City University of New York (CUNY) CUNY Academic Works

Dissertations and Theses City College of New York

2019

Pressure, temperature, and stress conditions of Southern Alpine Fault (New Zealand) mylonites

Jake T. Reitman CUNY City College

How does access to this work benefit ou?y Let us know!

More information about this work at: https://academicworks.cuny.edu/cc_etds_theses/767 Discover additional works at: https://academicworks.cuny.edu

This work is made publicly available by the City University of New York (CUNY). Contact: [email protected] Pressure, temperature, and stress conditions of Southern Alpine Fault (New Zealand) mylonites

City College of New York

Department of Earth and Atmospheric Sciences

Master’s Thesis

By Jake Reitman

Advisor: Dr. Steven Kidder

1 Abstract

The Alpine Fault, New Zealand, provides a unique window into the rheologic history of a seismically-active continental plate boundary. Due to lower exhumation rates along southern portions of the fault relative to the well-studied central Alpine Fault, deformation ages vary along strike by ~15 m.y.. Mylonites collected from two southern creeks near Haast, NZ, display a uniform recrystallized grain size of ~11 �m in mylonitic quartz indicating differential stress values in the south of ~100 MPa, roughly twice the stress values observed in the central Alpine

Fault. Constraints from Ti-in-quartz thermobarometry and Raman Spectroscopy of Carbonaceous

Material (RSCM) indicate mylonitization occurred at 390 – 495 °C, somewhat cooler than observed along the central Alpine Fault. Given similar bulk compositions, and no known variation in strain rate along strike, it is likely that these differences in stress and temperature are related to variation in geothermal gradient between the two areas, with the southern Alpine Fault escaping rapid exhumation that advected considerable heat into the upper crust of the central

Alpine Fault over the past ~6 m.y.. The observed magnitudes of peak stress in the two areas are consistent with those predicted by quartz flow laws using previously estimated geothermal gradients for the two areas, and suggest that integrated crustal strength has decreased by a factor of two in the central Alpine Fault during the Neogene.

1 Introduction

Numerous variables are likely to affect the strength of rocks during deformation including mineral composition, orientation, and size; fluid composition and concentration; geometrical arrangement of minerals; and temperature (Passchier and Trouw, 2005). The quantitative influences of many of these variables on rock strength are not well known, and as a result, numerical simulations of crustal deformation generally incorporate equations (“flow

2 laws”) in which temperature is the only variable responsible for changes in rock strength.

Further, in most flow laws, the relationship between temperature and rock strength is derived from experiments carried out exclusively on rocks containing only one mineral such as quartz

(e.g. Gleason and Tullis, 1995). In this study, we test the degree to which these simplifying assumptions are valid by comparing stress and temperature values recorded in polymineralogic rocks under different temperature regimes along New Zealand’s Alpine Fault (Figure 1). Due to significant variation in uplift rates along strike, the Alpine Fault can be thought of as a rare

“natural laboratory.” It comprises similar rocks that have been exhumed under different temperature-pressure conditions, but have similar strain, strain rate, and bulk composition.

This enables a direct comparison between predicted strength profiles from quartz flow laws and observations.

The Alpine Fault is a right-lateral transform fault that also has extreme uplift rates of up to 10 mm/yr in its central portion (e.g. Bull and Cooper, 1986; Wellman, 1979). The central

Alpine Fault exposes rocks that have been deformed in the middle crust as recently as recently as

3 Ma (hornblende and muscovite Ar40-Ar39) (Chamberlain et al., 1995; Little et al., 2005). The exhumational component varies along strike, such that mylonites deformed as long ago as 18 Ma

(hornblende Ar40-Ar39) are exposed to the south (Barth et al., 2013). This time transgression along strike allows for an analysis of various properties of rocks over the course of the evolution of the fault. In this thesis, we estimate the thermal and stress history of some of the older Alpine

Fault mylonites in the area of Haast, New Zealand. This remote area is less studied than the central Alpine fault, and differs in having a significantly lower exhumation rate. The goal of this study is to better constrain the geologic history of this little-studied portion of the Alpine Fault, with an emphasis on thermal and rheological aspects pertaining to the early history of the fault.

3 By comparison with previous analyses in the central area of the fault, these data allow us to assess the rheological and thermal evolution of the fault. Specifically, we test the hypothesis that the slower uplift rates and colder crustal temperatures in the southern portion of the Alpine Fault result in higher differential stresses and a significantly stronger crust than the central Alpine fault.

