20667 sill crd fcrd a L.00 0.3'14 0.101 sd(a) /a 0 0.1L755 0.27923 reacEio¡rs 1) 5q + 2sP = s¡d 21 3crd = 2py + 5q + 4si11 3) 3fcrd = 2a1m + 5q + 4si11 4) 2q + 3ilm + siI1 = alm + 3ru calculaEions P (T) sd dP/dT 1nK 1 6.4 1 03 o.0042 2.560 2 4.9 0 69 0.0023 -6.9]-4 3 6.0 0 50 0.0089 5.015 4 7.0 0 89 0. 0065 -0.931 Rock 5Bù[l : average pressrrres (for x(CO2) = 0.1 a¡rd x(H2O) = 0.4) LeasE Median of Squares - mulEiple end-mgriber diagrnosEics Àssunes tstlaE 9 acEivieies, chosen ra.ndomly, are oK, then solves for the rsnainder plus P and/or T. BesE resulÈs r,*re¡r median of acÈivity residuals ís leasE (LùfS). Glves protection againsE ¡nax of 1 bad activiEies. Suggested outliers only sèrious lf assocÍateil with different condiEions or lf Ll{S is large (>2?l sqrE LlttS = 0.89 giving P = 6.4 ldcars not solvable = 7t rþsL feaEu.red ouElier is gr of Ltre besE Ì[S resulEs (MS<1 .5 LÙfS), average P = 5.21òar, with scatstser 0.? )dcars (2 stgrna on disÈ.ribubion) Singte end-msriber diagrnostic lnf onnaÈlon av, sd, fiE are resulE of dor¡bling the uncerbainty on ln a : a ln a is suspect, ff any are v dlfferent fron lsq values. e* are In a residuals norrnalised to 1n a uncerEalnEies : large absolube values, say >2.5. points to suspect info. hat a¡e the diagonal elsnents of Etre hat ¡naLrix : large values, say >O.44, points Eo influenEial daEa. For 959 confidence, fits (- sd(fiE)) should be less tltan 1.61; hower¡er a larger value may be oK - look at the diagrnosElcs! av sd fib 1sq 5.83 0.44 t.l7 P sd fit e* hat' pv 6 07 0 36 0.87 t.4 0.23 alm 5 63 0 40 0.96 -1.0 0.17 q 5 83 0 44 t.ll 0 0 ilm 5 83 0 44 r.t1 0 0 ru 5 83 0 A4 I.7'l 0 0 sp 5 77 0 44 L.13 -0.6 0.08 si11 5 83 0 44 r.t'l 0 0 crd 5 88 0 46 1.15 -0.2 0.05 fcrd 5 84 0 53 L.t7 -0.0 0.35 ToC 600 650 700 7 50 80 850 9 0 0 lv P 5.2 5.¿ 5.6 5 .8 6.1 6.3 6 6 sd 0.41 0.39 0.40 0.,44 0.50 .57 0. 6 6 f 1.3 1.1 1.1 1. 1.3 t.4 1. 5 ry-3 Appendix 4, Thermocqlc output AVERAGE PRES SUREÆEMPERATURE CALCTJLATIONS Fox Ridge, Shear Zone Assemblage - Pelite, Sample 964-58 an independent. set' of reactions has been calculated Rock nane : 5Bsz (for x(CO2) = 0.8 and x(If2O) = 0.1) ksp py alm q ilm ru TI2O si11 a 1.00 0.00415 0.436 1 00 L.00 1 00 0 L00 1. 00 sd (a) /a 0 0.53398 0.15000 0 0 0 crd fcrd phl ann east a o.414 0.0897 0.0324 0.0348 0.0319 sd (a) /a 0.10302 0.29442 0.40344 0.40826 0.40481 reactions r,) 3fcrd=2alm+5q+4si11 2l 2q+3ilm+sil1=a1m+3ru 3) 4crd + 5eas! = py + 12si11 + 5ph1 4) 4si11 + 2ph1 = q + crd + 2east, s) 6fcrd + San¡r = 5ks¡r + 9alm + 5H2o + 3si11 6) 5a1m + 9q + 3east = 2W + 3fcrd + 3an¡r calculaElons P (T) sd dP/dr 1n-K 1 6.5 0 53 0.0095 5.574 2 7.5 0 89 0.0069 -0.830 3 6.0 1 34 0.0073 -L.879 4 7.8 2 09 0.0113 -0.913 5 4.2 0 86 -0.0019 23 -787 5 5.0 1 26 0.0141 -L3 -792 Rock SBsz : average E{l (for x(CO2) = 0.8 a¡rd x(H2O) = 0.1) LeasÈ Media¡r of Squares - mulÈiple end-msriber diagmostÍcs Assumes thâC 9 acEivities, chosern randomly, are OK, then solves for the rsnainder plus P and/or T. Best resulEs r,vhe¡r median of acElvlEy resÍduals is least (Lf'fS). cives proEection against ¡rac of 2 bad acEiviEies. SuggesEed out,lÍers only serious lf associated witsh different condibions or lf LlvlS is large (>2?) sqrE LlfS = 1.03 glving P = 6.7 )dca¡s and T = 7580C not. solvable = 2t ¡ncst. featu¡ed outlier is q Single end-msriber èÍagrnostic infor¡naEÍon avP, ar/I, sd's, cor, fiE are resulE of doubling the wrcerEainÈy on 1n a a ln a is sus¡recE lf any a¡e v different. frqn lsq values. e* a.r'e In a residuals nor¡naIÍsed Eo ln a uncerEainties : large absoluEe values, say >2.5, points to sus¡rect info. haE are the diagornl elsnenEs of the hat rnaErjx : Iarge values, say >0.46, poinE Eo influential data. For 95t confidence, fit (= sd(fit) ) should be less tìän 1.49 however a larger value rnay be OK - look at tJle diagmosEics ! avP sd ar/I sd cor fit rsq 4.8 0.8 636 57 0.8L0 t.69 P sd(P) T sd(T) cor fit e* hat ksp 4 75 0. 