Peak metamorphic temperature and thermal history of the Southern Alps (New Zealand) O. Beyssac, S.C. Cox, J. Vry, F. Herman

To cite this version:

O. Beyssac, S.C. Cox, J. Vry, F. Herman. Peak metamorphic temperature and thermal his- tory of the Southern Alps (New Zealand). Tectonophysics, Elsevier, 2016, 676, pp.229-249. ￿10.1016/j.tecto.2015.12.024￿. ￿hal-02133783￿

HAL Id: hal-02133783 https://hal.archives-ouvertes.fr/hal-02133783 Submitted on 19 May 2019

HAL is a multi-disciplinary open access L’archive ouverte pluridisciplinaire HAL, est archive for the deposit and dissemination of sci- destinée au dépôt et à la diffusion de documents entific research documents, whether they are pub- scientifiques de niveau recherche, publiés ou non, lished or not. The documents may come from émanant des établissements d’enseignement et de teaching and research institutions in France or recherche français ou étrangers, des laboratoires abroad, or from public or private research centers. publics ou privés. 1 Peak metamorphic temperature and thermal history of the 2 Southern Alps (New Zealand) 3 4 Beyssac O. (1),*, Cox S.C. (2), Vry J. (3), Herman F. (4) 5 6 (1) Institut de Minéralogie, de Physique des Matériaux, et de Cosmochimie, UMR 7 CNRS 7590, Sorbonne Universités – UPMC, Muséum National d’Histoire 8 Naturelle, IRD, 4 place Jussieu, 75005 Paris, France 9 (2) GNS Science, Private Bag 1930, Dunedin, New Zealand 10 (3) Victoria University of Wellington, P O Box 600, Wellington, New Zealand 11 (4) Institute of Earth Surface Dynamics, University of Lausanne, Switzerland 12 13 * Corresponding author: [email protected] 14 15 Submitted to Tectonophysics. Word count ca 14400 all included, 12 Figures, 2 tables. 16 17

18 Abstract 19 The Southern Alps of New Zealand result from late Cenozoic convergence between 20 the IndoAustralian and Pacific plates, and are one of the most active mountain belts in 21 the world. Metamorphic rocks carrying a polymetamorphic legacy, ranging from low- 22 greenschist to high-grade amphibolites, are exhumed in the hanging wall of the 23 Alpine Fault. On a regional scale, the metamorphic grade has previously been 24 described in terms of metamorphic zones and mineral isograds; application of 25 quantitative petrology being severely limited owing to unfavourable quartzo- 26 feldspathic lithologies. This study quantifies peak metamorphic temperatures (T) in a 27 300 x 20 km area, based on samples forming 13 transects along-strike from Haast in 28 the south to Hokitika in the north, using thermometry based on Raman spectroscopy 29 of carbonaceous material (RSCM). Peak metamorphic T decreases across each 30 transect from ≥ 640°C locally in the direct vicinity of the Alpine Fault to less than 31 330°C at the drainage divide 15-20 km southeast of the fault. Thermal field gradients 32 exhibit a degree of similarity from southernmost to northernmost transects, are greater 33 in low-grade semischist than high-grade schist, are affected by folding or 34 discontinuous juxtaposition of metamorphic zones, and contain limited information on 35 crustal-scale geothermal gradients. Temperatures derived by RSCM thermometry are 36 slightly (≤ 50°C) higher than those derived by traditional quantitative petrology using 37 garnet-biotite thermometry and THERMOCALC modeling. The age of RSCM T 38 appears to be mostly pre-Cenozoic over most of the area except in central Southern 39 Alps (Franz Josef-Fox area), where the amphibolite facies schists have T of likely 40 Cenozoic age. The RSCM T data place some constraints on the mode of exhumation 41 along the Alpine Fault and have implications for models of Southern Alps tectonics. 42 43 Keywords 44 45 Southern Alps; Alpine Fault; RSCM thermometry; Alpine Schist; exhumation

1 46 47 1. Introduction 48 49 The kinematics and thermal structure of orogenic wedges result from the coupling 50 between crustal and surface processes at convergent plate boundaries. Being one of 51 the most active mountain belts in terms of both tectonic and surface processes, the 52 Southern Alps of New Zealand offers a unique tectonophysical laboratory to 53 investigate these interactions. The rocks of this mountain belt were formed by 54 Paleozoic and Mesozoic subduction-accretion processes at the paleo-Pacific margin of 55 Gondwana, split from Gondwana and were thinned during the Late Cretaceous, then 56 rent by dextral strike-slip displacement as the Alpine Fault plate-boundary developed 57 during the Neogene. 58 The Southern Alps, which comprise much of the South Island (Figure 1), 59 began forming during the late Cenozoic as the IndoAustralian-Pacific plate motion 60 became increasingly convergent in the Pliocene-Pleistocene. These mountains form 61 against the Alpine Fault - a transpressive section of the Pacific and IndoAustralian 62 plate boundary (see Cox and Sutherland, 2007 for review). The Pacific Plate presently 63 appears to delaminate (e.g. Molnar et al., 1999) or subduct (e.g. Beaumont et al. 1994) 64 within the orogen, actively exhuming a belt of mid-upper crustal material obliquely 65 on the Alpine Fault, and accreting lower crustal material into a thickened crustal root 66 (e.g. Gerbault et al., 2002). The plate boundary is widely cited as a type-example of 67 deep geological processes and continent-continent collision (e.g. Okaya et al., 2007). 68 Over the past twenty years, there has been considerable scientific effort trying 69 to understand the architecture of the IndoAustralian-Pacific plate convergence in the 70 South Island (e.g. Okaya et al., 2007). Evidence has been gathered on the depth of the 71 crustal root, nature of lithosphere, and geometry of faults (see Okaya et al., 2007). 72 This effort has been complemented by thermochronologic work to decipher the timing 73 and thermal structure associated with mountain building and exhumation (e.g.; Tippett 74 and Kamp, 1993a,b; Batt et al., 2000; Herman et al., 2009). Other studies have noted 75 perturbations of the geotherm, producing high thermal gradients and hot spring 76 activity (e.g. Allis et al., 1979; Koons 1987; Allis and Shi 1995; Sutherland et al., 77 2012; Cox et al., 2015). However, while the general metamorphic structure of the 78 Southern Alps is qualitatively well established, there are very few quantitative 79 constraints on the thermal state and thermal history of the crust. An understanding of

2 80 the thermal history of the orogen is needed to constrain the information low- 81 temperature thermochronometers provide about erosion rates and the stability of 82 landforms, as well as the rheology of rocks, behavior of faults at seismogenic depth 83 (Toy et al., 2010), and ultimately seismic hazard (Sutherland et al., 2007). The lack of 84 thermal state information is largely attributable to the bulk rock compositions (mainly 85 metamorphosed quartzofeldspathic greywacke) that are chemically unfavourable for 86 precise metamorphic petrology, and complicated further by the polymetamorphic and 87 polydeformational history of the rocks and potential overprinting effects of fluid flow 88 (Koons et al., 1998; Vry et al., 2004; Menzies et al., 2014). 89 In this study we introduce thermometry based on Raman spectroscopy of 90 carbonaceous material (RSCM) (Beyssac et al. 2002) that allows the quantitative 91 estimate of peak metamorphic temperature (T) independently from the extent of 92 retrogression and presence of diagnostic mineral assemblages. Owing to widespread 93 presence of carbonaceous material in the local Alpine Schist and greywacke, this 94 technique has enabled the generation of a large dataset covering most of the Alpine 95 Fault hanging wall, both along strike and perpendicular to the fault. We present a 96 dataset of 142 new temperature estimates covering a 300 x 20 km area (Table 1). We 97 have also revisited traditional garnet-biotite thermometry results for some of the same 98 samples used for RSCM thermometry, or collected from nearby locations. We 99 provide those results for comparison, along with a few insights gained through 100 comparison of the observed mineral assemblages with their stability fields in P-T 101 pseudosections calculated using THERMOCALC. We then discuss the age of these 102 temperatures by reviewing existing geochronologic constraints to separate the 103 Mesozoic legacy from the late Cenozoic thermal overprint and the extent to which 104 this varies along the plate boundary. Finally, we highlight some constraints these 105 RSCM temperature distributions place on the style and nature of Southern Alps 106 tectonics.

107

108 2. Geological setting 109 110 2.1. General tectonics of the Southern Alps 111 Figure 1 depicts simplified geological and topographic maps of the South Island. 112 Pacific Plate motion relative to the IndoAustralian Plate is 39.7 ± 0.7 mm/a at 245 ±

3 113 1° in the central South Island (MORVEL model of DeMets et al., 2010). The vector is 114 12° anticlockwise of the Alpine Fault, which strikes 053° and is inferred to dip ~40- 115 60° SE (Norris and Cooper, 2007; Stern et al., 2007), extending downward to depths 116 of 25–30 km based on the presence of amphibolite facies schist exhumed in its 117 hanging wall (Grapes, 1995). The generally accepted crustal model depicts the Alpine 118 Fault shallowing eastward into a lower crustal décollement that delaminates the 119 Pacific Plate (Figure 2, e.g., Wellman, 1979; Norris et al., 1990; Okaya et al., 2007), 120 although there is no conclusive evidence for such a detachment. Thermochronological 121 modeling indicates uplift/cooling must be a two stage process first initiating on a 122 gently rising trajectory beneath the dry pro-side of the mountains then occurring 123 more-rapidly up the Alpine Fault ramp (Herman et al., 2009). While the maximum 124 metamorphic grade of exhumed rocks has been used to infer the approximate depth of 125 the Alpine Fault and Pacific Plate delamination, it is predicated on an assumed 126 geothermal gradient and the assumption that previously-stable metamorphic 127 assemblages were exhumed in the late Cenozoic. Although low-temperature 128 thermochronologic ages are clearly the result of Neogene-Quaternary cooling and 129 exhumation, many of the rocks reached peak metamorphic temperatures during the 130 Mesozoic, so much care is needed when using metamorphic assemblages to constrain 131 the present crustal structure.

132 133 2.2. Tectonostratigraphy and structural framework 134 Rocks of the Pacific Plate, southeast of the Alpine Fault, belong to the Eastern 135 Province and mostly to the Torlesse Composite Terrane/Supergroup (Mortimer et al., 136 2014). They are dominated by compositionally monotonous greywacke sandstone 137 and argillite sequences of the Rakaia Terrane that were deposited in an accretionary 138 prism on the margin of Gondwana during the Permian-Triassic (Mortimer, 2004). 139 Sediments were tectonically stacked and imbricated by accretion, and now locally 140 include some metavolcanic-rich and Cretaceous sequences (Cooper and Ireland, 141 2013). Metavolcanic, chert and micaceous-rich schist sequences have been 142 differentiated locally as the Aspiring lithologic association (Craw, 1984; Nathan et al., 143 2002; Cox and Barrell, 2007; Rattenbury et al., 2010) and may warrant separate 144 terrane status (Cooper and Ireland, 2013). Regional low-grade metamorphism affected 145 much of the Rakaia Terrane during the Jurassic (Mortimer 1993, 2000), possibly

4 146 involving several discrete metamorphic events (Adams 2003; Adams and Maas 2004). 147 Schist fabrics and most metamorphic mineral growth have occurred during Jurassic- 148 Cretaceous metamorphism (Cooper and Ireland, 2013; Mortimer, 2000; Vry et al., 149 2004). 150 The 300 km-long belt of schist adjacent to the Alpine Fault contains multiple 151 generations of metamorphic fabrics, folds, and syn- to post-metamorphic quartz veins 152 (Little et al. 2005; Cox and Barrell 2007). A near-continuous mid-upper crustal 153 section is exposed southeastwards across the Southern Alps (Grapes, 1995; Grapes 154 and Watanabe 1992; Little et al., 2005). Mid-crustal mylonites and amphibolite facies 155 schist adjacent to the fault contain evidence of a late Cenozoic ductile deformation 156 overprint that constructively reinforced and reoriented the pre-existing Mesozoic 157 metamorphic fabrics (Little et al. 2002a, 2007; Norris and Cooper 2003, 2007; Toy et 158 al., 2008). A steeply dipping array of late-stage shears that is present for 20 km in the 159 Franz Josef – Fox area (central Southern Alps) represents an exhumed, fossil, brittle- 160 ductile transition zone (BDTZ; Little et al., 2002b; Wightman and Little, 2007). The 161 zone separates schist from relatively undeformed greywacke and semischist 162 sequences that were metamorphosed during the Mesozoic but suffered only brittle 163 effects during the late Cenozoic IndoAustralian-Pacific plate transpression (Cox et al., 164 1997, 2012). 165 Two regionally extensive foliations are developed in schist beside the Alpine 166 Fault: An early foliation (S1 or S2) that is sub-parallel to remnant bedding and 167 metamorphic isograds, and a steeply dipping crenulation foliation (S3) that is axial 168 planar to folds of the S1/S2 fabric and bedding (Grindley, 1963; Little et al. 2002a). 169 At regional scale, metamorphic mineral zones are slightly oblique to boundaries in the 170 textural development of schistosity and cleavage (Little et al., 2005; Cox and Barrell, 171 2007; Rattenbury et al., 2010). The belt of schist varies in width along the orogen, 172 being narrowest (8 km) in the Franz Josef – Fox area of the central Southern Alps, 173 where tight to isoclinal folding has aligned the early schistosity with 035-045º strike 174 weakly oblique to the 055º Alpine Fault (Little et al., 2005). Elsewhere, the folding is 175 more open, early schistosity strikes at a high-angle (nearly perpendicular) to the 176 Alpine Fault, and the schist belt broadens. Mineral metamorphic grade changes are 177 indicated by the presence or absence of key minerals or mineral assemblages but 178 textural metamorphic zones, a semi-quantitative measure of the degree of cleavage, 179 foliation and metamorphic segregation development, are considerably easier to apply