2 Background

2.1 Geologic Setting

The Alpine Fault is a transpressional fault zone on the west coast of the South Island of

New Zealand where it comprises the Australian-Pacific plate boundary (Figure 1). It has long been thought that the fault originated in the Oligocene-Miocene as a dextral crustal shear zone

(Carter and Norris, 1976; Cooper et al., 1987; Kamp, 1997; Little and Mortimer, 2001;

Sutherland, 1999), however there is some evidence suggesting that the modern structure was predated by an earlier, proto-Alpine Fault (Lamb et al., 2016; Mortimer, 2018). Strike-slip rates are relatively constant along the Alpine Fault except where the fault branches to the north along various active splays (Figure 1). In contrast, modern dip slip rates vary substantially, with rapid exhumation rates of 6-10 mm/yr occurring in the central part of the fault, and slower rates of

<1.5 mm/yr in the southern region (Berryman et al., 1992; Little et al., 2005; Norris and Cooper,

2000). Exhumation rates have also varied with time. In the Paleogene, the Alpine Fault is thought to have been a dextral transform fault, while the Mid-Late Neogene (~6 – 8 Ma) brought upon the convergence of the Pacific plate overriding the Australian plate based off of Euler pole analyses and modeling (Cande and Stock, 2004; Sutherland, 1995; Walcott, 1998). This convergence resulted in a narrow transpressional fault zone, but a more accurate timing of the convergent onset is still extensively debated by many. The Alpine Fault Zone strikes ~055,

4 dipping between 40-60 SE to a depth of ~15 km, and eventually shallows along a decollement at

35 km depth (Stern et al., 2007). Some slip is also likely absorbed at deep levels on a vertical structure continuing into the mantle (Norris and Toy, 2014).

The protolith of the exposed mylonites in the study area within the Mataketake Range, as well as along the central Alpine Fault, is the Alpine Schist, a composite of mainly quartzofeldspathic schist with subordinate mafic schist and metachert. The protolith of the

Alpine Schist was deposited as part of the Mesozoic accretionary wedge (Mortimer et al., 2012) from Carboniferous to Cretaceous time (Ring et al., 2019). Cooper and Ireland (2013) determined age clusters of zircons in the Alpine Schist of the Hokitika River (~200 km Northeast of Haast) to be 108 ± 1.6 Ma and 71.9 ± 1.8 Ma. They interpreted the older age to signify the age of sedimentation and igneous activity while the younger age represents the timing of subsequent metamorphism. Monazite and garnet dating also indicate peak metamorphism occurred in the mid-Cretaceous (Briggs and Cottle, 2018; Little et al., 2007; Mortimer and

Cooper, 2004; Scott, 2015; Vry et al., 2004), potentially due to collision of the Hikurangi Plateau with the New Zealand portion of the Gondwana margin (Vry et al., 2004).

The base of the exhumed rock sequence exposed along the Alpine Fault is a 1.5 km thick ductile mylonite sequence that was deformed during the Neogene in a shear zone representing the down-dip continuation of the Alpine Fault (Norris and Cooper, 2007; Sibson et al., 1981).

Cataclasites and ultramylonites are found within close proximity to the fault, followed by mylonites, and eventually protomylonites (Figure 2). The protomylonites preserve fabric elements such as isoclinal folds from the Mesozoic (Gillam et al., 2013; Norris and Cooper,

2007; Toy et al., 2013; Toy et al., 2012). Due to variable uplift rates, southern Alpine Fault mylonitization predates deformation in the central Alpine Fault by ~15 m.y. with Ar40-Ar39

5 cooling ages around 18 Ma near Martyr River, ~50 km SW of the study area (Barth et al., 2013).

The study area comprises two creeks in the southern Mataketake Range: Robinson Creek, and a creek located near Mosquito Hill informally referred to as “Mosquito Creek” (Cooper and

Norris, 2011; Toy et al., 2011). Mosquito creek hosts an inverted metamorphic field gradient, i.e. an sequence of metamorphic zones with lower grade rocks at deeper levels (Cooper and Norris,

2011). Cooper and Norris (2011) proposed that the inverted gradient formed as a result of structural juxtaposition of variably-metamorphosed rocks exposed along the fault.

2.2 Alpine Fault P – T Time Constraints

The Alpine schist hosts a range of metamorphic zones. K-feldspar and Oligoclase zones are found proximal to the fault, followed by garnet, biotite, and chlorite zones, and eventually pumpellyite-actinolite and prehnite-pumpellyite facies further east (Cooper, 1972; Grapes and

Watanabe, 1992; Mason, 1962; Reed, 1958; Turner, 1933; Figure 1). Barrovian-facies rocks are present proximal to the fault indicating higher geothermal gradients than present in the higher- pressure intermediate facies rocks away from the fault (Cooper and Norris, 2011). The quantitative description of the metamorphism of this southern area and its regional extent along strike of the Alpine Fault are investigated further in this thesis.