83 636 57 0.810 1. 59 0 0 pv 5 04 0.96 653 64 0.862 t.63 -o.82 0.39 aIm 4 31 0.78 6L5 51 0.83L 1.43 1. s3 0. 1.9 q 4 75 0. 83 636 57 0.810 1.69 0 0 ilm 4 75 0. 83 636 57 0.810 t.69 0 0 ru 4 75 0. 83 636 57 0.810 1.69 0 0 t120 4 76 0. 84 636 6t 0.806 1.69 0 U¿ 0.06 si11 4 75 0. 83 636 57 0. 810 \.69 0 0 crd 4 85 0.81 632 55 0.770 1.62 0 62 0. 10 fcrd 4 66 0. 89 634 57 0.794 t.68 -0 42 o.26 ph1 4 '15 o.82 630 60 0.762 7.67 -0 44 0.30 ann 5 06 0.78 663 56 0.837 r.49 a 55 0.23 easb 4 69 0.79 624 57 0.804 t.67 -0 98 0.13 T = 635oC, sd = 57 P = {.8 ltbrr¡, sd = 0.8, cor = 0.810, f = !.69 IV-4 Appendix 4, Thermocølc output AVERAGE PRESSURE CALCULATIONS Fox Ridge, M1 Assemblage - Pelite, Sample 964-84 an independenb set of reactions has been calculated Rock name : 8A (suggested T = 750oC) (for x(CO2) = 0.1 and x(II2O) = 0.4) gr py alm S ilm ru an si11 2.82e-5 0. 0l-52 0.325 1. 00 1.00 1 00 0.480 1. 00 sd (a) /a 0.98561 0.50934 0.1_5000 0 0 ô 0.08112 0 crd fcrd a 0 .44e 0.0941 sd (a) /a 0.0 9141 0.28835 reacEions 1) 3crd = 2py + 5q + 4sil1 2\ 3fcrd = 2alm + 5q + 4si1l 3) 2q + 3ilm + si11 = alm + 3ru 41 Aalm + 5ru + 3an = gr + SiIm + 3fcrd caIculaElons P (T) sd dP/dT In-K 1 5.5 0 .60 0.0030 -5.964 z 5.9 o.52 0.0088 4.842 3 6.1 0.88 0.0055 -1,.L24 4 6.9 0.89 0.0083 -10.869 Rock 8A : average pressures (for x(Co2) = 0.1 and x(l¡2o) = 0.4) LeasE Media¡r of Squares - rmrlEiple end-medber diagrnostlcs Àssu¡nes tt¡at 7 acEiviÈies, ctrosen randcrnly, are oK, tlren solves for Ehe re¡nainder plus P and/or T. Best results vritren media¡r of activity resi.duals ls 1easts (LÞfS). Gives proEection against ¡nax of 1 bad acÈlvities. suggesÈed ouEliers only serious if associated with differenE conditions or If LlfS is large (>2?) sqrt LùlS = O.42 giving P = 5.9 kbars not solrrable = 1t mosE feaEured ouLlier is a¡t of the best ltfs results (l4S<1 .5 Ll"fS), average P = 6.0)dcar, witsÌr scatEer 0.3 )cbars (2 sigrna on dlstrlbut.ion) Single end-msriber diagnosElc informaEion av, sd, fit are result of dot¡bling Ehe u¡rcerEainty on In a : a In a fs suspect, lf arry are v dlfferent frcrn lsq values. e* are ln a reslduals nor¡nalised Eo In a uncerEalnties : large absoluEe values, say >2.5, poinE to suspecE info. ha! are tJle diagonal elgnenÈs of Ehe haÈ rnaErix : large values, say >0.40, points to influent.ial data. For 95t confidence, fit. 1= sd(fit)) should be less tltan 1.61; however a larger value rnay be oK - look aE t'l.e diagnosticsl av sd fit lsq 5.53 0.37 0. 96 P sd fit e* hat gr 5 73 0.38 0.58 1 5 0 05 pv 5 67 0.43 0.96 0 0 33 alm 5 54 0.41 0. 91 -0 5 0 16 q 5 63 0.37 0. 96 0 0 ilm 5 63 0.37 0. 95 0 0 ru 5 63 0.37 0. 95 0 0 an 5 65 0.37 0. 90 -0 4 0. 00 si11 5 63 0.3'1 0. 95 0 0 crd 5 64 0.38 o.96 -0 n o.02 fcrd 5 55 0.44 0.94 0 3 0.33 ToC 650 675 700 725 750 775 800 925 85 0 lv P 5.2 5.3 5.¿ 5.5 5.6 5.8 5.9 6.0 6.2 sd 0.43 0.40 0.38 0.36 0.37 0.38 0.39 0.39 0.40 f 1.3 t.2 1.1 1.0 1.0 0.9 0.9 1.0 1.0 ry-5 Appendix 4, Thetmocalc output AVERAGE PRESSURE / TEMPERATURE CALCULATIONS Fox Ridge, Shear Zone Assemblage - Pelite, sample 9ó4-84 an independent se! of reactions has been calculaEed Rock name : 8Àl'12 (for a(co2) = 0.1 and a(I{2o) = 0.21 ksp gr py alm q i1m ru H20 a 1. 