5 180 and map in the field (Bishop, 1974; Turnbull et al., 2001). This classification is 181 widely used throughout New Zealand, and offers a first-order representation of the 182 increasing deformation gradient on a regional scale towards the Alpine Fault. From 183 SE to NE, following increasing metamorphic grade, there are four main textural 184 zones: uncleaved greywacke (TZ1), cleaved greywacke (TZ2a) to well cleaved semi- 185 schists (TZ2b), foliated schists (TZ3) and well segregated schists (TZ4). 186 187 2.3. Metamorphism and geochronology in the Alpine Schist 188 The general metamorphic pattern is well mapped and qualitatively constrained in the 189 Southern Alps by the presence or absence of key index minerals or mineral 190 assemblages (Nathan et al., 2002; Cox and Barrell, 2007; Rattenbury et al., 2010). 191 There is a general SE-NW increasing metamorphic gradient, ranging from sub- 192 greenschist (prehnite-pumpellyite or pumpellyite-actinolite facies); to chlorite then 193 biotite zones (greenschist facies); and garnet-oligoclase then K-feldspar zones 194 (amphibolite facies). Metamorphic mineral zones have been mapped on the basis of 195 the first appearance of minerals, but many of the ‘isograd’ boundaries between zones 196 represent juxtaposition by brittle faults or ductile shear zones, such that the ‘first 197 appearance’ of an index mineral cannot necessarily be assumed to represent a 198 preserved mineral reaction surface (Craw, 1998). Blocks with distinct metamorphic 199 and structural histories are commonly bound by faults or shear zones. Locally, distinct 200 phases of biotite and garnet crystal growth can be distinguished. Fine-grained biotite 201 and very fine-grained grossular-spessartine garnet are associated with rocks attaining 202 garnet-biotite-albite zone of the greenschist facies during the Jurassic, and are 203 distinguished by the biotite-1 and garnet-1 ‘isograds’ (White, 1996; Mortimer, 2000; 204 Cox and Barrell, 2007; Rattenbury et al., 2010). Whilst common in Otago, in most 205 places in the Southern Alps these minerals were either (i) completely overgrown or 206 consumed by growth of porphyroblastic ‘biotite-2’ and ‘garnet-2’ almandine as 207 metamorphism reached amphibolite facies during the Cretaceous-Cenozoic, or (ii) 208 have been retrogressively replaced by chlorite. As with textural zonation, the Alpine 209 Schist metamorphic mineral zones are slightly oblique to the Alpine Fault and to the 210 main SW-NE structural trend of the mountains. 211 There are few detailed and quantitative petrology studies on Alpine Schist. P– 212 T estimates have been made using garnet–biotite thermometers and barometers that 213 involve partitioning of Ca between garnet and plagioclase (Cooper, 1980; Grapes and

6 214 Watanabe, 1992; Grapes, 1995). Peak metamorphic conditions have been constrained 215 to 600–700 °C and 9.2–10 kbar in the Franz Josef – Fox area (Grapes and Watanabe, 216 1992; Grapes, 1995). Such high temperatures were recently confirmed by Toy et al. 217 (2010) applying Ti-in-Biotite thermometry in the mylonitic rocks from the Alpine 218 Fault shear zone. Within a few kilometers of the Alpine Fault, peak P-T conditions 219 decrease progressively to lower P–T conditions, e.g. 400–540 °C and 4–7 kbar 220 towards the southeast (Cooper, 1980; Green, 1982; Grapes, 1995; Grapes and 221 Watanabe, 1992). To the north, in the Hokitika region, Vry et al. (2008) established 222 one of the rare P-T paths available for Alpine Schist, using pseudosection 223 calculations. They obtained a prograde P-T path from ca. 380°C / 2.5 kbar to ca. 224 490°C / 8.5 kbar followed by a slight increase of temperature during decompression 225 to reach ca. 500°C / 6.5 kbar. 226 Geochronological and thermochronological datasets available for the Southern 227 Alps cover a range from low to high closure temperature systems. Batt et al. (2000) 228 generated a compilation of new and previously available fission track data from 229 zircon and/or apatite (see also Tippett and Kamp, 1993a,b, 1995), and K-Ar and 40Ar- 230 39Ar data from muscovite, biotite and/or K-feldspar. This compilation was 231 subsequently complemented in Herman et al. (2009). These studies illustrated 2- 232 dimensional patterns of thermochronological ages with variation both along-strike and 233 perpendicular to the Alpine Fault. Along strike, old apparent ages are almost 234 systematically observed with all techniques close to Haast River in the southwest, 235 representing a complex burial and exhumation history during the Cretaceous and 236 Cenozoic, prior to the development of the modern tectonic regime. Ages decrease 237 towards the northeast and are mostly late Cenozoic and ‘reset’ in Alpine Schist of the 238 central Southern Alps. For instance, Chamberlain et al. (1995) derived 40Ar-39Ar dates 239 for hornblende in the Alpine Schist and obtained reset Neogene (3-6 Ma) ages in the 240 Fox-Franz Josef area, as well as some disturbed 25 Ma ages with plateau ages 241 suggestive of partial gas retention. Although two Cretaceous 40Ar-39Ar ages on 242 hornblende were subsequently obtained by Little et al. (2005) in Fox Valley and 243 Waikukupa River, other hornblende ages and all other thermochronometers for schist 244 between Copland and Whataroa Rivers record a Neogene exhumation history for the 245 central Southern Alps. Further northeast, higher temperature thermochronometers 246 record increasingly older ages, partially reset by the thermal regime of the present 247 plate boundary. Perpendicular to the Alpine Fault, all thermochronometers show a

7 248 zone of young ages adjacent to the fault, which increase southeastwards across the 249 mountains through a series of exhumed partial annealing zones. The pattern of 250 resetting is systematic, with annealing zones located further from the Alpine Fault, but 251 also along strike away from the central Southern Alps, with decreasing closure 252 temperature of the thermochronometer (Batt et al. 2000; Herman et al., 2009; Little et 253 al., 2005). 254 The long and complex burial and exhumation history of Alpine Schist makes it 255 hard to decipher and distinguish the thermal history during late Cenozoic evolution of 256 the plate boundary, and/or formation of the Southern Alps (Mortimer, 2004; Cox and 257 Sutherland, 2007). Schists in the Southern Alps are generally thought to be correlated 258 with those in Otago – both being part of the Haast Schist Group (see Mortimer et al., 259 2014). But while metamorphic assemblages in Otago were formed during Jurassic- 260 Early Cretaceous and schist uplifted in the Early Cretaceous, some metamorphic 261 growth in Alpine Schist occurred during the Late Cretaceous and a component of 262 ductile-brittle deformation overprinted, but did not involve complete recrystallization 263 of these fabrics during the Neogene (Little et al., 2002a,b). A few studies have 264 suggested that locally in the central Southern Alps, peak metamorphic temperatures 265 may have been reached in the Alpine Schist during the Neogene. For instance, on a 266 large scale, by reconsidering available thermochonological data in light of geological 267 and geophysical observations, Little et al. (2005) proposed the local 'Alpine' 268 exhumation of lower crustal rocks (T>500°C) in a narrow zone where 40Ar-39Ar ages 269 on hornblende are totally reset. This narrow zone is located between Franz Josef – 270 Fox and is not located right on the Alpine Fault but a few kilometers from it to the SE. 271 Combining detailed petrological and geochronological investigations on garnet 272 porphyroblasts collected farther to the north, from schist near the Alpine Fault in the 273 Hokitika region, Vry et al. (2004) were able to show that at least locally, peak 274 metamorphic temperatures of ca. 600°C were reached during the late Cenozoic and 275 recorded in the external growing zone of garnets. For this, they dated the different 276 zones of the garnet porphyroblasts using Sm-Nd and Lu-Hf systems and calculated for 277 each zone the P-T conditions of crystallization using garnet-biotite thermometry and 278 garnet-plagioclase-muscovite-biotite barometry. 279 280 3. Methods

8 281 282 3.1. Sampling 283 Samples collected for this study have been contributed to the national New Zealand 284 rock and mineral collection PETLAB (pet.gns.cri.nz). Samples already in the 285 collection, obtained during many years of mapping and research investigations in the 286 western Southern Alps (e.g. Grindley, 1963; Cox and Barrell, 2007; Rattenbury et al., 287 2010), were also analysed. Schist and semischist samples have been systematically 288 characterized for their lithology, main mineral assemblage, textural fabric 289 development, and are georeferenced with standard coordinate systems. Any research 290 action on the samples, such as geochronology, geochemistry or quantitative petrology, 291 has been recorded, and substantial information has been returned to the PETLAB 292 database where it is publically accessible. Supplementary material contains a kmz file 293 of samples, sample descriptions and RSCM T results. Numbers prefixed by P, OU or 294 VU refer to unique sample identifiers used by PETLAB. 295 Samples were collected from the field, or selected from PETLAB, to form 296 thirteen transects across the Southern Alps named according to major rivers or places, 297 to quantify variations both across and along the plate boundary. For each sample 298 location, we calculated a three-dimensional structural distance D (in km) to the Alpine 299 Fault plane by assuming the fault dips at 45° from its mapped surface trace and using 300 the topographic altitude of the sample site, as depicted on Figure 2. Possible variation 301 in D, associated with uncertainty in the subsurface geometry of the fault, was also 302 assessed by varying the dip by ±15°, to 30° and 60° (see below). 303 304 3.2. RSCM thermometry 305 RSCM thermometry is based on the quantitative study of the degree of graphitization 306 of carbonaceous material (CM) that is a reliable indicator of metamorphic temperature 307 (T). Because of the irreversible character of graphitization, CM structure is not 308 sensitive to the retrograde path during exhumation of rocks and records the maximum 309 T reached during metamorphism (Beyssac et al., 2002). Absolute T can be determined 310 in the range 330-650°C with a precision of ± 50 °C due to uncertainties on petrologic 311 data used for the calibration. Relative uncertainties on T are however much smaller, 312 in the range 10-15 °C (Beyssac et al., 2004).

9 313 Raman spectra were obtained using a Renishaw InVIA Reflex 314 microspectrometer (ENS Paris). We used a 514 nm Spectra Physics argon laser in 315 circular polarization. The laser was focused on the sample by a DMLM Leica 316 microscope with a 100 × objective (NA=0.85), and the laser power at the sample 317 surface was set around 1 mW. The Rayleigh diffusion was eliminated by edge filters, 318 and to achieve nearly confocal configuration the entrance slit was closed down to 10- 319 15 µm. The signal was finally dispersed using a 1800 gr/mm grating and analyzed by 320 a Peltier cooled RENCAM CCD detector. Before each session, the spectrometer was 321 calibrated with a silicon standard. Because Raman spectroscopy of CM can be 322 affected by several analytical mismatches, we followed closely the analytical and 323 fitting procedures described by Beyssac et al. (2002, 2003). Measurements were done 324 on polished thin sections cut perpendicularly to the main metamorphic rock fabrics 325 (mostly S1 or S2) and CM was systematically analyzed below a transparent adjacent 326 mineral, generally quartz. More than 15 spectra were recorded for each sample in the 327 static or extended scanning mode (1000-2000 cm-1) with acquisition times varying 328 from 30 (static) to 60 - 150 (extended) seconds. Spectra were then processed using the 329 software Peakfit using a fitting procedure with 3 bands with Voigt profiles (Beyssac 330 et al., 2003; Beyssac and Lazzeri, 2012).