Cretaceous ages of garnet growth near the southern Alpine Schist (97.3 ± 0.3 Ma to 75.4

± 1.3 Ma) demonstrate that peak metamorphism occurred prior to the formation of the Alpine

Fault (Briggs and Cottle, 2018; Vry et al., 2004). The timing of cooling from high temperatures, however, is not well constrained. Thermochronologic data along the southern Alpine Fault (Batt and Braun, 1999) suggest substantial cooling over the period of 60 – 20 Ma. Alternatively, a garnet rim age in the Northern Alpine Fault suggests growth of garnet and elevated temperatures

~600 °C during early Alpine Fault deformation (Vry et al., 2004). These differences indicate

6 insufficient data to precisely resolve post-Cretaceous cooling, and may also point to along-strike differences in metamorphism. Distinguishing peak temperatures from the temperatures that existed during Alpine Fault mylonitization is an important challenge of this study, and motivates the use of Ti-in-quartz thermobarometry which can potentially distinguish deformation temperatures (e.g. Kidder et al., 2018; Kohn and Northrup, 2009).

The dip-slip component of Alpine Fault motion in the south is much smaller than it is in the northern Alpine Fault region: 2.25 ± 0.5 mm/year and >12 mm/year, respectively (Norris and

Cooper, 2000), leading to slower southern cooling rates and older rock and mineral cooling ages.

Zircon fission track ages determined by Batt et al. (2000) in our area of study decrease from 7.4

± 1.4 Ma at Haast Pass to 0.2 ± 0.2 Ma in the Copland, Fox Glacier, and Franz Josef Glacier

Valleys. Muscovite (present within schist) Ar40-Ar39 ages are also higher in the Mataketake

Range (Figure 1) with values of 12.8 +/- 0.1 Ma, 13.2 +/- 0.1 Ma, and 15.4 +/- 0.2 Ma compared to ages in the Fox Glacier and Copland Valley ranging from 5.2 +/- 0.1 Ma to 9.24 +/- 0.04 Ma

(Batt et al., 2000). As previously mentioned, Barth et al. (2013) determined a hornblende Ar40-

Ar39 cooling age within a mylonite 1 km northeast of the Martyr River, (~50 km southwest of the study area), of 18 +/- 4 Ma. Hornblende Ar40-Ar39 ages in the central area of are ~3 Ma

(Chamberlain et al., 1995; Little et al., 2005). The rapid uplift rate in the central Alpine Fault is understood to have begun at ~5 Ma, and is likely responsible for not only younger cooling ages, but considerable heat advection as lower crustal rocks are exhumed to the surface (e.g. Koons,

1987). Thus, we interpret that the present-day central Alpine Fault geothermal gradient (Kidder et al., 2018; Toy et al., 2010) varies significantly from its pre-5 Ma condition as well as geotherms further south (near Haast, NZ) that have not undergone such rapid uplift.

7 2.2.1 Central Alpine Fault thermal history

It is likely that the central and southern portions of the Alpine Fault share certain qualities with respect to their metamorphic histories, based on differences in uplift rates between the two areas that were unlikely to have existed prior to the increased convergence ~5 Ma. Due to the exhumational component of the Alpine Fault, it deforms rocks at a wide range of temperatures and pressures. Scott (2015) found temperature and pressure conditions in the oligoclase zone of

Kini Creek (~67 km northeast of Haast) to be ~600 – 700 °C at 10 kbar. In relation, peak temperatures associated with early mylonitization appear to be ~600 °C (Vry et al., 2004) occurring at depths around 35 km (Grapes, 1995; Vry et al., 2004). Rocks in the central region show cooling ages through hornblende and muscovite closure temperatures of ~500 and ~350 °C respectively, to be as young as <3 Ma (muscovite in Ar) (Little et al., 2005). K-Ar and Ar40/Ar39 hornblende analyses a few km from the Alpine Fault show ages of 3-10 Ma including a hornblende reset zone (~4 km from the fault) with ages <6 Ma (Chamberlain et al., 1995; Little et al., 2005).

Several recent studies have suggested that the geothermal gradient in the Central Alpine

Fault is kinked, with a deep crustal gradient of 5-10 °C km-1 shifting to an average upper crustal gradient of ~45-60 °C km-1 (Cross et al., 2015; Kidder et al., 2018; Toy et al., 2010). These kinked geothermal gradients resemble thermal profiles resulting from thermal advection largely outpacing heat diffusion along a rapidly-exhuming fault (Koons, 1987). Measured geothermal gradients in the upper-most crust are highly variable and can be extreme (95 °C km-1 (Shi et al.,

1996), 63 °C km-1 (Sutherland et al., 2012), 125 °C km-1 (Sutherland et al., 2017b)), most likely due to significant heat transfer by topographically-influenced groundwater flow (Sutherland et al., 2017a).