00 5.2 6e-5 0 .00322 0.469 l-.00 1.00 1.00 0.200 sd (a) /a 0 0.8 3382 0 .65382 0.15000 0 0 0 an si11 crd fcrd ph1 ann easE a 0.430 L. 00 0.459 0. 0865 0. 0598 0.0207 0.0328 sd(a) /a 0 -o9747 0 0.08780 0.2989't 0.34243 0.47362 0.40235 reacEions 1) 3a.n=gr+q+2si11 21 3fcrd = 2alm + 5q + 4si11 3) 2q + 3i1m + sill = alm + 3ru 41 4crd + Seast = py + 12sil1 + 5ph1 5) 9q + 6si11 + 2plr.1 = 2ksp + 2H2O + 3crd 61 3q + 3fcrd + ¿ann = 4ksp + 6alm + 4H2O 71 5gr + 5si11 + 3crd = 2py + 15a.n calculations P (T) sd dP/dT lnK 1 4.8 t.37 0.0101 -7.321 z 5.7 0.51 0.0090 5.829 ? 7.5 0 .87 0.0070 -o.757 4 6.3 t.20 0.0080 o.379 5 5.3 0 .43 0. 0163 3.297 6 5.5 1.09 -0. 0037 18.310 7 6.5 3.44 0.0227 2'7 .464 Rock 8Àl{2 : average HI (for a(CO2) = 0.1 and a(II2O) = 0.21 Least Median of squa-res - mlltslple end-mgrìber diagrnost.ics Assr.unes tàåE 10 acElviEies, chosen randomly, are oK. Ehen solves for Ehe rgnainder plus P and/or T. BesE resulEs v*ren median of act.ivity residuals is leasE (Ll4S). cives protection against max of 3 bad activities. suggested outliers only serious 1f associated wittr different conditions or lf Llvtrs ls large (>2?\ sqrt, I¡fS = 0.27 giving P = 3.4 ldca¡s and T = 5'77oC noE solrrable = 234 tnost feacured ouÈ1ier is phl Single end-mqnber êiagmostic infor¡ration avP, aúI, sd's, cor, fit are resulE of dotrbllng Ehe uncerEainEy on ln a a In a is suspecE if any are v differenE from lsq values. e* a.re In a reslduals norrnalised to In a txrcertainties : large absolute values, say >2.5, poinE Eo suspect info. hat a¡e ttre diagonal elsnents of ttre baE matri¡< : large values, say >O.47, poinE to influenbial data. For 95t confídence, fit (= sd(fit)) should be less than 1.45 however a larger value ray be oK - look aÈ the diagrnosbics! avP sd ar/I sd cor fit. 1sq 4.9 0.9 674 68 0.856 1.74 P sd (P) T sd (T) cor fit e* hat. ksp 4 91. 0. 93 674 68 0. 856 t.7 4 0 0 gr 4 92 0 94 674 68 0. 853 1 74 -0.23 0 04 ¡¡/ 5 77 0 94 727 67 0. 898 1 45 -2 -0L 0 42 aIm 4 27 0 81 640 57 0.872 1 37 2 -07 0 T9 q 4 9t 0 93 674 68 0.856 1 74 0 0 ilm 4 9I 0 93 6't4 68 0.856 1 74 0 0 ru 4 91 0 93 674 68 0.856 1 74 0 0 1120 4 91 0 93 674 68 0.856 L 74 0 n an 4 91 0 93 674 68 0.855 L 74 0.08 0 00 si11 4 9t 0 93 674 68 0. 856 1 74 0 0 crd 4 96 0 92 670 67 0.825 L 70 0.51 0 09 fcrd 4 84 L 00 672 69 0.848 l- 74 -0.30 0 23 ph1 4 92 0 o, 583 73 0. 811 L '73 0.47 0 36 ann 4 97 0 99 679 74 0. 873 1 '14 0.34 0 I9 easC 4 79 0 88 65'7 67 0. 853 1 63 -1.30 0 13 fr 674oe, sd = 58 P¡ {.9 l¡bu¡, sd = 0.9, cor = 0.856, f = t.'l 4 IV.6 Appendix 4, Thermocøh output AVERAGE PRESSURE / TEMPERATURE CALCULATIONS Fox Ridge, Shear Zone Assemblage - Biotite Sillimanite Gneiss, Sample 9&-8C an independenE se! of reactions has been calculated Rock name : 8C (for a(CO2) = 0.2 and a(H2O) = 0.5) ksp q ilm ru rr20 si11 crd fcrd 1.00 1. 