331 332 3.3. Petrology 333 The principal aim of this paper was to investigate metamorphic temperatures using 334 RSCM thermometry on a large scale. However we also felt it is important to place 335 these data in the context of previously available petrology, albeit without providing a 336 complete and parallel petrological assessment. Because published petrological data 337 was at least one, and in most cases three, decades old we undertook some new 338 analytical work and recalculations to update the available data. 339 Major element analyses of bulk-rock samples were carried out by X-ray 340 fluorescence spectrometry, with Fe2+ analyses by titration. Early results had been 341 obtained at Victoria University (Grapes et al., 1982). More recent results were 342 obtained at SpectraChem Analytical Ltd., Wellington, New Zealand, using methods 343 described in Vry et al. (2008), with Fe2+ analyses performed at the Albert-Ludwigs- 344 Universität, Freiburg, Germany, following the techniques outlined in Heinrichs and 345 Herrmann (1990). For these latter samples, the standard CRM2115 was also analysed,

10 346 and the result was Fe2+ = 8.54% compared to the stated value of Fe2+ = 8.50%; the 347 FeO determinations have a 1σ error of 0.08%. The powdered rock samples were 348 prepared at Victoria University using a TEMA tungsten carbide mill. All analytical 349 data are available from the PETLAB database (www.petlab.gns.cri.nz). 350 Electron probe microanalyses of rim compositions of garnet, plagioclase, and 351 coexisting matrix biotite were obtained by R. Grapes, mainly in the early 1990's 352 (Grapes and Watanabe, 1992; Grapes, 1995) using a JEOL JXA-733 SuperProbe, with 353 Moran Scientific software that incorporates modified ZAF and Bence-Albee matrix 354 corrections. The operating conditions used were 15 kV accelerating voltage, 12 nA 355 sample current, and a beam diameter of 1-3 µm. Some representative analyses are 356 given in supplementary material. All analysis positions were located using 357 backscattered electron imaging. Some representative data are given in supplementary 358 material and are available from the PETLAB database (www.petlab.gns.cri.nz). 359 To provide insights into the P-T conditions for the stability fields of relevant 360 mineral assemblages, P-T pseudosections were calculated in the 11 components

361 system MnNCKFMASHTO (MnO, Na2O, CaO, K2O, FeO, MgO, Al2O3, SiO2, H2O,

362 TiO2, O) using THERMOCALC v. 3.35 (Powell & Holland, 1988) and the internally 363 consistent thermodynamic dataset 5.5 (Holland & Powell, 1998). Mineral mixing 364 models and nonideality parameters are based on Holland & Powell (1998) and Powell 365 & Holland (1999). The activity-composition relationships used for garnet 366 (CaMnFMAS), white mica and paragonite (NKFMASH), plagioclase (NCAS), biotite 3+ 367 (KFMASHTO), epidote (CaFe ASH), and chlorite (MnFMASH) are as described in

368 Vry et al. (2008). The activity-composition model for magnetite (FTO) is from White 369 et al. (2000), and the ilmenite (MnFTO) is from White et al. (2005), but with non 370 ideality described by W(ordered ilmenite, pyrophanite) = 2 kJ, W(disordered ilmenite, 371 pyrophanite) = 2 kJ, and W(hematite, pyrophanite) = 25 kJ (R. Powell, personal

372 communication, 2012). Albite, rutile, sphene, quartz, and H2O were treated as pure 373 end-members.

374

375 4. Results

376 377 4.1. Graphitic carbon in the Southern Alps

11 378 Graphitic carbon is widespread in rocks of the Southern Alps, generally dispersed in 379 the mineral matrix and totally absent only at a few localities. The latter localities may 380 correspond to unfavourable lithologies or lithologies affected by intense fluid 381 circulation which may be responsible for bleaching of carbonaceous material in the 382 rocks. In addition, distinctive hydrothermal graphite has been found in association 383 with orogenic gold-quartz mineral deposits in Otago (Pitcairn et al., 2005; Henne and 384 Craw, 2012; Hu et al. 2015). We found no such carbonaceous material in our samples 385 or fieldwork in the central Southern Alps. Figure 3 depicts representative Raman 386 spectra from the Southern Alps and demonstrates the general gradient of 387 graphitization following increasing metamorphism from SE to NW towards the 388 Alpine Fault. In the greywacke east of the drainage divide (Main Divide), 389 carbonaceous material exhibits Raman spectrum with several defect bands (e.g. D1, 390 D2, D3 and D4). Such spectra are characteristic of very disordered graphitic carbon 391 that were transformed under low-grade metamorphism at temperature below 330°C 392 (Beyssac et al., 2002; Lahfid et al., 2010). In this study, the peak temperature has been 393 assumed to be less than 330°C in these samples (Figure 4). To the west, there is a 394 progressive increase of the degree of graphitization through semischist and schist 395 towards the Alpine Fault. Close to the Main Divide, semischist samples exhibit 396 spectra with an intense main defect band as well as a strong D2 defect band as a 397 shoulder on the G peak: such spectra correspond to disordered graphitic carbon in 398 which the tridimensionnal aromatic skeleton remains poorly developed. Going 399 towards the Alpine Fault, both D1 and D2 bands decrease progressively with 400 increasing metamorphic grade and finally completely disappear in some of the highest 401 grade schist samples near the Alpine Fault. This shows a progressive graphitization 402 process that is completely achieved on the Alpine Fault. Importantly, detrital graphitic 403 carbon has been found in samples at all metamorphic grades. It can be easily 404 distinguished from in situ graphitizing organic matter based on: (i) morphological 405 criteria - as it generally appears as isolated grains or flakes; and (ii) Raman spectra - 406 as it usually exhibits a high crystallinity except in very high grade samples where it is 407 difficult to distinguish from organic matter from the simple observation of the Raman 408 spectra. Detrital graphite spectra were not included in RSCM T determinations. Note 409 that the presence of detrital graphite throughout the sequence has been observed in 410 many other metamorphic belts because graphite is easily recycled during the

12 411 erosion/weathering cycle (see Galy et al., 2008), and occurs widely in sedimentary 412 rocks. 413 More specifically, we have also carried out detailed investigation of inclusions 414 of graphitic carbon in the garnet porphyroblasts studied by Vry et al. (2004), sample 415 MA2 (VU37559 in Table 1). Graphitic carbon provides a marker of the garnet zoning 416 as it is present in some zones and absent from others, matching the chemical zonation 417 of garnets. The Raman spectrum of all such inclusions is constant and representative 418 of highly crystalline graphite. Calculating the temperature for such spectra yields ca. 419 575°C (Table 1) in good agreement with the maximum T obtained for the external rim 420 of garnet (ca. 600°C) and late Cenozoic ages. We conclude that all graphitic carbon in 421 this sample recorded the maximum T while garnet composition was only equilibrated 422 in the rims and not in the core during increasing metamorphism. 423 424 4.2. RSCM temperatures in the Southern Alps 425 All RSCM T are listed in Table 1 with key parameters such as location, geological 426 information, the number of spectra per sample, the mean R2 ratio parameter and T 427 with associated uncertainties. For very disordered graphitic carbon that is found in 428 least metamorphosed rocks, we assign T<330°C which is the lower bound of the 429 calibration by Beyssac et al., (2002). RSCM T are plotted on a regional-scale map 430 (Figure 4) and on three local maps depicting the main textural zones, metamorphic 431 mineral zone ‘isograd’ boundaries, faults and folds along the northern (Figure 5, 432 Wanganui to Taramakau), central (Figure 6, Copland to Whataroa) and southern 433 (Figure 7, Haast to Karangarua) segments of the Alpine Fault hanging wall. The latter 434 figures also include T profiles against the structural distance (D) to the Alpine Fault, 435 together with some approximate thermal field gradients drawn manually through the 436 data points. Curve fitting was deemed unwarranted for these transect gradients, due to 437 the relatively small numbers of samples involved in each transect. In addition, some 438 of the RSCM T data are depicted along four geological cross sections (Figure 8) 439 representing variations in geology along the Southern Alps: Waitaha river – Rakaia 440 valley, Whataroa river – Havelock river, Franz Josef – Godley valley, Karangarua - 441 Mt Cook village. Last, all RSCM T were plotted in frequency histograms for the 442 various metamorphic and textural zones (Figure 9). 443

13 444 Based on the dataset, these figures and Table 1, we make the following general 445 observations: 446 - Highest RSCM T measurements were from K-feldspar zone mylonites and schist 447 beside the Alpine Fault, where rocks locally contain only perfect graphite and are 448 inferred to have reached a minimum T of 640°C. Such very high T values are 449 observed in the Haast and Moeraki transects in the south, but also in Copland, 450 Waikukupa and Whataroa transects. There are two samples from garnet-oligoclase 451 zone rocks that contain only highly crystalline graphitic material yielding high T, 452 one from just south of Karangarua and the other at Wanganui river quarry. 453 - Lowest RSCM T measurements, where values have been assigned 330°C to 454 represent the current lower bound of the Beyssac et al. (2002) calibration were 455 observed in the vicinity of the Main Divide or to the southeast. Here the rocks are 456 TZ2a cleaved greywacke or textural zone TZ2b semischist, metamorphosed to 457 either sub-greenschist or chlorite zone. 458 - Plotting all RSCM T data versus metamorphic or textural zones shows the general 459 systematic T increase with increasing metamorphic grade and deformation (Figure 460 9). We note that some metamorphic (e.g. biotite or garnet-oligoclase) or textural 461 (e.g. TZ3 or TZ4) zones are characterized by a relative clustering of the T data, 462 whereas other zones (e.g. chlorite zone or textural zone 2B) exhibit a significantly 463 wider range of RSCM T measurements. This in part reflects the presence of some 464 relatively undeformed TZ2b rocks of the Aspiring lithologic association between 465 Waitaha-Arahura rivers, that are unusual in that they have reached amphibolite 466 facies metamorphism and yet retain remnants of original sedimentary structures 467 (Cox and Barrell, 2007; Cooper and Ireland, 2013). 468 - The RSCM T profiles exhibit a degree of similarity from the southernmost to the 469 northernmost transects. Plots of T as a function of D (Figures 5,6,7) nearly all 470 show RSCM thermal field gradients through the sub-greenschist to greenschist 471 facies (chlorite and biotite zone) rocks that are higher (>35°C/km) than field 472 gradients through amphibolite facies (garnet and K-feldspar zone) (<20°C/km). 473 474 In detail, the RSCM T data along the four geological sections of Figure 8 yields some 475 insight on the thermal evolution of the Alpine Schist. To the south, along the 476 Karangarua – Mt Cook village section (Figure 8d) there is extensive exposure of TZ3 477 and TZ4 rocks, corresponding to biotite and garnet zones, which comprise a map

14 478 thickness of more than 10 km. The rocks exhibit a fan-like structure, marked on a 479 broad scale by the opposite vergence of the Alpine Fault, which dips towards the SE, 480 and faults in the southeast, including the Main Divide Fault Zone (Cox and Findlay, 481 1985), which dip NW. This fan shape structure is mimicked by the main S3 482 crenulation cleavage and schistosity, which dips towards the SE close to the Alpine 483 Fault, is nearly vertical in the garnet zone and then progressively changes to dip 484 towards the NW in the garnet and biotite zones. The enveloping surface of folded, 485 early (S1/2) fabric and lithological variation dips gently southwest. Along this 486 geological section and nearby profiles (e.g. Copland or Karangarua), the field 487 gradients appear systematic with progressive increase in RSCM T towards the Alpine 488 Fault (Figures 6, 7). Yet in detail metamorphic sequences in pumpellyite-actinolite to 489 biotite zone rocks (TZ2a-3) near the Main Divide are locally inverted by juxtaposition 490 on NW-dipping faults (Cox et al., 1997; Craw, 1998) – a brittle juxtaposition which 491 results in steep thermal field gradients. Along the Moeraki profile there are also 492 kilometer-scale late- or post-metamorphic antiform and synform structures plunging 493 10-20° SW that fold the S2 surface (Rattenbury et al., 2010). Here strong T reversals 494 are present in RSCM measurements of garnet zone rocks (Figure 7). Other smaller 495 temperature reversals occur locally in data from the Otoko and Haast profiles, which 496 can also be attributed to late- or post-metamorphic folds (Figure 7; Cooper 1974; 497 Rattenbury et al., 2010). Where such folds are prominent, resulting thermal field 498 gradients are low. The Otoko profile also crosses an area of garnet-biotite-albite zone 499 samples within the greenschist facies, that contain very fine-grained (<1 mm) 500 grossular-spessartine garnets that are thought to be a remanent of Mesozoic 501 metamorphism exposed more widely in Otago (Figure 1; Mortimer, 2000; Rattenbury 502 et al., 2010). There do not appear to be any distinct steps or obvious thermal effects in 503 the dataset associated with this zone, which is distinguished by the ‘garnet-1 isograd’ 504 in Figure 7. 505 In the Franz Josef – Fox area, the high-grade TZ3 and TZ4 schist units 506 metamorphosed to biotite zone and above, are much narrower (8 km, Figure 6), and 507 have been juxtaposed against semischist sequences by the BDTZ with escalator-like 508 component of dip-slip motion (Figures 6, 8c; Little et al., 2002b; Wightman and 509 Little, 2007). Here the distance between the Alpine Fault and the Main Divide is also 510 the smallest, being the region where the late Quaternary dip-slip rates are the greatest 511 on the Alpine Fault (>12 mm/year – Norris and Cooper, 2001), and uplift and