8

2.2.2 Southern Alpine Fault thermal History

Age constraints relevant for the high-grade rocks of the Mataketake range include a garnet Lu-Hf age of 97+/-0.3 (Briggs et al., 2018) near Haast. Briggs and Cottle (2018) define an “unmodified” (unrecrystallized) igneous monazite age range of 79.5 ± 0.4 – 49.8 ± 0.2 Ma indicating initial crystallization timespan from anatectic crustal melt of deeper rocks than are presently exposed (Briggs and Cottle, 2018; Chamberlain et al., 1995). The initiation of this crystallization, hypothesized by Scott (2015), was likely due to the Late Cretaceous crust undergoing dehydration and partial melting near its base. As the fault comprises an uplifting component, retrograde metamorphism occurred cooling the rocks as they are exhumed to the surface. Batt et al. (1999) interpreted Rb-Sr muscovite and Mu-K-feldspar isochron ages of ~60-

70 Ma of pegmatites present in the northern Mataketake range as indicating cooling to temperatures of 500 +/- 50°C. Cooling through temperatures of 300-350° (where quartz ceases to deform plastically) occurred between 5-12 Ma based on a Rb-Sr age on biotite ~12 Ma, and a zircon fission track age of ~4 Ma (Batt and Braun, 1999).

Peak pressure/temperature constraints of the Alpine Schist within the Mataketake Range region are 4.7-7.3 kbar at 625-680 °C within the K-feldspar zone schists (Beyssac et al., 2016;

Grapes, 1995; Wallace, 1974), and 6.1 +/- 2 kbar and 558 +/- 50 °C in lower grade rock in the oligoclase zone schists in the Moeraki River area (Scott, 2015; Figure 1). Beyssac et al. (2016) compared RSCM peak temperatures associated with differing metamorphic zones along strike of the central and southern regions of the Alpine Fault displaying rather similar values.

The peristerite transitional zone (otherwise known as the peristerite gap (Grapes and

Otsuki, 1983)) seen in Figure 2 represents the transition of albite feldspar to oligoclase feldspar.

9 The onset of amphibolite facies metamorphism is formerly defined as the oligoclase isograd, as it records the cessation of albite (Grapes and Watanabe, 1992). Peristerites are typical indicators of garnet zone metamorphic settings as well as epidote-amphibolite facies (Toy, 2007).

Compositions of peristerite-gap feldspars indicate differences between the central and southern

Alpine schists that likely correspond to different paleo-geothermal gradients during peak metamorphism, with higher geotherms in the north (Grapes and Otsuki, 1983).

Shi et al. (1996) determined a modern geothermal gradient in the upper several hundred meters of crust of approximately 37 °C km-1 based off of borehole measurements located near

Haast at a distance of ~6 km from the Alpine Fault, which is considerably lower than the above- mentioned modern geothermal gradients seen at similar distances from the fault in the central

Alpine Fault.

2.3 Background on Ti-in-quartz thermobarometry

Ti-in-quartz (TitaniQ) thermobarometry (Huang and Audétat, 2012; Thomas et al., 2010;

Thomas et al., 2015) links trace concentrations of titanium in quartz to variations in temperature and pressure. If the Ti-activity and geothermal gradient are known, the thermobarometer is theoretically precise to around ± 50 °C. Resulting temperatures and pressures relate to conditions during crystallization that quartz has experienced either from its initial formation or possibly a deformational event involving recrystallization.

While some papers applying Ti-in-quartz to deformed rocks have found results consistent with expectations (Behr and Platt, 2011; Kidder et al., 2013; Kohn and Northrup, 2009) it is now recognized that there are potential issues that interfere with simple applications of the Ti-in- quartz thermobarometer. First among these is the issue of Ti activity, which has a large influence on calculated temperatures and has been interpreted to vary considerably without relation to

10 mineral assemblage. For example, in order to produce temperatures consistent with independent constraints in rutile-bearing rocks, Ti activity in the Hsueshan range, Taiwan, was found to be

1.0 (Kidder et al., 2012). In contrast, rutile-bearing rocks in the central Alpine Fault require a Ti activity an order of magnitude smaller (0.1). As a result, independent constraints on pressure- temperature conditions may be required in order to utilize Ti-in-quartz data to make quantitative pressure-temperature estimates in deformed rocks.