00 1.00 1 00 0.600 1.00 0.3]-2 0.160 sd (a) /a 000 0 0 0.1-4200 0.21600 phl ann east a 0.0259 0.0393 0.0361 sd(a) /a 0.42242 0.39205 0.39366 reacbions 1) 3q + 2i1m + 2si11 = 2ru + fcrd 21 9q + 6si11 + 2plr.1 = 2ksp + 2Wo + 3crd 3) 5q + si11 + east = ksp + H2O + crd 41 9q + 6si11 + 2a¡ur = 2ksP + 21120 + 3fcrd calculaEions P (T) sd dP/dT 1n-K 1 4-8 0 .54 0.0087 -L. 833 z 5.1 0 .55 0.0152 3.813 3 4.3 0.77 0.0L72 2.L57 4 5.5 0 .6L o.0]-52 0.975 Rock 8C : average Pt (for a(CO2) = 0.2 and a(I{2o) = 0.6) LeasE Median of Squares - rmrltsiple end-msriber diagnosEics Àssrrrnes t}laE. 9 activltsies, chosen randomly, are OK, Ehen solves for the rgnainder plus P and/or T. BesE resulEs r,'¡hen media¡r of acÈiviEy residuals is least (L¡4s). Gives protection against nrax of 1 bad activlEles. SuggesEed outliers only serious j.f associaEed wlth differents condibions or if LlfS is large (>2?) qrE L¡{S = 1.0e+4 glvtng P = NaN l Single end-msriber d.lagmos tic inf orrnation avP, ar/l, sd's, cor, fiE are resulh of doubling the uncerEainly on 1n a a ln a is suspect if any are v different from 1sq values. e* are In a residuals nornralised to ln a uncerEainEies : large absoluEe values, say >2.5, point t'o sus¡rect info. hat are the diagonal elsnents of tJre hat rnaErix : large values, say >0.36, polnt bo influenbial claca. For 95t confidence, fit (- sd(fit)) should be less Ehan 1.61 hower¡er a larger value rnay be oK - look at the diagmostlcs! avP sd aúI sd cor fir lsq 4.4 1.0 704 67 0.935 0.81 P sd(P) T sd(T) cor fib s* hat ksp 4 40 0.96 704 67 0.935 0 81 0 0 s 4 40 0.95 704 67 0.935 0 81 0 0 ilm 4 40 0.96 704 67 0.935 0 81 0 0 l1r 4 40 0.96 704 67 0.935 0 81 0 0 H20 4 40 0.96 704 67 0.935 0 81 0 0 si11 4 40 0.96 't04 6'7 0.935 0 81 0 0 crd 4 39 0.96 699 69 0.9r7 0 79 0.15 o.t4 fcrd 4 18 1.26 693 79 o.946 0 79 -0. L9 0 .50 ph1 4 38 0.96 705 67 0.916 0 79 0.22 0.24 ann 4 77 1.11 735 81. 0.950 0 65 0.59 0.39 easE 4 23 0.98 685 7t 0.933 0 56 -0.76 0.16 fr 704oC, sd = 67 Pr {.{ lçbrr¡, sd = 1.0, cor = 0.935, f = 0.81 T\I-7 Appendfu 4, Thermocalc output AVERAGE PRESSURE CALCULATIONS Fox Ridge, Shear Zone Assemblage - Biotite Sillimanite Gneiss, Sample 9&-8C an inconpleEe independenE set of reacEions has been calculated Rock name : 8C (suggested T = 700oC) (for x(co2) = 0.3 and x(IÐo) = 0.7) ksp q iln ru 1120 si11 crd fcrd a 1.00 1. 00 L.00 1_. 00 0.700 1.00 o.3t2 0.160 sd (a) /a 0 0 0 0 0 o-74200 0.21600 phl ann east a 0.0259 0.0393 0.0361 sd(a) /a 0.42242 0.39205 0.39366 reactions 1) 4sill + 2phl = q + crd + 2east 21 2ilm + 14si11 + 6ph1 = 2ru + 3crd + fcrd + 6east 3) 3ilm + 21si1l + 10ph1 = 3ru + 6crd + ann + 9easC calculations P (T) sd dP/dr lnK 1 6.8 2 03 0.0109 -0.50L 2 6.4 1 59 0.0105 -3.335 3 6.2 1 69 0.0105 -3.583 Rock 8C : average pressures (for x(co2) = 0.3 and x(I{2o) = 0.7) teasL Median of squares - rmrltiple end-msriber diagrnosEics Assì.unes thaÈ 9 activiEles, chosen randornly, are OK, Ehen soLves for Ltre rsnainder plus P and/or T. BesE results raÈren median of acEivity residuals is leasÈ (f,¡dS). Gives proEecEion against ¡narc of 1 bad activiEles. suggested oulliers only serious if assoclaÈed with dlfferenL conditions or if Ll{S is Ia¡ge (>2?) sqrt LùfS = 0.91 giving P = 5.2 ktrars not solr¡able = lt mosÈ featu¡ed ouElier ls ksp of t¡e best ltlS results (tifs P sd fiC e* haE q 62 0 .51 1.13 0 0 ilm 62 0.51 1.13 0 0 l1r 62 0 .5L 1.L3 0 0 si11 62 0.51 L.13 0 0 crd 67 0.51 1.10 -0.2 0.02 fcrd 44 0 .68 1.09 0.3 0.48 ph1 58 0.53 1.11 -0.3 0.09 ¿utn 77 0.44 o.92 -0. I 0.09 east 63 0.35 0 .78 1. 1 0.00 ToC 600 625 650 675 700 725 750 775 800 tv P {.1 ,l .3 t.4 {.6 {.8 5.0 5.2 5.{ sd 0.49 0.50 0.50 0.51 0.51 0.52 0.52 0.53 0. 53 f 7.2 7.2 r.¿ 1.1 1.L L.1 1.1 1.1 1.1 rv-8 Appmdix 4, Thermocalc output AVERAGE PRESSURE CALCTJLATIONS Fox Ridge, M1 assemblage, Sample 9&-208 an independenE seE of reaclions has been calculaEed Rock name : (sugqesEed T = TOOoC) (for x(CO2) = 0.2 and x(H2O) = 0.4) gr py alm q ilm ru an si11 a 9.10e-5 0. 