15 512 exhumation rates of schist also appear to be the highest along the plate boundary 513 (Little et al., 2005). The high grade schist units contain an S3 crenulation cleavage 514 dipping towards the SE, weakly oblique to the Alpine Fault, which steepens with 515 distance from the fault (Little et al., 2002a). The geological section across the schist 516 units contains upright SE plunging folds and a tight fan-like structure, with an abrupt 517 truncation of the structural trend towards the NW at the BDTZ. Along this section, 518 high RSCM T (graphitic material at >640°C) was observed in the Alpine Fault 519 mylonite zone, but RSCM T are relatively constant in the range 525-570°C in garnet 520 zone and biotite zone rocks structurally below the BDTZ (see Fox, Waikukupa, Franz 521 profiles on Figure 6). Structurally above the BDTZ, where foliation and bedding have 522 a predominantly NW dip, the thermal field gradients are much higher (>35°C/km, 523 potentially reaching ~90°C/km). Field gradients are also steep through chlorite and 524 biotite zone rocks of the Whataroa profile (Figure 6), which we interpret to be a 525 combination of the effects of topography, flat lying S2 fabric, and fault juxtaposition 526 of different metamorphic blocks (Figure 8b). 527 Features observed in the southern and central profiles are also present in the 528 northern region (Figures 5, 8a,b). The high-grade TZ3 and TZ4 schist units widen to 529 14 km in both map view and geological section, and there are a number of mapped 530 kilometer-scale late- or post-metamorphic antiform and synform structures which 531 appear to fold isotect and isograd boundaries between Wanganui and Hokitika rivers. 532 Such folds are relatively tight with steep limbs in foliated schist sequences, but 533 typically open in the semischists (Figure 8a). The result appears to be a relatively flat- 534 lying S2 enveloping surface, similar to the Whataroa-Havelock section (Figure 8b), 535 plunging gently-moderately to the northeast, which produces temperature reversals in 536 the Waitaha profile (Figure 8a). TZ2a and TZ2b semischist sequences (greenschist 537 facies) are relatively flat-lying, with structural blocks juxtaposed by subhorizontal 538 folding or faults (Andrews et al., 1974). Thermal field gradients are high (>45°C/km) 539 in the sub-greenschist to biotite zone rocks, and very low (<10°C/km) across schist 540 sequences. 541 542 4.3. Petrological constraints 543 Petrological data were collated for samples from the Franz Josef-Fox area, including 544 older studies (Grapes and Watanabe, 1992; Grapes, 1995). Traditional garnet-biotite 545 geothermometry, and temperature estimates based on results of the P-T pseudosection

16 546 calculations new for this study are presented in Table 2. An example of pseudosection 547 calculation is given as supplementary material. These data are shown on Figure 10 548 along with P-T estimates extracted from Figure 4 in Grapes and Watanabe (1992), and 549 shown in map view on Figure 11. Note that garnet-biotite temperatures in our study 550 are generally higher than those from Grapes and Watanabe (1992), based on the same 551 mineral analyses. In our study, we used the calibration by Hodges and Spear (1982) 552 based on rim analyses using pressure estimates based on results of garnet-biotite- 553 muscovite-plagioclase barometry (Hoisch, 1990, Fe-endmember). Grapes and 554 Watanabe (1992) used the calibration by Ferry and Spear (1978), as modified for Ca 555 content by Hodges and Spear (1982) and Hoinkes (1986), and the garnet-biotite- 556 muscovite-plagioclase geobarometer of Ghent and Stout (1981), with modification by 557 Hodges and Crowley (1985). The P and T uncertainties (± 1 kbar and ± 50°C; Grapes 558 and Watanabe, 1992), as estimated from standard deviations, apply for both studies. 559 560 5. Discussion

561 562 5.1. Comparison of RSCM data with petrological constraints 563 At some localities close to the Alpine Fault, in the high-temperature range, there is 564 good agreement between RSCM thermometry and the peak T estimated by petrology. 565 Nearly pure crystalline graphite is present yielding RSCM T above 580°C close to the 566 Alpine Fault, and in many places perfect graphite is found in the central and southern 567 areas yielding RSCM T above 640°C, the higher limit of the calibration by Beyssac et 568 al. (2002). This is in agreement with the T estimates by Vry et al. (2004) who showed 569 that the maximum T recorded by a garnet porphyroblast in Hokitika area (Mac’s 570 Creek) is ca. 600°C using garnet-biotite thermometry. The lowest RSCM T occur 571 from the Main Divide southeastwards, where most samples have been assigned 572 <330°C based on the lower limit of the Beyssac et al. (2002) calibration. There are no 573 quantitative petrological estimates of T for quartzofeldspathic lithologies 574 metamorphosed to such low temperatures. However the observed low temperatures 575 are in agreement with: 1) the observed metamorphism of rocks at prehnite- 576 pumpellyite and pumpellyite-actinolite facies. For instance, prehnite occurring 577 together with pumpellyite typically indicates temperatures in the range 250-300 °C 578 and pressures below about 2 kbar (e.g. Willner et al., 2013); 2) the presence of

17 579 partially annealed fission tracks in zircons, with zircon ages which suggest rocks had 580 exceeded 240°C closure temperatures and have been uplifted from partial annealing 581 zone or deeper (Batt et al., 2000; Cox and Findlay 1995; Tippett and Kamp 1993a,b); 582 3) K-Ar ages that suggest the rocks have partial retention of gas and remained below 583 temperatures of ca. 300°C (Batt et al., 2000). 584 All methods record the decrease in peak temperature away from the Alpine 585 Fault towards the southeast (Figure 11), but RSCM T are generally higher than any 586 related estimate from petrology and yield significantly lower apparent field 587 metamorphic gradients through the high-grade schist. This is most-clearly illustrated 588 on Figure 10 where all RSCM T data for the southern, central and northern profiles 589 are shown together with petrological T estimates from Grapes and Watanabe (1992) 590 and this study (see Table 2) from the Franz Josef – Fox area. The T estimates by 591 Grapes and Watanabe (1992) are definitely lower by several tens of degrees compared 592 to RSCM T (central Southern Alps profiles) except those in the direct vicinity of the 593 Alpine Fault where they converge towards ca. 600°C. THERMOCALC T estimates 594 are higher than those of Grapes and Watanabe (1992) and for the most-part are lower 595 than RSCM T. 596 A possible explanation of the difference between RSCM T and petrological 597 estimates is due the irreversibility of graphitization (Beyssac et al. 2002) such that 598 RSCM records the peak T and is not sensitive to the retrograde/exhumation path of 599 the rocks. By way of contrast, mineral assemblages can re-equilibrate their chemistry 600 during retrogression, modifying and erasing the peak T signal, especially when fluid 601 circulation and deformation are important. In the Alpine Schist, garnet compositions 602 are not simple Fe-Mg end-member mixtures, and can vary considerably from rock to 603 rock, and the biotite can be subject to re-equilibration and regrowth during uplift and 604 cooling. Fluid flow has locally affected some of the rocks, and the effects and timing 605 of this may pass unrecognized, and have not been quantified. The MnO contents in 606 the bulk rock compositions vary, and the first appearance of garnet in P-T 607 pseudosections is very sensitive to this, as well as the choices of activity-composition 608 models. In any case, we note that temperature estimates based on results of P-T 609 pseudosection calculations for rock samples that contain relatively high-grade mineral 610 assemblages (containing ilmenite, oligoclase, ± garnet), are typically only slightly 611 lower than, and within error of, the RSCM T obtained from nearby samples. 612 However, for the samples that contain lower-grade mineral assemblages with sphene,

18 613 there is a larger temperature difference, with higher T estimates being obtained from 614 the RSCM data (see Figure 10). 615 At some localities (see Table 2), the discrepancy between RSCM T and 616 petrological T estimates is, nonetheless, somewhat surprising. RSCM thermometry 617 has now been used in many various geological contexts and generally exhibits good to 618 excellent concordance with conventional petrology, including garnet-biotite 619 thermometry (e.g. see Plunder et al., 2012 and references therein). However, in most 620 available studies, RSCM T was applied to simple thermal histories, i.e. 621 monometamorphic history with one single thermal event, except the notable examples 622 of some internal units of Taiwan (Beyssac et al. 2007). But given the high quality and 623 consistency of the RSCM spectral data in the Alpine Schist, and considering that 624 minerals may not record the true peak metamorphic temperature, we consider that 625 RSCM thermometry yields a first quantification on a large scale of peak metamorphic 626 T. 627 628 5.2. RSCM T pattern in the Southern Alps 629 Figure 10 is a compilation of all local-scale profiles presented on Figures 5, 6 and 7 630 for southern, central and northern Alpine Schist. RSCM T data from all metamorphic 631 zones has also been projected onto an along-strike section parallel to the Alpine Fault 632 (Figure 12). Also shown are the approximate position of key metamorphic mineral 633 ‘isograds’ (first appearance of biotite; garnet and K-feldspar), and the zone where 634 young (<6 Ma) 40Ar-39Ar and K/Ar ages have been obtained in the central Southern 635 Alps. It is useful compare the along-strike projection against Figure 10 that shows the 636 same data plotted perpendicular to the Alpine Fault. From Figures 10 and 12, it seems 637 there may be a decrease in peak T along the Alpine Fault in the K-feldspar zone going 638 towards the northeast, although this trend may in part be apparent due to the absence 639 of samples analysed from the northernmost area. Both the garnet-oligoclase and 640 biotite zones exhibit a relatively clustered pattern of peak T around 560°C and 520°C 641 respectively all along strike. There is dispersion in both cases around the mean values 642 for each zone especially in the central Southern Alps although this may also be due to 643 a denser sampling bias in this area. The chlorite zone all along the Alpine Fault covers 644 an extremely wide range of T from ca. 360°C reaching as high as ca. 550°C. This is a 645 very wide T range for a classical chlorite greenschist facies zone, especially towards 646 the high T. One possibility is that the lithology/bulk chemistry may not have allowed