Another complication is the degree of resetting, or lack of resetting of Ti during deformation and recrystallization. The metamorphic characteristics, such as fluid composition, that affect the re-equilibration of Ti-in-quartz are not well known. Based on experimental data from Nachlas et al. (2018), grains deformed to higher strains with little recrystallization show minimal to no re-equilibration, while recrystallized grains, due to grain boundary processes, show equilibrated Ti concentrations reflecting pressure and temperature conditions during deformation. Dynamic recrystallization via grain boundary migration (GBM) tends to reset Ti-in- quartz concentrations more efficiently than subgrain rotation recrystallization (SGR) (Grujic et al., 2011; Haertel et al., 2013; Kidder et al., 2013; Nachlas and Hirth, 2015). During deformation at low temperatures (300-500°C), Ti-in-quartz appears to reset rather sluggishly, such that most analyses are only partially reset from values associated with an older period of Ti stability (e.g.

Bestmann et al., 2016; Kidder et al., 2013; Nevitt et al., 2017).

3 Methods

3.1 TitaniQ Thermometry

Ti-in-quartz analyses were carried out at the California Institute of Technology using a

Cameca 7f secondary ion mass spectrometer (SIMS) with a 16O- primary ion beam using the same procedures as Kidder et al. (2018). A beam current of 10-15 nA was used with a mass

11 revolving power of 4000, and a field aperture of 100 mm. Masses of 27Al, 30Si, 44Ca, 47Ti were analyzed, and in some samples 56Fe. Before each analysis, a 50 x 50 mm area was rastered for

60 s. The effective spot size for all analyses was approximately 10 �m. A regression line based on 12 analyses of National Institute of Standards (NIST) glasses 610 and 612 (434 ± 15 and 44 ±

5 ppm TiO2, respectively; (Jochum et al., 2011)), and constrained through the origin, was used to determine Ti concentrations. Standards data and calibration curve are provided in Kidder et al.

(2018). A correction factor of 0.67 was used to account for matrix effects between quartz and

NIST glass (Behr et al., 2010). A Herkimer “Diamond”, a natural quartz containing 4 – 5 ppb

Ti, was used as a Ti blank (Kidder et al., 2013). Analyses of the Herkimer “Diamond” (Ti blank) suggest of an effective detection limit of 78 ± 27 ppb. No blank correction was made. The

SIMS analysis procedure of this study was based off of Kidder et al. (2013), which successfully generated known Ti concentrations of two low-Ti quartz standards. Following Kidder et al.

(2018), analyses with high Fe values (56Fe/30Si > 0.007) were removed from the dataset due to likely contamination by Ti-oxide. Temperatures were calculated using equations from the

Thomas et al. (2010) and Thomas et al. (2015) calibration of the TitaniQ thermobarometer.

3.2 Quartz Grain Size Estimates

29 thin sections were examined from mainly quartzofeldspathic schists and four metacherts present in Mosquito Creek and Robinson Creek (Table 1), and all of the samples from

Mosquito Creek are logged in the Otago University collection. The samples are representative of the Alpine Fault mylonitic sequence. Grain size estimates for all quartz grains (recrystallized and unrecrystallized grains) and recrystallized grains were determined in all samples. Geometric mean grain sizes and uncertainties were determined using the linear intercept method (e.g. Exner,

1972), under cross-polarized light in thin sections with thickness of ~30 �m. Multiple linear

12 intercepts between 100-300+ grains per sample were counted (Tables 2, 3) in each sample.

Line intercepts were made through quartz rich areas within the samples. For the grain size analysis, any visible sharp change in luminosity was as a grain boundary, thus optically visible subgrains were included in the analysis.

3.3 RSCM

Raman spectra were obtained using a Renishaw inVia microspectrometer equipped with a 514-mm argon laser housed in Tatung University, Taipei, Taiwan. The laser was focused on the sample using a DMLM Leica microscope with a 50× objective. Laser power at the sample surface was set to be no more than 1 mW. The signal was filtered by edge filters and finally dispersed using an 1800 gr/mm grating to be analyzed by a Peltier cooled CCD detector.

Before each session, the spectrometer was calibrated using a silicon standard. Analytical and fitting procedures described by Beyssac et al. (2002) were strictly followed to avoid analytical pitfalls. For each sample, more than 10 spectra were measured in the extended scanning mode

(1000-2000 cm-1) with acquisition times from 10 to 30 s. Spectra were processed using the software Peakfit following Beyssac et al. (2003).