0108 0.356 1 00 1.00 1. 00 0.590 L 00 sd(a)/a 0.82086 0.54610 0.15000 00 0 0.05043 0 crd fcrd a 0.523 0.0629 sd (a) /a 0.06826 0.33678 reacÈions 1) 3crd = 2py + 5q + 4si11 21 3fcrd = 2a1m + 5q + 4si11 3) 2q + 3i1m + si1l = alm + 3ru 41 3q + 6ilm + 3an = gr + 2a1m + 6l:t¡ calculations P (T) sd (P) a sd a) b c In_K sd(1n-K) 1 4.9 0 50 39.33 2 1) 0.L 0071 -15.383 -7.LL2 1 L11 z 5.7 0 53 -59.51 1 33 0.1 0539 -L6.r54 6.233 1 054 3 6.4 0 68 L.52 0 39 0.0 193 I -t.87'l -1.033 0 150 4 5.2 0 80 -23.37 t 24 0.1 5358 -9.050 -9.787 0 887 Rock : average pressrlres (for x(Co2) = 0'2 and x(H2o) = 0.4) Single end-mqnber diagrnostic lnfor¡naElon av, sd, fit a¡e resul! of dor¡bling the wrcerEainty on 1n a : a ln a is suspecE if a¡ry are v differenc from 1sq values. e* are In a residuals nor¡¡alised Eo ln a uncerÈ.ainties : large absolule values, say >2.5, point to sus¡recE info. haE are the diagonal elgnenÈs of t.l.e haE ¡naErix : large values, say >0.40, poinE Eo influentsial data. For 95t confidence, fit 1= sd(fit)) should be less than 1.61; however a larger value may be oK - look a! the diagmosEics! av sd flÈ lsq 5.47 0.39 1.08 P sd fit s* hats gr 5.53 0.36 0 98 0.9 0.07 ptr 5.55 0 .39 0 95 0.9 0.32 aIm 5.22 0.31 0 76 -1.3 o.26 q 5.47 0.39 1 08 0 0 ilrn 5.47 0.39 L 08 0 0 ru 5.4',7 0.39 1 08 0 0 an 5.47 0.38 L 07 -o.2 0.00 siLl 5.47 0.39 1 08 0 0 crd 5.48 0.39 T 07 -0.2 0.01 fcrd 5.45 o.44 T 08 0.1 0.29 1!oC 620 5{0 660 580 700 7 20 7to 7 60 7e0 800 rv I 5.1 3.2 5.3 5. { 5.5 5 6 s.7 5 .8 5.9 6.0 sd 0.37 0.37 0.37 0.37 0.39 0 40 0.43 0 .45 0.49 0.52 f 1.1 L.1 1.1 1.1 1.1 L 1 r.2 1 .2 1.3 1.3 rv-9 Appendix 4, Thermocalc output P-T Estimates Trost Rocks AVERAGE PRESSURE CALCIJLATIONS TROST ROCKS, Mt ASSEMBLAGE, SAMPLE 964-TR8 an inconplete independenE seE of reactions has been calculated for sd(P) 1i¡niE = 4.0 ldcars : Rock name : TR8M1 (suggested T = 700oC) gr py alm s ilm ru a 5.70e-5 0. 0134 0.297 1. 00 1. 00 1.00 sd(1n a) 0.83206 0.52324 0.09968 0 0 0 herc sp an si11 d o.707 0 .292 0.505 1.00 sd(1n a) 0.0256'7 0.15017 0.07351 0 reactions 1) alm + 3a¡r = çtr + 6q + 3herc 21 2q + 3i1m + sill = alm + 3ru calculations P (T) sd dP/dT lnK 1 5.9 2.24 0. 0084 -7 .54',7 2 5.8 0.50 0.0050 -r.214 Rock lR8M1 : average pressìlres Diagrnos tic inf orrnaEion av, sd, fiE are resulE of dot¡b1ing (firsts cable) / halving (second Eable) uncerEainEy on 1n a : a In a is suspecE if any are v different from 1sq values. e* are 1n a residuals nor¡nalised Eo In a uncertainEies : large ab,so1uEe values, say >2.5, poinE to sus¡recE info. hat are tl.e diagonal elsnents of the haE ¡natri:< : large values, say >0.25, poinE to influentsial data. For 95t confidence, fit 1= sd(fiE)) should be less thä.n 1.96 av sd fit 1sq 5.'77 0.48 0.07 P sd fiE e* hat gr 5 .76 0.50 0.04 -0.1 0. 09 alm 5 .'79 0 76 0.07 0.0 0. 65 q 5 .77 0 8 0.07 0 0 ilm 5 77 0 I 0.07 0 0 ru 5 77 0 I 0.07 0 0 herc 5 77 0 I 0.07 -0.0 0.00 an 5 77 0 I 0.07 0.0 0.01 si11 5 77 0 I 0.07 0 0 P sd fit e* hats gr 5.79 0 44 0 t2 -0 1 0 09 aIm 5.76 0 33 0 08 0 0 0 65 q 5.77 0 48 0 0'l 0 0 ilm 5.'t7 0 48 0 07 0 0 ru 5.77 0 48 0 07 0 n herc s.'t7 0 48 0 07 -0 0 0. 00 an 5.77 0 48 0 08 0 0 0. 01 si11 5 -77 0 48 0 07 0 0 ?!oc 500 550 600 650 700 750 800 850 900 tv P { .8 5.0 5.3 5.5 5.8 5.0 6.3 5.5 6.9 sd 0.41 0.42 0.44 0.46 0.48 0.50 0.51 0.53 0.55 f 0.3 0.2 0.1 0.0 0.1 0.1 0.2 0.2 0.3 ry-10 Appendix 4, ThetmocøIc output AVERAGE TEMPERATURE CALCULATIONS Trost Rocks, Cordierite-Bearing Assemblage (Mz), Sample 964-TR11 An inconplete independent set of reactions has been calculated Rock nane : TR11l"f2 (suggested P = 6.0 kbar) (for x(CO2) = 0.3 and x(H2O) = 0.7) ksp q iIm ru 1l2O herc sp 1. 