19 647 growth of index minerals (garnet and/or biotite, see Vry et al., 2008), or the rocks 648 suffered complete retrogression obliterating some high-grade minerals, resulting in 649 the possibility that samples could have been ‘misclassified’ as being apparently of 650 lower metamorphic grade than the temperature they actually experienced. 651 Figure 10 shows that higher RSCM T extends farther from the Alpine Fault 652 along the southern and, to a lesser degree, the northern compared to the central 653 profiles. This is due to the wider map extension of high-grade units (biotite and garnet 654 oligoclase) in the south and north, whereas these units are far less thick and steeper in 655 the central Southern Alps. Although the data are somewhat scattered, it seems that the 656 RSCM T pattern along both the central and northern profiles has nearly constant high 657 T within the first kilometers close to the Alpine Fault (ca. 6 km for the central and ca. 658 8 km for the north) followed by a steep decrease of T in the eastern biotite and 659 chlorite zones. In the south, the T profile is more continuous and progressive, with no 660 break in the slope, and projection toward Alpine Fault suggest rocks here may have 661 reached temperatures ~700°C. This is consistent with widespread occurrences of 662 pegmatite in the Mataketake Range and less commonly elsewhere between Haast and 663 Moeraki rivers, which reflect partial melting of the schist that occurred during the 664 Late Cretaceous (Chamberlain et al., 1995). The calculation of D for these data 665 assumes the Alpine Fault has a constant dip of 45° (Figure 2). Also shown in Figure 666 10 are the range of possible D values and field gradients that could occur if the fault 667 dipped as shallow as 30° or as steeply as 60°. 668 669 5.3. Age of RSCM temperatures in the Southern Alps 670 RSCM thermometry records the peak metamorphic T undergone by carbonaceous 671 material and the host rock during the burial history but it carries no age information 672 by itself. In the case of complex poly-phased metamorphic histories like in the Alpine 673 Schist with at least two major thermal events, either associated with tectonics on the 674 margin of Gondwana during Jurassic – Cretaceous or with evolution of the present 675 plate boundary during late Cenozoic, an important issue remains with regards to the 676 age of the recorded RSCM T. This is a significant question if one wants to use such 677 data to infer and constrain models of the thermal structure of the Southern Alps, or 678 rheology of Alpine Fault rocks at depth (e.g. Toy et al. 2010), as the T measurements 679 will be most directly relevant if they are late Cenozoic in age. Confirming the RSCM 680 T measurements as late Cenozoic is made more difficult by the complete absence of

20 681 Cenozoic-aged metasedimentary units in the Southern Alps that would only contain 682 the expression of late Cenozoic metamorphism. Age of mineral assemblages and of 683 peak P-T conditions recorded in the Alpine Schist have therefore long been discussed 684 in the literature (e.g. Chamberlain et al., 1995; Little et al. 2002a,b, 2005; Vry et al. 685 2004, 2008), but there is no real consensus as to the exact locations where peak P-T 686 conditions recorded by the rocks are late Cenozoic. There are a number of arguments 687 however that supports this to be the case, at least locally, for Alpine Schist between 688 Fox and Franz Josef, and potentially Copland – Whataroa rivers, as we discuss below. 689 On a large scale, the spatial distribution of RSCM T is strongly correlated with 690 the geometry of the main contacts and tectono-stratigraphic structures that were 691 acquired during formation of the Southern Alps. But the classical two-dimensional 692 model of exhumation by Wellman (1979), exhuming deep old rocks up a ramp formed 693 by the Alpine Fault, would yield the same results whether the peak T was late 694 Cenozoic, or a product of older Jurassic-Cretaceous metamorphism with passive 695 exhumation up the Alpine Fault ramp during the late Cenozoic. Instead, 696 thermochronologic data at least record when the rocks have cooled below some 697 threshold temperature (closure temperature), depending on the selected 698 thermochronometer. Except for U-Pb geochronology on zircon yielding 699 crystallization ages, the highest temperature thermochronometer so far applied has 700 been 40Ar-39Ar ages on hornblende. Between Franz Josef and Fox, hornblende from 701 mylonite beside the Alpine Fault has 40Ar-39Ar ages older than Neogene formation of 702 the plate boundary, yet garnet-oligoclase zone Alpine Schist been reset to < 6 Ma 703 when the Southern Alps were uplifted (Figures 4, 11; Chamberlain et al., 1995; Little 704 et al., 2005). The reset ages are distributed through a zone where measured RSCM T 705 range between 530-565°C (Figure 11). These RSCM temperatures are therefore 706 equivalent to, or higher, than the commonly assumed value for the closure 707 temperature of 40Ar-39Ar ages on hornblende which is ca. 500-550°C. At least, in this 708 zone it seems reasonable to assume that RSCM T represents a peak T which was 709 reached during the late Cenozoic. Further support is given by the close 710 correspondence between (i) peak T of ca. 600°C recorded by the most external rim of 711 garnet porphyroblasts dated to Cenozoic ages (Vry et al., 2004) and (ii) RSCM T of 712 ca. 575°C obtained on all graphitic inclusions on the same garnet from the core to the 713 rim.

21 714 To the south or north, however, there are no other places where a late 715 Cenozoic age for thermochronometers or RSCM T can be easily inferred. There are 716 also observations, as outlined above, showing that peak RSCM T values are 717 Mesozoic: (i) 40Ar-39Ar ages on hornblende are Miocene or older to the south, (ii) 718 RSCM T reversals occur on profiles where the regional S1/S2 foliation has been 719 folded and cross-cut by the S3 crenulation cleavage, suggesting the RSCM T data 720 record the peak metamorphism associated with this older fabric, and (iii) that 721 projection of RSCM temperatures on the Haast and Otoko profiles suggest the schists 722 may have reached maximum T ~700°C, where partial melting that produced 723 pegmatites have been dated as late Cretaceous. This does not mean that all RSCM T 724 are Mesozoic, as for instance the coverage for high T dating systems is restricted to 725 the Franz-Josef Fox region. In the south there are also some local complications that 726 are yet to be fully mapped and understood, such as inversion of the metamorphic 727 sequence in the mylonite zone caused by distributed shear on the Alpine Fault 728 (Cooper and Norris, 2011). 729 730 5.4. Some tectonic implications 731 The belt of schist along the Southern Alps has a seemingly continuous westward 732 increase in metamorphic grade toward the Alpine Fault. It has long been considered 733 (e.g. Suggate, 1963) to be due to an increase in depth of exhumation, with some 20-30 734 km of exhumation adjacent to the fault. It is commonly assumed that the metamorphic 735 and textural boundaries in the schist were once sub-horizontal, as observed in the 736 Otago region of southeast New Zealand, and have been rotated during Neogene 737 tectonics. The notion of rock uplift or exhumation and cooling in the Neogene has 738 been corroborated by thermochronologic studies, which demonstrated young ‘reset’ 739 ages adjacent to the Alpine Fault (e.g. Tippett and Kamp 1995; Batt et al. 2000; Little 740 et al. 2005; Herman et al. 2007; 2009). Structural models of the Southern Alps now 741 nearly all infer some form of deformation and rotation of upper crustal material within 742 the Southern Alps, as the Pacific Plate is delaminated and translated Alpine Fault 743 ramp (see Okaya et al., 2007 for a review). Petrological and field observations 744 provide evidence the Alpine Fault hanging wall has been tilted southeastward in the 745 central Southern Alps (Figure 2). Cumulative vertical displacements on an array of 746 fractures in the BDTZ effect a 22 ± 8º of bulk SE tilt (Wightman and Little, 2007; 747 Little et al. 2007). Structurally below the BDTZ, garnet zone rocks also preserve

22 748 microstructural evidence of distributed ductile shear strain (Ɣ = 0.6, down to the east) 749 with sufficient magnitude that could account for ~32° SE tilt of the schist sequence 750 (Holcombe and Little, 2001; Little et al., 2002a). 751 Although seemingly continuous at a regional-scale, in detail the pattern of 752 metamorphic zones represents a complex disrupted metamorphic pile, which involves 753 slices that are variably affected by folding, and transitions (‘isograd’ boundaries) that 754 involve juxtaposition on faults and shear zones. The presence of the BDTZ in the 755 Franz Josef-Fox area, where an escalator-like back shearing process has occurred, is 756 an important observation and reference frame (Wightman and Little, 2007; Little et al. 757 2007). Structurally highest chlorite and biotite zone rocks near the Main Divide, 758 remained above the brittle-ductile transition, so only record a brittle expression of the 759 modern phase of oblique convergence (e.g. Cox and Findlay 1995; Cox et al., 1997). 760 Below the BDTZ, ductile deformation resulted in constructive reinforcement of pre- 761 existing fabrics rather than superposition of a new foliation, and the SE tilt of the rock 762 sequence. Clearly, field metamorphic gradients measured across major boundaries 763 such as the BDTZ will contain little or no information about crustal-scale geothermal 764 gradients, or pressure-temperature relationships during metamorphism. Measured 765 field gradients might also be expected to be very different either side of the BDTZ 766 due to the difference in nature and style of deformation either side of the boundary. 767 We observed high (>35°C/km) RSCM thermal field gradients through the sub- 768 greenschist to greenschist facies (chlorite and biotite zone) rocks and low (<20°C/km) 769 field gradients through amphibolite facies (garnet and K-feldspar zone) (Figures 770 5,6,7,10). 771 The flux of Pacific Plate rock through the deforming zone has been suggested 772 to have two distinct domains/stages, with pure-shear style motion and thickening in 773 the eastern outboard domain, then inclined out-of-plane non-coaxial simple shear in 774 the inboard domain (Little, 2004; Cox et al., 2012). The metamorphic and textural 775 transition in the Franz Josef-Fox area (central Southern Alps) is considerably 776 narrower in map view relative to schist than further to the north and south along the 777 Alpine Fault (Figure 2), potentially the result of differing geometry of the fault at 778 depth (Little et al., 2005). Schist and semi-schist sequences in the central Southern 779 Alps appear to have been thinned relative to the north and south, but although 780 condensed, the westward prograde metamorphic mineral zonation has remained in 781 sequential order. There have been arguments for structural thinning (Grapes and

23 782 Watanabe 1992) and ‘extrusion’ of lower crustal material (Walcott, 1998) associated 783 with uplift of the Alpine Fault hanging wall. On the basis of geobarometry (garnet- 784 biotite-muscovite-plagioclase) and geothermometry (garnet-biotite) of high-grade 785 schists, Grapes and Watanabe (1992) argued the crustal section in the central 786 Southern Alps has been thinned to one third its original thickness. However, because 787 there are significant uncertainties (at least ±1 kbar) in each pressure estimate, the 0.33 788 thinning ratio carries high uncertainty. If tilting of high-grade Alpine Schist involved 789 thinning or shortening perpendicular to the Alpine Fault, we might expect to see it 790 represented in metamorphic peak T field gradients recorded by the RSCM T data. For 791 example, if thinning has been as substantial as the 0.33 ratio suggested by Grapes and 792 Watanabe (1992), the field gradients might be expected to greatly exceed ‘normal’ 793 crustal geothermal gradients of ~20-40°C/km. Instead, the RSCM T field gradients 794 observed for garnet-oligoclase and K-feldspar zone schist in this study were only 795 ~20°C/km perpendicular to the Alpine Fault in the central Southern Alps, and around 796 10°C/km or lower to the north and to the south (Figures 5,6,7). 797 Local temperature reversals in the Otoko and Waitaha profiles clearly reflect 798 upright antiform and synform structures folding the S2 foliation surface and lowering 799 observed field gradients. Here the enveloping surface of the isograds must have a 800 relatively shallow SW dip and is near horizontal when considered perpendicular to the 801 Alpine Fault. Importantly, there is little evidence in the RSCM T data anywhere for 802 substantial thinning of the high-grade (garnet and K-feldspar zone) Alpine Schist 803 sequence perpendicular to the fault, whether the RSCM T data represent Mesozoic or 804 late Cenozoic peak temperatures. A thinning ratio of 0.33 would increase the apparent 805 geothermal gradient by a factor of 3, although the 0.9 thinning ratio proposed in a 3-D 806 kinematic model (Little, 2004) only requires a 1.11 increase, and is potentially 807 supported by RSCM-T. We suggest any Neogene deformation that occurred to the 808 high-grade schist sequence below the brittle-ductile transition is unlikely to have 809 involved more than 0.5 thinning relative to a fixed Alpine Fault reference frame 810 dipping 45°. Deformation by inclined simple shear would meet such criteria. By way 811 of contrast, chlorite and biotite zone semischist sequences show moderate to high (40 812 - 90°C/km) RSCM T field gradients. Here the rocks have remained above the brittle- 813 ductile transition and Neogene deformation was accommodated by oblique dip-slip 814 faulting (backthrusts), imbricated duplex-like stacking and local reversal of 815 metamorphic grade (Cox and Findlay, 1995; Cox et al., 1997; Craw, 1998). The