3.4 CL Imaging

Cathodoluminescence (CL) analyses were obtained using the Zeiss variable-pressure field-emission scanning electron microscope (SEM) at the University of Otago. Thin sections were carbon-coated and analyzed using a variable-pressure secondary-electron (VPSE) detector operated at high-vacuum, 30 kV accelerating voltage and 7nA beam current. The detector is sensitive in the range of 300-650 nm. One hyperspectral CL map (MacRae et al., 2013) was generated (sample 78177) using a JEOL 8500F electron microprobe equipped with ocean optics

QEPro spectrometer tuned to collect from 200-960 nm. The sample used for hyperspectral CL

13 mapping was polished with colloidal silica, coated with 15 nm of carbon, and analyzed at 20kV accelerating voltage, 30nA beam current, a dwell time of 40 ms per pixel, and a defocused beam and pixel size of 2 �m. The hyperspectral CL map was displayed and processed using the in- house software, Chimage, (Leeman et al., 2012) and then fitted at each pixel using a least squares method with a set of three Gaussian distributions (Ti4+ peak at 470 nm (Leeman et al., 2012); near infrared peak at 729 nm and the non-bridging oxygen hole center at 646 nm (Kalceff and

Phillips, 1995)).

4 Descriptions:

The mylonitic section of the southern Mataketake range contains diverse microstructures due to different peak metamorphic grades of the protoliths (Norris and Cooper, 2007), variable degrees of mylonitization associated with Alpine Fault deformation, as well as the local effects of deformation associated with seismic activity and pseudotachylite formation (Toy et al., 2011).

Samples from Mosquito Creek and Robinson Creek both span a distance of ~800 meters to the southeast of the Alpine Fault. Based on their microstructural characteristics, samples can be categorized into two groups: 1) Near-fault samples of mainly greenschist facies ultramylonites and mylonites found within ~300 m of the fault; 2) Coarse-grained protomylonite samples located at larger distances from the fault. Note that all sample distances referred to in this study are map view distances unless otherwise stated.

4.1 Near-fault samples

Samples within ~300 meters of the Alpine fault contain a variety of common low- temperature mylonitic microstructures such as S-C fabrics and C’ shear bands (e.g. Trouw et al.,

2010). Foliation planes are deflected about rounded or lens-shaped feldspar and garnet

14 porphyroclasts that range in size from approximately 100 to 600 µm. The porphyroclasts commonly include mica, quartz, and layers of fine-grained opaque material that vary in orientation from grain to grain. Such inclusion trails are interpreted as remnants of an older foliation and are more common here than further north along the Alpine Fault where feldspar also tends to be finer-grained. Feldspars contain minor undulose extinction, but deformation twins, bending of twins, subgrains, and other evidence of crystal plastic deformation are lacking.

Feldspar porphyroclasts are sometimes fractured and deformed by bookshelf faulting. Chlorite is common in some samples and absent in others, and is not well correlated with any particular microstructures such as C’ shear bands (Gillam et al., 2013; Kidder et al., 2018). Two samples

(mylonite sample 78171 and protomylonite sample NZ78) contain arrays of chlorite-bearing quartz veins and microfaults at a high angle to foliation. Quartz in these 100 – 200 µm thick veins has undulose extinction and minor amounts of recrystallization (Figure 3).

Less-deformed samples, at all structural levels, commonly contain quartz grains of size

100-1000 µm. Areas of samples containing larger grains commonly contain evidence of high grain boundary mobility such as island grains, strong crystallographic preferred orientation

(CPO), and foliation-parallel mica inclusions (Jessel, 1987; Figure 4). Very close to the fault

(<30 m), areas of pure quartz are rare due to phase mixing associated with cataclastic and ultramylonite deformation, but where present quartz is generally fully recrystallized with a grain size of 25-50 µm (Figure 5). Quartz grain size (average of all recrystallized and unrecrystallized grains) increases steadily away from the fault reaching a size of ~180 µm at 200 m from the fault

(Figure 5). This increase is due to increasing sizes and concentrations of coarser grains

(generally 100-500 um) away from the fault. Fine recrystallized grains (5-20 um) are found in all near-fault samples. Figure 5c, which shows the average grain size of the smallest line intercept

15 for each sample, indicates that the size of the smallest grains present decreases gradually away from the fault from about 10-20 µm near the fault to 5-8 µm at ~200 m. Quartz grains tend to be elongate in the foliation plane or form oblique S-C fabrics. Axial ratios average from 1:1 in smaller (~10 µm) grains, and increase gradually to an average of 1:2.5 in 1 mm grains. Larger ribbon grains with axial ratios averaging about 1:7 are common and more prominent in Mosquito

Creek. These grains are comprised of subgrains that are similar in size to many of the recrystallized grains forming at high angle grain boundaries and have a size of ~12 µm. All quartz-rich zones contain a very-strong CPO as evidenced by insertion of the analyzer under cross-polarized light (e.g. Figure 4). Undulose extinction occurs in all coarse grains, and varies from sweeping to patchy. Grain boundaries are generally serrated at a fine wavelength similar to the recrystallized grain size. Recrystallization mechanisms associated with the finer-grained (~11 um) quartz in the near-fault samples is dominated by bulging recrystallization (e.g. Passchier and

Trouw, 2005; Stipp et al., 2002), with a minor amount of subgrain rotation recrystallization.