00 L. 00 1. 00 1.00 0.700 0.72t 0.2'79 sd (a) /a 0 0 0 0 0.05000 0.15595 si11 ph1 ann easE crd a 1.00 0.0902 0.0138 0.00104 0.586 sd (a) /a 0 0.29372 0.52003 0.620]-7 0.05142 fcrd a 0.0375 sd (a) /a 0.39012 reacEions 1) 9q + 6si11 + 2an¡r = 2ksp + 2H2o + 3fcrd 21 3sp + ann = 3herc + phl 3) 5m + 3sp + ann + fcrd = 5i1m + 5si11 + phl calculations T (P) sd dP/dT lnK 1 683 50 0.0171 r.284 2 650 276 o.0822 4.726 3 707 84 0.0191 8.990 Rock !R.11ID : average Esq)eratures (for x(CO2) = 0.3 and x(lI2O) = 0.7) LeasÈ Median of squares - mult.iple end-msriber diagrnostfcs Assrnnes thaE 15 aclivities, chosen randonly, are OK, then solves for the rgnainder plus P and/or T. Best. resulÈ.s wtren nredian of acÈivity residuals is leasts (LlfS). Gives proteclion against max of 1 bad acEiviEies. suggested ouÈliers only serious if associaEed with different conditions or if Lllfs is large (>2?) sqrE Ll'lS = 0.75 giving T = 688oC noE solvable = 4t nps! feaEured ouÈ,Iier ls ksp of the besE MS results (l4S<1.5 LÀfS), average T = 680oC, wlEh scatter 41oC (2 sigrna on disLribution) Single end-mgriber diagrnostic informaÈlon av, sd, fiE are result of doubllng the wrcerlainEy on In a : a ln a is suspecE if arry are v different frqn lsq values. e* a-re In a residuals nor¡nalised Eo 1n a u¡rcertainÈies : large ab;soluEe values, say >2.5, point to suspecE info. hat are Lhe diagonal elqnenEs of the hat rnatrix : large values, say >0.27, poinE Eo influenEial daÈ.a. For 95t confÍdence, fiÈ (= sd(fi.t,)) shoulil be less tlnn 1'73; however a larger value rnay be OK - look aÈ the dÍagnosEics I av sd fit 1sq 683 42 0.66 T sd fir. g* ltat Ìrsp 683 42 0.66 00 q 583 42 0.65 00 ilm 583 42 0.66 00 ru 583 42 0.66 00 H20 683 42 0.66 0.0 0.00 herc 685 42 0. 53 -0.4 0.01 sp 687 46 0 .64 0.2 0.11 si11 683 42 0.65 00 ph1 685 44 0. 55 -0.1 0.05 ann 687 44 o.62 -0.3 0. 08 fcrd 66]- 67 0. 60 -0.3 0. 65 P {.0 4.4 {.8 5.2 5.6 6.0 6.¿ 6.8 7.2 7.6 I .0 rv T 577 599 620 641 662 683 703 72t 744 764 7 8¿ sd 42 42 42 42 42 42 42 42 42 42 42 f 0.6 0.6 0.6 0.6 0.6 o.7 o.7 0.7 0.7 o.7 0. 8 IV-ll Appendix 4, Thermocalc ouþut AVERAGE PRES SURE CALCIJLATIONS Trost Rocks, Ml Assemblage, Sample 964-TR4 an inconqrlete independenE set of reactions has been calculaEed Rock name : TR4M1 (suqgesEed T = 750oC) py alm ilm ru a 0 -0273 0.251 1.00 1. 00 sd (a) /a 0.43961 0.15000 0 0 herc sp cor si11 a 0 .662 0.338 1.00 1.00 sd (a) /a 0.05000 0.I3t47 0 reactions L) 5ilm + 3si11 = alm + 5ru + 2herc 2) 3i1m + 3si11 = alm + 3ru + 2cor calculaEions P (T) sd dP/dT In-K 1 't .6 1 43 -0.0013 -2.207 2 6.6 0 79 0.0053 -]-.382 Rock TR4M1 : average pressures Single end-msdber d.iagrnosEic inf ormation av, sd, fÍÈ are resulE of doublfng Ehe uncertainEy on ln a : a 1n a 1s suspect 1f any a¡e v different frcrn 1sq values. e* are 1n a residuals nor¡nalised to ]n a wrcerEainEies : large absolute values, say >2.5, points Eo sus¡recE info. hat are the diagonal elsnenEs of tåe haE rnatrix ; large values. say >0.33, point to lnfluenEial data. For 95t confidence, fit 1= sd(fiL)) should be less tltan 1.96; however a larger value rnay be OK - look at Ehe diagrnosEics! av sd fit 1sq 6 00 o .82 L.26 P sd fiÈ e* hats alm 5 68 r.28 7.20 -0.2 o.62 Í1n 6 00 0.82 L.26 0 0 Lu 6 00 0.82 t.26 0 0 lrerc 6 29 0.64 0.88 -0.8 0 I7 cor 6 00 0.82 7.26 0 0 sí11 6 00 0.82 7.26 0 0 îoc 550 675 700 725 750 775 000 825 850 tv 5.1 5.3 5.6 5.8 6.0 6.2 6.5 6.7 6.9 sd t.28 I.I7 1. 