24 816 RSCM T field gradients in semischist sequences appear boosted by juxtaposition of 817 metamorphic zones and we believe they are unlikely to represent any true geothermal 818 gradients in the crust. 819 Our observations are presented using a calculated structural distance (D) with 820 regard to an assumed Alpine Fault reference frame, fixed at 45° dip. What was 821 perhaps surprising was how low the resulting thermal field gradients were in the high- 822 grade Alpine Schist, particularly that nowhere can they be considered to have 823 exceeded 20°C/km. Had a 60° Alpine Fault dip been selected, it would have resulted 824 in even lower calculated field gradients (to about 10°C/km, depending locally on 825 topography) and at 30° dip the field gradients would still not have exceeded 30°C/km. 826 Since equally low (5-10°C/km) thermal gradients have been independently predicted 827 during evolution of the Alpine Fault mylonite zone (Toy et al., 2010; Cross et al., 828 2015), it encourages us to think the low RSCM T field gradients might actually reflect 829 the thermal state of high-grade Alpine Schist prior to uplift and exhumation. If not 830 real, then such low geothermal gradients near the Alpine Fault can alternatively be 831 explained by vertical thickening (Little, 2004), a crustal drag structure (Little et al., 832 2005) or imbricated reversals associated with distributed oblique-slip (Cooper and 833 Norris, 2011). Perhaps the simplest alternative explanation is that the RSCM T field 834 gradients dominantly represent an oblique slice through the Mesozoic crustal pile, and 835 that the enveloping surface of Mesozoic isograds is much shallower than the Alpine 836 Fault. A corollary is that the degree of rotation of the hanging wall was limited and it 837 has not been completely rotated into parallelism with the Alpine Fault (as suggested 838 by Wellman fig. 4a,c 1979, or Walcott fig 15, 1998). There are various ways this 839 could be achieved. One is that the dip-slip displacement distributed on backthrust 840 structures almost matches that on the Alpine Fault, so that blocks bound by faults and 841 shear zones are only weakly rotated as they are exhumed up the Alpine Fault ramp 842 (Wellman figure 4b, 1979). A corollary is that faults in the Southern Alps hanging 843 wall must be active and have relatively high cumulative slip rates, which is important 844 for seismic hazard assessment (see Wallace et al., 2007; Cox et al., 2012). An 845 alternative hypothesis, that is not easily addressed with RSCM T data and is beyond 846 the scope of our study, is that the Alpine Fault could itself have been rotated, or 847 evolved to a shallower dip during the Neogene and Quaternary (Koons et al., 2003). 848

25 849 6. Conclusions 850 In this study, we present a dataset of peak metamorphic temperatures experienced by 851 Alpine Schist, semischist and greywacke now exhumed in the hangingwall of the 852 Alpine Fault. Carbonaceous material has been analysed in 142 samples, from 13 low- 853 to high-grade transects, in which peak metamorphic temperatures decrease from ca. 854 650-700°C near the Alpine Fault to less than 330°C at the main drainage divide, about 855 15-20 km southeast from the fault. The temperature decrease is relatively uniform in 856 the south, but distinct thermal field gradients are present across the central Southern 857 Alps. This is the first systematic and consistent dataset at the scale of the entire 858 Southern Alps with quantitative values for the peak metamorphic T experienced by 859 various textural and metamorphic zones. RSCM T increase with metamorphic and 860 textural grade, with reversals occurring only locally across folds and any apparent 861 steps where there are faults. Peak temperatures recorded by the RSCM method are 862 generally higher by ≤ 50°C than existing temperature estimates from petrology. 863 Biotite-in, garnet-in and K-feldspar-in first appearance ‘isograds’ occur at different 864 temperatures along the schist belt, which could reflect variable ages of peak 865 metamorphism, or potentially some truncation and juxtaposition of metamorphic 866 zones by faults and shear zones. RSCM T are mostly pre-Cenozoic except in the 867 Franz Josef - Fox area of the central Southern Alps, where these T are likely Cenozoic 868 in age. 869 The RSCM temperatures place limited constraints on thermal conditions 870 experienced within the orogen, with field temperature gradients potentially carrying 871 information on amounts of tilting and structural re-organisation of the Pacific Plate in 872 the Alpine Fault hangingwall, albeit disrupted by fault and shear zone juxtaposition. 873 Plots of RSCM T with respect to structural thickness (D) perpendicular to the Alpine 874 Fault, assuming a 45° dip, yield thermal field gradients that are consistently low, <20 875 °C/km, within the garnet-oligoclase and K-feldspar zones. It suggests these rocks 876 were neither fully rotated, nor structurally thinned, during exhumation. Given the 877 number and consistency of thermochronological data and geological observations 878 available, this dataset can constitute a basis to test thermokinematic and/or 879 thermomechanical models of mountain building processes in the Southern Alps. Such 880 models may have important implications in terms of thermal structure of the crust

26 881 before, during and after orogenic processes as well for our knowledge of crustal 882 rheology. 883

884 Acknowledgments

885 New samples from Westland National Park were collected under Department of 886 Conservation permit WC-22994-GEO, including material collected by Richard 887 Jongens, Mark Rattenbury and Lukas Nibourel. Older samples were sourced from 888 PETLAB collections at GNS Science, Victoria University of Wellington and 889 University of Otago. Rodney Grapes also provided access to archival samples and 890 analytical material. Holly Godfrey and Belinda Smith Lyttle provided technical 891 support. We also wish to thank our colleagues Tim Little, John Townend, and Rupert 892 Sutherland for discussions and helpful comments during the gestation of this work, 893 although not necessarily implying they agree with all of our interpretations and 894 conclusions. Olivier Beyssac acknowledges funding from ANR (GeoCARBONS 895 project), Sorbonne Universités (PERSU program) and CNRS-INSU. Simon Cox was 896 funded under GNS Science’s ‘Impacts of Global Plate Tectonics in and around New 897 Zealand Programme’ (PGST Contract C05X0203). Frederic Herman was funded by 898 the Swiss National Fund (grant PP00P2_138956). We thank Dave Craw and an 899 anonymous reviewer for very constructive help, and Jean-Philippe Avouac for his 900 editorial support. 901

27 902

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32 1118 Norris, R.J., Koons, P.O., Cooper A.F. (1990) The obliquely convergent plate 1119 boundary in the South Island of New Zealand: implications for ancient collision 1120 zones, New Zealand. Journal of Structural Geology, 12, 715-725. 1121 Okaya, D., Stern, T., Davey, F., Henrys, S., Cox, S. C. (2007) Continent-continent 1122 collision at the Pacific/Indo-Australian plate boundary: background, motivation, and 1123 principal results, in A Continental Plate Boundary: Tectonics at South Island, New 1124 Zealand edited by D. Okaya, T. Stern, F. Davey. AGU Geophysical Monograph 1125 Series, 175, 1-18. 1126 Pitcairn, I K, Roberts, S, Teagle, D A H & Craw, D. 2005. Detecting hydrothermal 1127 graphite deposition during metamorphism and gold mineralisation. Journal of the 1128 Geological Society, London 162: 429-432. 1129 Plunder, A., Agard, P., Dubacq, B., Chopin, C., Bellanger, M. (2012) How continuous 1130 and precise is the record of P-T paths? Insights from combined thermobarometry and 1131 thermodynamic modelling into subduction dynamics (Schistes Lustrés, W. Alps). 1132 Journal of Metamorphic Geology, 30, 323–346. 1133 Powell, R., Holland, T.J.B. (1988) An internally consistent dataset with uncertainties 1134 and correlations: 3, Applications to geobarometry, worked examples and a computer 1135 program. Journal of Metamorphic Geology, 6, 173–204. 1136 Powell, R. & Holland, T.J.B. (1999) Relating formulations of the thermodynamics of 1137 mineral solid solutions; activity modeling of pyroxenes, amphiboles, and micas. 1138 American Mineralogist, 84(1–2), 1–14. 1139 Rattenbury, M.S., Jongens, R., Cox S.C. (compilers) (2010) Geology of the Haast 1140 area. Institute of Geological and Nuclear Sciences 1:250,000 geological map 14. 1141 Lower Hutt, New Zealand: Institute of Geological and Nuclear Sciences. 67 pages + 1 1142 folded map. 1143 Stern, T., Okaya, D., Kleffman, S., Scherwath, M., Henrys, S., Davey, F. (2007) 1144 Geophysical exploration and dynamics of the Alpine Fault zone, in A Continental 1145 Plate Boundary: Tectonics at South Island, New Zealand edited by D. Okaya, T. 1146 Stern, F. Davey. AGU Geophysical Monograph Series, 175, 207-233. 1147 Suggate, R.P. (1963) The Alpine Fault. Transactions of the Royal Society of New 1148 Zealand (Geology), 2, 105-129. 1149 Sutherland, R., Eberhart-Phillips, D., Harris, R.A., Stern, T.A., Beavan, R.J., Ellis, 1150 S.M., Henrys, S.A., Cox, S.C., Norris, R.J., Berryman, K.R., Townend, J., Bannister, 1151 S.C., Pettinga, J., Leitner, B., Wallace, L.M., Little, T.A., Cooper, A.F., Yetton, M., 1152 Stirling M.W. (2007) Do great earthquakes occur on the Alpine Fault in central South 1153 Island, New Zealand?, in A Continental Plate Boundary: Tectonics at South Island, 1154 New Zealand edited by D. Okaya, T. Stern, F. Davey. AGU Geophysical Monograph 1155 Series, 175, 235-251. 1156 Sutherland, R., Toy, V.G., Townend, J., Cox, S.C., Eccles, J.D., Faulkner, D.R., Prior, 1157 D.J., Norris. R.J., Mariani, E., Boulton, C., Carpenter, B.M., Menzies, C.D., Little, 1158 T.A., Hastings, M., De Pascale, G.P., Langridge, R.M., Scott, H.R., Lindroos, Z.R., 1159 Fleming, B., Kopf, A.J. (2012) Drilling reveals fluid control on architecture and 1160 rupture of the Alpine Fault, New Zealand. Geology, 40(12), 1143-1146; doi: 1161 1110.1130/G33614.33611.

33 1162 Tippett, J.M., Kamp, P.J.J. (1993a) The role of faulting in rock uplift in the Southern 1163 Alps. New Zealand Journal of Geology and Geophysics 36(4), 497−504. 1164 Tippett, J.M., Kamp, P.J.J. (1993b) Fission track analysis of the late Cenozoic vertical 1165 kinematics of continental Pacific crust, South Island, New Zealand. Journal of 1166 Geophysical Research, 98(B9), 16119−16148. 1167 Tippett, M.J., Kamp, P.J.J. (1995) Quantitative relationships between uplift and relief 1168 parameters for the Southern Alps, New Zealand, as determined by fission track 1169 analysis. Earth Surface Processes and Landforms, 20, 153-176. 1170 Toy, V.G., Prior, D.J., Norris R.J. (2008) Quartz textures in Alpine Fault mylonites: 1171 influence of pre-existing preferred orientations on fabric development during uplift. 1172 Journal of Structural Geology, 30(1), 602-621. 1173 Toy, V.G., Craw, D., Cooper, A.F., Norris, R.J. (2010) Thermal regime in the central 1174 Alpine Fault zone, New Zealand : constraints from microstructures, biotite chemistry 1175 and fluid inclusion data. Tectonophysics, 485(1-4): 178-192; 1176 doi:10.1016/j.tecto.2009.12.013 1177 Turnbull, I.M., Mortimer, N., Craw, D. (2001) Textural zonations in the Haast Schist - 1178 a reappraisal, New Zealand. Journal of Geology and Geophysics, 44(1), 171-183. 1179 Vry, J.K., Baker, J., Maas, R., Little, T.A., Grapes, R., Dixon, M. (2004) Zoned 1180 (Cretaceous and Cenozoic) garnet and the timing of high grade metamorphism, 1181 Southern Alps, New Zealand. Journal of Metamorphic Geology 22(3), 137-157. 1182 Vry, J.K., Powell, R., Williams, J. (2008) Establishing the P-T path for Alpine Schist, 1183 Southern Alps near Hokitika, New Zealand. Journal of Metamorphic Geology, 26(1), 1184 81-97. 1185 Walcott, R.I. (1998) Modes of oblique compression: Late Cenozoic tectonics of the 1186 South Island New Zealand. Reviews in Geophysics 36(1): 1-26. 1187 Wallace, L.M., Beavan, R.J., McCaffrey, R., Berryman, K.R., Denys P. (2007) 1188 Balancing the plate motion budget in the South Island, New Zealand using GPS, 1189 geological and seismological data. Geophysical Journal International, 168(1), 332- 1190 352, doi:10.1111/j.1365-246X.2006.03183.x. 1191 Wellman, H.W. (1979), An uplift map for the South Island of New Zealand and a 1192 model for uplift of the Southern Alps, in The Origin of the Southern Alps, edited by 1193 R. I. Walcott, and M. M. Cresswell, Royal Society New Zealand Bulletin, 1, 13−20. 1194 White, R.W., Powell, R., Holland, T.J.B., Worley, B.A. (2000) The effect of TiO2 1195 and Fe2O3 on metapelitic assemblages at greenschist and amphibolite facies 1196 conditions: mineral equilibria calculations in the system K2O-FeO-MgO-Al2O3- 1197 SiO2-H2O-TiO2-Fe2O3. Journal of Metamorphic Geology, 18, 497-511. 1198 White, R.W., Pomroy, N.E., Powell, R. (2005) An in-situ metatexite-diatexite 1199 transition in upper amphibolite facies rocks from Broken Hill, Australia. Journal of 1200 Metamorphic Geology, 23, 579-602. 1201 White, S. (1996) Composition and zoning of garnet and plagioclase in Haast Schist, 1202 northwest Otago, New Zealand: implications for progressive regional metamorphism. 1203 New Zealand Journal of Geology and Geophysics, 39(4), 515−531.