Average and recrystallized grain sizes follow similar trends in both studied creeks (Figure 5).

4.2 Fault-distal samples

Samples found at distances >300 m from the Alpine Fault, also contain a wide variety of microstructures, but are generally coarser grained than the near-fault samples. A notable exception to this is a number of samples collected from the same outcrop as the pseudotachylite described by Toy et al. (2011). These samples, and a few others not associated with the observed pseudotachylites, contain areas of very-fine grain sizes, prominent veining and cataclastic deformation.

The mylonite sequence hosts a variety of grain sizes including unrecrystallized quartz grains, large recrystallized quartz grains, and fine recrystallized quartz grains. The

16 unrecrystallized quartz grains range in size from around 900 – 1300 µm, large recrystallized grains range from ~45 – 95 µm, and the fine recrystallized grains are approximately 8 – 20 µm.

Layers of coarser quartz grains have interlobate or blocky grain boundaries with ~90° angles between grains. Subgrains are common. Amoeboid shaped grains and other high temperature features such as Zener pinning associated with micas in quartz are abundant (Jessel, 1987). These characteristics correspond to dislocation creep regimes 2 and 3 of Hirth and Tullis (1992).

Quartz grain sizes in the fault-distal samples are much less consistent from sample to sample than closer to the fault (Figure 5). Both average quartz grain size, and the size of the smallest fraction of grains vary over more than an order of magnitude (Figure 5).

5 Results

5.1 RSCM Results

Raman Spectroscopy of Carbonaceous Material (RSCM) data yield peak metamorphic temperatures and is sensitive up to temperatures of 641 °C (Beyssac et al., 2016). RSCM temperatures in the study area show the inverted metamorphic gradient in Mosquito Creek increasing from 513 °C near the fault to ≥641 °C as a function of increasing distance from the fault to the southeast (Table 4, Figure 2). In Robinson Creek (Figure 1), an inverted metamorphic gradient is not present in the mylonite zone. Temperature values there range from

495 to 557 °C with the two highest temperature found close to the fault (Figure 1). The temperatures and mineral assemblage of the Robinson creek samples resemble the garnet-grade rocks of Mosquito creek. Although the Robinson Creek mylonite zone was previously mapped as

K-feldspar grade (e.g. Rattenbury et al., 2010), the observed temperatures and a lack of oligoclase or K-feldspar indicate lower-grade metamorphism in this area.

17

5.2 Ti-in-Quartz Patterns

Ti-in-quartz samples (all from Mosquito creek) show an overall pattern of decreasing Ti concentrations towards the Alpine Fault (Figure 6) that mimics the length-scale and trend of the

RSCM data. Within individual samples, there is considerable scatter of data, however some trends are apparent (Figure 7). The two highest grade samples (78177 and 78182) show a decrease in Ti concentrations from values of ~3-5 ppm to values of about 1-3 ppm in smaller grains. The lower-Ti, near-fault samples (78169 and 78171) have Ti values distributed over a range of 0.3 - ~2 ppm (Figure 7), but small grains (<20 µm) in both samples have higher average values in the range of 1 – 2 ppm (e.g. similar to the higher grade samples). Sample 78176 is a cataclastically-deformed sample adjacent to a pseudotachylite outcrop, and does not show a clear pattern in Ti-in-quartz values.

Veins are present in samples 78171 and 78176 that cross-cut foliation at high angles and contain minor amounts of recrystallization (e.g. Figure 3). The veins in sample in 78171, which have the lowest Ti-in-quartz values of any material sampled from Alpine Fault mylonites (Cross et al., 2015; Kidder et al., 2018), are relatively planar, while in 78176 they are somewhat folded and have Ti-in-quartz values overlapping the non-vein material in the sample. Average values Ti concentrations in the various veins and host material are given in Table 5.

5.2.1 Cathodoluminesce

CL intensity in quartz has a direct correlation to Ti concentrations, particularly at blue wavelengths (Wark and Spear, 2005). In some cases, CL images can be viewed as Ti maps

(Bestmann and Pennacchioni, 2015), with Ti concentrations tending to be higher in brighter regions of the CL map. Figure 8 contains a detailed blue wavelength CL map of sample 78177 overlain by Ti values in specific localities. The image further demonstrates the aforementioned

18 relationship between Ti content and CL intensity by showing that brighter regions generally have higher Ti concentrations while the dull edges of the grains have lower concentrations. The full- spectrum CL map of this sample shows uniform illumination and this pattern is not visible.