06 o.94 0.82 0.70 0.68 0.69 0.70 f. 2.1 10 1.7 1.5 1.3 1.1 0.9 0.7 0.5 rv-12 Appendix 4, ThetmocøIc output P-T estimates, Jettv Peninsula AVERAGE PRESSURE / TEMPERATURE ESTIMATES Jetty Peninsula" cordierite-bearing assemblage - 964-59 Àn independenE seb of reactions has been calculaEed for sd(P) li:nit = 2.0 ]dcars : Rock name : PCM59 (for a(CO2) = 0.8 and a(I{2O) = 0.2) sp herc py alm crd fcrd Phl ann a 0.180 0.568 0. 0186 0 .347 0.507 0 .0778 0.0600 0 .0282 sd(In a) 0.201-72 0.05599 0.48626 0.07954 0.07291 0.3L19 0.3420 0.43550 naph ab ksp q si11 tno a 0. 00803 0.104 0. 896 1. 00 1. 00 0 .200 sd(1n a) 7.24533 0.24084 0.OO324 0 0 reactions 1) 2sp + 5q = q¡d 2l py + 2si11 = sp + crd 3) alm + 2si11 = herc + fcrd 4\ 3herc + 5q = alm + 2si11 5) 9fcrd + 10ann = 3herc + 15alm + L0ksp + 10H2O 6) 2'lherc + 9fcrd + 10ph1 - 30sp + 15alm + 10ksp + 10H2o 7l 3sp + 3fcrd+ 2naph+ 2ks¡r = 3py + 2an¡r+ 2ab + 6si11 calculaEions P (T) sd dP/dT lnK I 5.5 0 .98 0.0046 2.750 2 I 0.78 0 .0033 1.591 3 z 0.49 0 .0L04 -2.06I 4 I 0.86 0 .0148 0.538 5 6 0.93 -0 .0030 39.995 6 7 t.42 -0 .0063 -2.030 7 9 7-52 0 .0028 -0.943 Rock PC!459 : average Pt (for a(co2) = 0.8 a¡rd a(IÐo) = o.2\ DiagrnosEic lnforrnaEion avP, an/I, sd's, cor, fiE are result of doubling (first table) / halving (second table) uncertainEy on ln a : a In a is suspec! if any are v different from 1sq values. e* are 1n a residuals norrnalised to ln a w¡certainEies : large absolute values, say >2.5, point to sus¡rects info. hat a¡e the diagonal elsnenEs of Ehe haE maErix : large values, say >0.50, point Eo influenÈial da!a. For 95t confidence, fit 1= sd(fiÈ) ) should be less Lltan 1.¿5 avP sd aúI sd cor fit 1sq 5.6 0.5 676 38 0.823 1.09 P sd P) T sd (r) cor fiÈ hat. sp 5.53 0 54 674 39 0.834 1.08 o -29 0 10 herc 5.56 0 48 669 35 0.812 0.93 t 15 0 .07 pv 5.49 0 64 6'12 44 0.873 1.08 0 19 0 .43 aIm 5.66 0 49 678 35 0.817 L.00 -0 64 0 .03 crd 5.5s 0 54 676 38 0.795 1.09 -0 08 0 .08 fcrd 5.44 0 53 673 37 0.811 1.04 -0 68 0 .22 phl 5.53 0 52 666 44 0.'t75 I.O7 -0 37 0 .49 ann 5.79 0 52 696 39 0.851 0.95 1 13 0 25 naph 5.58 0 48 680 35 0.821 0.99 1 t4 0 03 ab 5.57 0 51 677 38 0.822 1.08 -0 22 0 00 ksp 5.s'l 0 52 676 38 0.823 1.09 -0 00 0 00 q 5.57 0 52 676 38 0.823 1.09 0 0 si11 5.57 0 52 676 38 0.823 1.09 0 0 H20 5.57 0 52 676 38 0.823 1. 09 0 0 fr 676oC, sd = 38 Pr 5.5 kbrr¡, sd = 0,5, cor = 0.823, f = 1.09 rv-13 Appendix 4, Thermocalc output AVERAGE PRES SURE ESTIMATES Jetty Peninsula" cordierite-bearing assemblage - opx-bearing pelite, sample 9&-655 an independenE set of reactions has been calculated for sd(P) li¡nit = 4.0 ldcars : Rock name r 655 (suggested T = 750oC) (for x(co2) = 0.5 and x(IÐo) - 0.5) an ph1 ann easE gr py alm en a 0.411 0.246 0.oo374 0.0379 l-.10e-4 o.0279 0.208 o.372 sd(ln a) 0.10408 0.1-5242 0.64228 0.3891-2 0.81582 0.43686 0.1,4945 0.11832 fs mgbs crd fcrd q iIm ru r120 a 0.0998 0.0343 0. 609 0. 0288 1.00 1.00 1.00 0.500 sd(ln a) 0.28079 0.279'7't 0. 04586 0.4L363 0 0 U reacEions 1) gr + 3mgts + 3q = 3an + py 2\ 3an + en = çtr + 2mgts + 3q 3) 6en + 3fcrd = 4py + 2alm + 9q 41 6fs + 3crd = 2py + 4alm + 9q 5) 2alm + 4nr = fcrd + q + 4i1m 6) 6a¡r + 3fs + 6i1n = 2gr + 4alm + 6m 7l ann + crd = east + alm + 3q 8) 3east + 3alm + 9q = 2ph1 + ann + 3fcrd calculations P (T) sd dPldT ln_K L 7.0 2.57 -0. 0016 12.986 z 5.6 L.7 4 -0.0004 -t2.204 3 5.3 L.02 0. 