34 1204 Willner, A.P., Massonne, H.-J., Barr, S.M., White, C.E. (2013) Very low- to low- 1205 grade metamorphic processes related to the collisional assembly of Avalonia in SE 1206 Cape Breton Island (Nova Scotia, Canada). Journal of Petrology 54(9) 1849-1874. 1207 Wightman, R., Little, T.A. (2007), Deformation of the Pacific Plate above the Alpine 1208 Fault ramp and its relationship to expulsion of metamorphic fluids: An array of 1209 backshears, in A Continental Plate Boundary: Tectonics at South Island, New Zealand 1210 edited by D. Okaya, T. Stern, F. Davey. AGU Geophysical Monograph Series, 175, 1211 177-205. 1212

35 1213 1214 Table captions 1215 1216 Table 1 – RSCM temperature data obtained along various profiles in the Southern 1217 Alps. For each sample, information provided are: sample name (PETLAB database, 1218 http://pet.gns.cri.nz), Easting and Northing (New Zealand Transverse Mercator using 1219 NZGD2000), altitude z (in meters), Distance D to the Alpine Fault (in kilometers), 1220 Textural Zone (TZ), Metamorphic Zone, N number of Raman spectra, R2 ratio and 1221 associated standard deviation SDV, Temperature T and associated standard error SE 1222 (Standard error is the standard deviation divided by √N). See text for further details.

1223 1224 Table 2 – Summary of petrologic data obtained in this study along Franz, Fox and 1225 Waikukupa profiles. For each sample, information provided: sample name (PETLAB 1226 database, http://pet.gns.cri.nz), Easting and Northing (New Zealand Transverse 1227 Mercator using NZGD2000), altitude z (in meters), Distance D to the Alpine Fault (in 1228 kilometers), P-T conditions from Grapes and Watanabe (1992), Garnet-biotite 1229 geothermometry (Hodges and Spear, 1982) results (grt-bio T) based on rim analyses, 1230 using pressure estimates based on results of garnet-biotite-muscovite-plagioclase 1231 barometry (Hoisch, 1990, Fe-endmember) for the same or nearby samples, from 1232 Grapes and Watanabe (1992). X (grs+sps) is (Ca +Mn)/(Fe + Mg + Ca + Mn) in 1233 garnet, values <0.2 are generally more suitable for garnet-biotite thermometry. T TC 1234 is maximum temperature estimate based on observed mineral assemblages and results 1235 of P-T pseudosection calculations (this study), for pressure estimates based on nearby 1236 samples and the calculated mineral assemblage stability field. See the PETLAB 1237 database for other analytical data. 1238

36 1239 1240 1241 Figure captions 1242 1243 Figure 1 – Simplified geological (left) and topographic map of the South Island of 1244 New Zealand. The left map depicts the main textural zones in schists and location for 1245 the other figures. The right map shows the topography of the South Island and the 1246 main kinematic vectors for the Pacific Plate relative to the IndoAustralian Plate (De 1247 Mets et al., 2010), and late Quaternary slip on the Alpine Fault (Norris and Cooper, 1248 2001).

1249 Figure 2 – Simplified geological cross-section across the central Southern Alps (see 1250 Figure 1 for location). Bottom is a general cross-section modified after Cox et al. 1251 (2012). Top is a sketch illustrating the calculation of the distance D for sampling 1252 points to the Alpine Fault at depth (see also Figure 8).

1253 Figure 3 – Representative Raman spectra for rocks from the Southern Alps. For each 1254 spectrum, the calculated RSCM T is given. Note that the spectral window represented 1255 was reduced for the figure. P numbers refer to samples from the New Zealand 1256 National Rock and Mineral Collection (PetLab).

1257 Figure 4 – Simplified geological (left) and metamorphic (right) map of the central 1258 Southern Alps (see Figure 1 for location). Left depicts the main textural zones for the 1259 Alpine Schist with the RSCM T data and main ‘isograd’ boundaries between 1260 metamorphic mineral zones. Right depicts RSCM T data with the main metamorphic 1261 zones based on index mineral assemblages and the chrontours (lines of equal age) for 1262 40Ar-39Ar thermochronology on biotite, muscovite and hornblende (after Little et al., 1263 2005). New Zealand Transverse Mercator projection, with NZGD2000 grid.

1264 Figure 5 – Simplified geological map and RSCM T data for the northern segment 1265 (Wanganui to Taramakau valleys) of the Alpine Fault hanging-wall (see Figure 1 for 1266 location). Sample locations are color coded by metamorphic zone, annotated with 1267 RSCM T in °C (as provided in Table 1). Selected metamorphic mineral ‘isograd’ zone 1268 boundaries and fold axes are also shown. On the right, RSCM T versus distance D to 1269 the Alpine Fault (see Figure 2) for the main river profiles. Rotated New Zealand 1270 Transverse Mercator projection, with NZGD2000 grid.

37 1271 Figure 6 – Simplified geological map and RSCM T data for the central segment 1272 (Copland to Whataroa valleys) of the Alpine Fault hanging-wall (see Figure 1 for 1273 location, Figures 5 or 7 for a legend). Sample locations are color coded by 1274 metamorphic zone, annotated with RSCM T in °C (as provided in Table 1). Selected 1275 metamorphic mineral ‘isograd’ zone boundaries, fold axes and the exhumed brittle- 1276 ductile transition zone (BDTZ) are also shown. On the right, RSCM T versus distance 1277 D to the Alpine Fault (see Figure 2) are represented for the main river profiles. 1278 Rotated New Zealand Transverse Mercator projection, with NZGD2000 grid.

1279 Figure 7 – Simplified geological map and RSCM T data for the southern segment 1280 (Karangarua to Haast valleys) of the Alpine Fault hanging-wall (see Figure 1 for 1281 location). Sample locations are color coded by metamorphic zone, annotated with 1282 RSCM T in °C (as provided in Table 1). Selected metamorphic mineral ‘isograd’ zone 1283 boundaries and fold axes are also shown. On the right, RSCM T versus distance D to 1284 the Alpine Fault (see Figure 2) are represented for the main river profiles. Rotated 1285 New Zealand Transverse Mercator projection, with NZGD2000 grid.

1286 Figure 8 – Geological cross-sections and RSCM T data from north to douth: for the 1287 profiles Waitaha to Rakaia (AA’), Whataroa to Havelock (BB’), Franz Josef to 1288 Godley (CC’) and Karangarua to Mount Cook village (DD’) (see Figure 1 for 1289 location). On each section, RSCM T data are represented according to metamorphic 1290 zones (see color code in Figures 5-7) with T value annotations in °C. Main structural 1291 features such as faults or fold axes are also depicted. BB’,CC’, DD’ are modified after 1292 Little et al. (2005).

1293 Figure 9 – Histograms of frequency distribution for all RSCM T data as a function of 1294 metamorphic zones (left) and textural zones (right).

1295 Figure 10 – Diagram with all RSCM T data for the three main segments along the 1296 Southern Alps investigated versus the distance D to the Alpine Fault, assuming the 1297 fault dips at 45° (see Figure 2) with bars representing ± 15° uncertainty for the dip. 1298 On this figure, temperature data (central area) from garnet-biotite thermometry of 1299 Grapes and Watanabe (1992), or reanalysed and recalculated from their dataset, and 1300 maximum temperature from THERMOCALC modeling (this study) are also shown 1301 (see Table 2 and supplementary material).

38 1302 Figure 11 – Enlarged map of the Franz Josef-Fox area, central Southern Alps, 1303 providing a comparison between RSCM T data (circles, Table 1), coloured according 1304 to temperature, and equivalent petrologic results (Table 2) derived from garnet-biotite 1305 geothermometry (squares, recalculated for this study using Hodges & Spear 1982) and 1306 THERMOCALC (squares). The zone of reset 40Ar-39Ar hornblende ages (Hnbl <6 1307 Ma) from Little et al. (2005) is also shown. New Zealand Transverse Mercator 1308 projection, with NZGD2000 grid.

1309 Figure 12 – Diagram with all RSCM T data represented for the different profiles 1310 along the strike of the Alpine Fault by metamorphic zones. Main mineral first 1311 appearance ‘isograds’, drawn by hand, do not necessarily represent preserved 1312 metamorphic mineral reaction surfaces (see text for discussion). Zones where 1313 complete resetting is observed for 40Ar-39Ar data on hornblende (yellow) and biotite 1314 (white) are indicated (after Little et al. 2005).

1315

39 Table 1 (1/3)

Sample E (NZTM) N (NZTM) z (m) D (km) TZ Metamorphic zone N R2 SDV T SDV SE

Haast P77792 1309906 5117483 242 18.1 TZ3 chlorite 16 0.55 0.05 398 21 5 P77793 1312947 5124939 81 14.7 TZ4 biotite 26 0.26 0.08 527 35 7 P77797 1312393 5125154 108 14.4 TZ4 biotite 24 0.35 0.07 487 29 6 P77794 1301364 5128179 80 8.4 TZ4 garnet oligoclase 14 0.13 0.05 583 20 5 P77795 1298599 5126195 79 8.6 TZ4 garnet oligoclase 16 0.15 0.05 573 23 6 P77798 1301490 5128191 79 8.4 TZ4 garnet oligoclase 24 0.13 0.07 583 30 6 P77799 1299029 5126727 87 8.4 TZ4 garnet oligoclase 28 0.14 0.07 580 32 6 P77800 1295689 5124480 72 8.7 TZ4 garnet oligoclase 16 0.15 0.04 576 19 5 P77801 1295229 5124479 72 8.4 TZ4 garnet oligoclase 25 0.15 0.06 574 26 5 P77802 1290858 5127238 45 4.9 TZ4 garnet oligoclase 21 0.14 0.05 580 23 5 P77796 1287199 5128876 30 2.4 TZ4 K feldspar T>640

Moeraki P77450 1324408 5134539 193 14.4 TZ2B chlorite 27 0.34 0.07 490 31 6 P77451 1326594 5135106 327 15.1 TZ2B chlorite 12 0.31 0.03 504 16 4 P77758 1321340 5138581 1787 11.9 TZ2B chlorite 13 0.28 0.05 519 23 6 P77440 1318690 5138159 923 10.2 TZ3 biotite 19 0.20 0.07 552 30 7 P77438 1315991 5138968 1184 8.7 TZ3 garnet oligoclase 13 0.12 0.05 587 23 6 P77439 1317878 5138564 878 9.6 TZ3 garnet oligoclase 15 0.14 0.05 578 24 6 P76968 1316152 5139908 961 8.1 TZ4 garnet oligoclase 34 0.13 0.06 583 25 4 P77004 1308795 5143528 219 2.3 TZ4 K feldspar T>640

Otoko P77471 1335909 5143423 1031 14.9 TZ2B chlorite 15 0.61 0.03 372 15 4 P77472 1334212 5144070 815 13.5 TZ2B chlorite 12 0.40 0.03 462 13 4 P77473 1333527 5144828 753 12.7 TZ3 chlorite 16 0.22 0.05 545 23 6 P77476 1331794 5146308 819 11.2 TZ3 garnet oligoclase 26 0.22 0.05 543 22 4 P77480 1327895 5148161 379 7.9 TZ3 garnet oligoclase 21 0.28 0.07 515 29 6 P77482 1325519 5149565 246 5.8 TZ4 garnet oligoclase 15 0.11 0.05 593 24 6 P77478 1329716 5146734 341 9.5 TZ4 garnet oligoclase 18 0.10 0.05 597 24 6

Karangarua P76825 1357320 5150369 2267 19.7 TZ2B lower greenschist 14 0.64 0.04 357 16 4 P76826 1357720 5150269 2040 19.7 TZ2B lower greenschist 15 0.64 0.02 355 10 2 P63409 1346616 5154270 1539 12.7 TZ4 biotite 18 0.26 0.06 527 26 6 P63418 1351818 5156572 1551 13.3 TZ4 biotite 13 0.25 0.04 529 20 6 P63413 1340214 5161773 1928 6.0 TZ4 garnet oligoclase T>640

Copland P70750 1370092 5157385 906 19.8 TZ1 lower greenschist T<330 P70757 1367001 5162257 1231 16.0 TZ1 lower greenschist T<330 P70727 1364421 5162674 1695 15.1 TZ2B lower greenschist T<330 P76830 1357466 5163538 556 10.9 TZ4 biotite 16 0.35 0.06 487 28 7 P77803 1354440 5163478 591 9.7 TZ4 garnet oligoclase 18 0.17 0.05 565 23 5 P77806 1343254 5168737 62 1.8 TZ4 K feldspar T>640 P78955 1343359 5168653 82 1.9 TZ4 K feldspar 10 0.05 0.06 618 26 8

Fox P76824 1372824 5175278 2487 11.6 TZ1 lower greenschist T<330 OU68321 1372624 5176378 2392 10.9 TZ2A lower greenschist 13 0.63 0.04 359 16 4 P76817 1373524 5173878 2889 12.9 TZ2A lower greenschist T<330 P76818 1373824 5174178 3019 13.0 TZ2A lower greenschist T<330 OU68325 1372624 5176378 2392 10.9 TZ2A lower greenschist T<330 OU68282 1372824 5178379 2543 10.0 TZ2A chlorite T<330 OU68284 1368623 5178579 1963 7.5 TZ2B biotite 17 0.36 0.04 480 19 4 OU68312 1365822 5178679 962 5.6 TZ4 biotite 15 0.22 0.05 542 23 6 OU68317 1366222 5179279 1472 5.8 TZ4 biotite 20 0.19 0.07 555 29 6 OU68313 1365522 5179079 833 5.2 TZ4 garnet oligoclase 17 0.29 0.05 510 21 5 OU68314 1366122 5179479 1350 5.5 TZ4 garnet oligoclase 14 0.18 0.05 560 23 6 OU68316 1366122 5179379 1406 5.6 TZ4 garnet oligoclase 12 0.16 0.05 569 21 6 OU68318 1362021 5178679 460 3.6 TZ4 garnet oligoclase 12 0.21 0.05 546 22 6 P76819 1362021 5178679 460 3.6 TZ4 garnet oligoclase 14 0.26 0.05 526 21 6 P76820 1362021 5178679 460 3.6 TZ4 garnet oligoclase 18 0.24 0.06 535 27 6 P77816 1361578 5179921 356 2.8 TZ4 garnet oligoclase 23 0.20 0.06 551 28 6 Table 1 (2/3)

Sample E (NZTM) N (NZTM) z (m) D (km) TZ Metamorphic zone N R2 SDV T SDV SE

Waikukupa P63364 1371230 5182241 2194 6.9 TZ2B biotite 14 0.26 0.05 524 21 6 P63365 1371341 5181877 2192 7.1 TZ2B biotite 18 0.41 0.05 459 23 5 P63367 1371341 5181877 2192 7.1 TZ2B biotite 13 0.39 0.03 467 14 4 P63368 1371397 5181431 2245 7.4 TZ2B biotite 16 0.55 0.04 395 19 5 P63369 1370423 5181580 1991 6.8 TZ2B biotite 20 0.31 0.06 504 27 6 P63370 1370423 5183580 1545 5.3 TZ3 biotite 14 0.25 0.05 529 24 6 P63381 1369923 5183780 1688 5.1 TZ3 biotite 14 0.24 0.05 536 21 6 P63382 1370037 5183217 1738 5.5 TZ3 biotite 19 0.24 0.06 534 26 6 P63384 1369523 5182880 1775 5.5 TZ3 biotite 12 0.27 0.05 519 22 6 P63385 1369523 5182880 1775 5.5 TZ3 biotite 14 0.23 0.04 539 19 5 P63377 1367323 5183781 1714 4.2 TZ3 garnet oligoclase 14 0.17 0.04 565 20 5 P63375 1368523 5183580 1708 4.7 TZ4 garnet oligoclase 16 0.19 0.06 558 25 6 P63374 1368723 5183380 1763 4.9 TZ4 garnet oligoclase 14 0.23 0.05 540 23 6 P63379 1369623 5184080 1902 5.0 TZ4 garnet oligoclase 16 0.26 0.05 525 22 6 OU68345 1363522 5185881 274 0.1 TZ4 K feldspar T>640 P77818 1363485 5185857 290 0.1 TZ4 K feldspar T>640

Franz P76832 1377624 5175378 1815 13.4 TZ1 lower greenschist T<330 P76831 1376124 5179379 2274 10.8 TZ2A lower greenschist 15 0.65 0.01 350 3 1 P63406 1378625 5186281 1193 6.9 TZ2B chlorite 16 0.35 0.05 487 24 6 P76828 1374424 5182780 1798 7.8 TZ2B biotite 23 0.40 0.07 461 33 7 P63403 1376525 5185280 1522 6.8 TZ2B biotite 14 0.33 0.05 496 21 6 OU68292 1373724 5185381 1778 6.1 TZ3 biotite 15 0.26 0.05 527 21 5 P77807 1370977 5186723 247 3.0 TZ4 garnet oligoclase 18 0.23 0.05 540 23 5 P77808 1370831 5187348 239 2.6 TZ4 garnet oligoclase 16 0.25 0.07 530 29 7 P77809 1370859 5187795 221 2.5 TZ4 garnet oligoclase 20 0.23 0.05 540 23 5 P77810 1370129 5190055 625 1.1 TZ4 garnet oligoclase 19 0.16 0.06 571 27 6 P77811 1369400 5187859 1280 2.4 TZ4 garnet oligoclase 13 0.21 0.05 548 23 6 P77812 1369169 5188205 1182 2.1 TZ4 garnet oligoclase 28 0.19 0.08 558 37 7 P77813 1369402 5188415 1014 1.9 TZ4 garnet oligoclase 14 0.18 0.05 559 21 6 P77814 1369612 5188792 897 1.8 TZ4 garnet oligoclase 17 0.18 0.06 560 27 6 P77819 1373643 5192307 259 0.8 TZ4 garnet oligoclase 26 0.18 0.05 563 23 5

Whataroa P63388 1395627 5190281 1050 11.9 TZ2A chlorite 12 0.66 0.02 346 11 3 P63387 1395627 5190281 1050 11.9 TZ2A chlorite 12 0.65 0.05 353 22 6 Wat_B5 1396792 5191461 947 11.7 TZ2B chlorite 15 0.65 0.01 351 4 1 Wat_B6 1392842 5188901 676 11.2 TZ2B chlorite 13 0.58 0.02 382 7 2 Wat_B4 1391168 5193360 238 7.8 TZ3 chlorite 14 0.22 0.03 544 15 4 Wat_B7 1387148 5187922 840 9.7 TZ3 chlorite 14 0.37 0.05 478 23 6 Wat_B8 1388791 5189073 519 9.4 TZ3 chlorite 14 0.28 0.06 518 25 7 P63394 1398328 5197582 244 8.7 TZ4 chlorite 14 0.20 0.04 553 19 5 P63395 1396728 5197382 176 8.3 TZ4 biotite 14 0.19 0.04 555 16 4 P63397 1392928 5196282 159 6.8 TZ4 biotite 13 0.16 0.04 568 19 5 OU68345 1363522 5185881 274 0.0 TZ4 K feldspar T>640

Wanganui P67288 1417830 5210581 301 9.6 TZ3 biotite 19 0.23 0.05 540 21 5 P63659 1411730 5214782 238 4.9 TZ3 garnet oligoclase 13 0.14 0.05 577 23 6 P67289 1418030 5212081 249 8.8 TZ4 biotite 16 0.19 0.05 557 22 6 P67290 1415130 5212681 252 7.3 TZ4 garnet oligoclase 19 0.19 0.05 557 22 5 Table 1 (3/3)

Sample E (NZTM) N (NZTM) z (m) D (km) TZ Metamorphic zone N R2 SDV T SDV SE

Toaroha P70698 1450376 5235386 1095 9.0 TZ2B chlorite 15 0.43 0.04 449 19 5 P70699 1452031 5236876 1702 9.6 TZ2B chlorite 32 0.52 0.05 409 21 4 P70700 1452031 5236876 1702 9.6 TZ2B chlorite 30 0.56 0.04 390 18 3 P70701 1450530 5234877 1219 9.4 TZ2B chlorite 22 0.44 0.06 447 25 5 P70702 1445132 5237877 1688 5.7 TZ2B garnet oligoclase 20 0.20 0.04 551 17 4 P70703 1445132 5237877 1688 5.7 TZ2B garnet oligoclase 19 0.20 0.05 554 23 5 P70694 1448983 5237444 756 7.1 TZ3 biotite 16 0.28 0.04 519 16 4 P70695 1448983 5237444 756 7.1 TZ3 biotite 22 0.25 0.05 530 24 5 P70696 1448983 5237444 756 7.1 TZ3 biotite 12 0.26 0.04 523 17 5 P61286 1439682 5240627 170 0.6 TZ4 K feldspar 29 0.19 0.07 556 32 6

Waitaha P67564 1428030 5215780 1934 11.3 TZ2A lower greenschist 18 0.50 0.06 417 26 6 P67565 1426331 5216680 1663 10.0 TZ2B biotite 13 0.25 0.04 531 18 5 P67566 1424231 5216480 1255 9.3 TZ3 biotite 15 0.19 0.05 557 24 6 P67571 1431231 5219779 2125 10.3 TZ3 biotite 20 0.27 0.06 519 28 6 P67572 1426431 5222380 1117 6.2 TZ3 biotite 19 0.21 0.06 546 27 6 P67574 1424531 5224180 1680 4.7 TZ3 garnet oligoclase 17 0.11 0.06 591 27 7 P67575 1421532 5225681 1275 2.4 TZ3 garnet oligoclase 26 0.18 0.07 559 33 6 P67570 1429731 5222880 1552 7.6 TZ3 garnet oligoclase 24 0.21 0.05 548 23 5 P67569 1431031 5225779 1973 6.5 TZ3 garnet oligoclase 12 0.20 0.05 550 22 6 P67576 1420532 5226581 1496 1.7 TZ3 K feldspar 22 0.12 0.06 587 28 6 P67577 1420432 5226981 1309 1.3 TZ3 K feldspar 17 0.12 0.06 589 27 7 P67294 1417731 5222081 277 2.6 TZ4 garnet oligoclase 22 0.18 0.07 560 30 6

Taramakau-Arahura VU40234 1463178 5257853 1278 3.0 TZ3 garnet oligoclase 17 0.22 0.04 543 19 5 VU40235 1463178 5257853 1278 3.0 TZ3 garnet oligoclase 15 0.18 0.05 563 20 5 VU40239 1463178 5257853 1278 3.0 TZ3 garnet oligoclase 15 0.19 0.04 557 20 5 VU40240 1463178 5257853 1278 3.0 TZ3 garnet oligoclase 15 0.21 0.06 549 25 6 VU40273 1461586 5257639 1305 2.5 TZ4 garnet oligoclase 18 0.18 0.05 559 23 5 VU37559 1471761 5266809 274 0.3 TZ4 garnet oligoclase 14 0.15 0.05 574 22 6 P77820 1472653 5267524 165 0.1 TZ4 garnet oligoclase 16 0.13 0.06 582 25 6

Cook-Godley P70748 1373471 5166526 1078 15.9 TZ1 lower greenschist T<330 P70753 1395761 5172009 945 22.5 TZ1 lower greenschist T<330 P70759 1374724 5169968 1060 14.4 TZ1 lower greenschist T<330 P70760 1381425 5172177 2459 17.2 TZ1 lower greenschist T<330 P70763 1395228 5171182 1428 22.9 TZ1 lower greenschist T<330 P76829 1378424 5171377 1922 15.9 TZ1 lower greenschist T<330 P70731 1372686 5170557 2190 14.0 TZ1 lower greenschist T<330 63370 1383279 5170123 1787 18.7 TZ2A lower greenschist T<330 P70758 1374200 5170262 1483 14.3 TZ2A lower greenschist T<330 P70762 1391483 5170977 2114 21.7 TZ2A lower greenschist T<330

Otago P77790 1299757 5071010 363 41.4 TZ3 chlorite 12 0.50 0.03 421 15 4 P77791 1297425 5072182 344 39.7 TZ4 chlorite 20 0.35 0.06 485 26 6