0031 -0.882 4 6.3 0.87 0.0 067 L.876 5 6.9 0 .57 0.0 086 -0.407 6 6.2 1.05 0.0 ]^05 -72.262 7 6.7 L.23 0.0 100 L.242 8 5.6 1.04 0.0 084 -4.507 Rock 655 : average pressures (for x(CO2) = 0.5 a¡rd x(H2O) = 0.5) Diagmos Eic inf ormation av, sd, fiE are resulE of doublfng (firsE table) / halving (second table) uncertainEy on ln a : a ln a is sus¡rect if any are v differents from Isq values. e* are In a residuals norrnalised t'o 1n a uncertainties : large absoluÈe values, say >2.5, polnt to suspect info. hat a¡e the diagonal elsnenEs of tl:e hat ¡natrix : large values, say >0,53, point to ÍnfluenÈial daEa. For 95t confÍdence, fit 1= sd(fit)) should be less than L.42 av sd fiU 1sq 6.86 0.41 I.25 P sd fit e* håt an 6.87 0.40 r.23 -0 40 00 ph1 6. 88 0 .37 r.I4 -0 90 00 ann 6.85 0 .39 L.20 1 00 00 east 6.9L 0 .35 1.06 1 80 01 gr 6.89 0 .39 1.20 0 90 01 pv 7.00 0 .37 1.09 t 80 08 alm 6.46 0 .58 1.18 -0 80 60 en 6.84 0 -42 1.24 -0 20 04 fs 6. 86 0 .41, r.25 -0 L0 00 rElts 6.85 0 42 r.24 -0 20 05 crd 6.86 0 42 r.25 0 00 02 fcrd 6.88 0 4L r.24 -0 40 01 g 6.86 0 4T 1,.25 0 0 i1m 6.86 0 4l L.25 0 0 l:11 6.86 0 4T 7.25 0 0 ToC 550 600 650 700 750 800 850 900 950 rvP 5.6 5.9 6.2 6.5 6.9 7.2 7.5 7.4 e.2 sd 0 .47 0.42 0.40 0.39 0.41 0.44 o.49 0.55 0.62 f L.7 1.5 1.3 L.2 t.2 1.3 7-4 1.5 L.6 rv-14 Appendix 4, Thermocalc outPut AVERAGE PRESSURE / TEMPERATURE ESTIMATES Jetty Peninsula, M2 assemblage - Opx-bearing pelite, Sample 9&-655 an independenE set of reacEions has been calculated for sd(P) timit = 4.0 ldcars : Rock name : 655 (for x(CO2) = 0.3 and x(H2O) = 0.3) an ph1 ann east gr PY alm en a 0.411 0.246 0.003?4 0.0379 1.10e-4 0.0279 0.208 0.372 sd(1n a) o.1o4o8 0.t5242 0.64228 0.39692 0.81582 0.43686 0.18818 0.11832 fs mgts crd fcrd q Í120 a 0.0998 0.0343 0.609 0.0288 1.00 0.300 sd(ln a) 0.28079 0.27977 0.04586 0.4L363 0 reactions 1) gr+3mgCs+3q=3an+PY 2l 3an+en=gÍ+2mgts+3q 3) 3an+3fs=gr+2alm+3q 4l 6en + 3fcrd = 4py + 2a1m + 9q s) 6fs+3crd=2F{y+4alm+9q 6) ann+crd=east+a1m+3q 7) 2ph1 + 3fcrd = 2east + 2a1m + ãngEs + 9q calculations P (T) sd dPldT ln-K 1 7.0 2.51 -0 .0 015 ]-2.986 5.6 t.70 -0 .0 004 -L2.204 3 6.2 1.40 0 .0 1_30 -2.674 4 4.9 1.00 0 .0 026 -0.882 5 5.8 0.87 0 .0 062 r.876 6 5.9 r.22 0 .0 094 1.242 7 4.9 0.93 0 .0 033 -2.984 Rock 655 : average E{t (for x(CO2) = 0.3 and x(lnO) = 0.3) Dlagnos tic inf or¡naEion avP, arÆ, sd's, cor, f1È are result of doubling (first table) / halving (second Eable) uncerEainty on In a : a ln a is suspecE lf any are v different. from lsq values. e* a.re 1n a residuals nor¡nalised to Ln a uncertainEles : large absoluÈe values, say >2.5, point t'o suspecE lnfo. hat' are Etre diagonal elgnenEs of the ÌnE rnat,rix : large values, say >0.54. point tso lnfluential data. For 95t confidence, flts (- sd(fit)) should be less tharl 1.{5 avP sd a\/I sd cor flE 1sq 5.6 635 93 0.318 1.51 P sd(P) T sd(T) cor fÍt, e* håt an 5 5'l 0 82 637 930 .3t7 1.49 0.33 0 01 phl 5 65 0 75 646 850 .32'l 1.36 0.97 0 01 ann 5 50 0 79 660 950 .258 1.43 -1.06 0 20 easÈ 5 85 0 70 633 76 0 .304 L.23 -L.8'7 0 07 gr 5 56 0 8L 625 930 .3r4 t.46 -0.86 0 05 pv 5 90 0 8'l 5s6 900 .392 7.41 -I.2t 0 26 a1m 5 44 0 85 662 105 0 .]-07 r.47 0.44 0 23 en 5 54 0 83 636 930 .3L2 1.50 -0.32 0 01 fs 5 54 0 86 639 99 0 .197 1.51 -0.16 0 26 rr4 bs 4 83 0 93 560 100 0 .568 1.31 t.37 0 50 crd 5 .57 0 84 635 940 .330 1.51 -0.02 0 01 fcrd 5 .55 0 84 639 970 .254 1.51 -0.20 0 t6 q 5 .5'7 0 83 635 930 318 1.51- 0 0 fr 635oC, sd = 93 Pr 5.6 tçbrr¡, sd = 0.8, cor = 0.318, f = 1.51 ry-15