6 Discussion

6.1 Significance of the Distribution of Metamorphic Zones along the Alpine Fault

Early interpretations of the occurrence of K-feldspar in mylonitic rocks in the central region of the Alpine Fault suggested an association between high-grade metamorphism and deformation (e.g. Grapes and Watanabe, 1994; Grapes, 1995). Cooper and Norris (2011) showed that K-feldspar is absent in the mylonites in the Haast area, and we have shown that the low-grade metamorphic zone extends at least ~8 km northeast along the fault to Robinson Creek.

K-feldspar is present in this region in unmylonitized rock to the east of the fault, including in granitic pegmatites that crystallized between ~80-50 Ma (Briggs and Cottle, 2018). While a comparison of ages of mylonitization and peak metamorphism indicate that K-feldspar in the central Alpine Fault mylonites predates Alpine Fault deformation, it is also evident from the geometric relationship in our map area that the K-feldspar zone metamorphism and mylonitization are unrelated because they do not occur in the same rocks.

K-feldspar metamorphism in the central Alpine fault most-likely results from lateral smearing of previously metamorphosed high-grade metamorphic rocks along the fault (Cooper and Norris, 2011). The thinning and redistribution of metamorphic zones by strain along the fault is identical to the mechanism responsible for the inverted metamorphic sequence in Mosquito

Creek (Cooper and Norris, 2011). Norris and Cooper (2003) used the distribution and thicknesses of pegmatitic rocks in Alpine Fault mylonites as far north as the Waikukupa River

19 (just southwest of the Franz Josef Glacier) to calculate extremely high strain values in Alpine

Fault mylonites and corresponding highly-resolution strain rates. We note that A. Cross (personal communication, 2013) discovered mylonitic pegmatite in float in Harold Creek, i.e. 50 km further north than previously recognized.

6.2 Variations in Peak Temperature

The majority of peak temperatures from RSCM in the two creeks are from garnet-grade rocks with values ranging from 495 °C to 557 °C. In Robinson Creek, these values fall within a

~40° range from 495 °C to 536 °C over much of the mylonite zone, with temperatures of ~555

°C adjacent to the fault. In Mosquito Creek, RSCM temperatures increase steadily over roughly

1 km from values of 513 °C to ≥642 °C. The Mosquito Creek RSCM data thus confirm that the inverted metamorphic field gradient observed by Cooper and Norris (2011) represents variations in peak metamorphic temperature. The apparent inverted gradient calculated from our data of

188 °C km-1 (108°C/.573 km) is on the high end of such gradients observed elsewhere (England and Molnar, 1993), and following Cooper and Norris (2011), we consider it to be structurally assembled. Robinson Creek does not show an inverted metamorphic gradient in the mylonite zone, however, higher-grade K-feldspar zone rocks are mapped to the southeast of the section we studied (Grapes and Watanabe, 1994; Rattenbury et al., 2010). Thus, at larger scales Robinson

Creek most-likely displays inverted metamorphism. Given their proximity, we attribute the differences between the two creeks to local variations in near-surface Alpine Fault kinematics

(e.g. Little et al., 2002) with the increased thickness of garnet-grade rocks in Robinson Creek potentially resulting from structural duplication along previously-mapped step-over faults

(Figure 1).

6.3 Stability of Ti during late deformation

20 Due to the exhumational component of the Alpine Fault, presently-exposed rocks cooled as they were deformed and transported to the surface. Such cooling generally leads to grain size reduction since colder rocks are stronger in the ductile regime (e.g. Kohlstedt et al., 1995), and recrystallized grain size decreases with increasing stress in rocks deforming by dislocation creep

(e.g. Cooper and Norris, 2011; Stipp et al., 2006). As observed elsewhere (Behr and Platt, 2011;

Kidder et al., 2013), the smallest Alpine Fault quartz grains are most likely to record Ti-in-quartz concentrations stable during the latest period of plastic deformation of quartz. The differences in

Ti-in-quartz between small grains and larger, older, grains may thus reveal evidence of evolving pressure-temperature conditions of Alpine Fault deformation. Below, we refer to Ti concentrations in coarse quartz grains as “initial” values, although we recognize that such values could potentially represent partially re-equilibrated values.

The Ti-in-quartz data collected from Mosquito Creek shows a degree of scatter that is typical of Ti-in-quartz data in rocks deformed at low temperature (Kidder et al., 2013; Kidder et al., 2018). Part of this variation may be due to undetected contamination from analysis of micro- impurities. Such analyses would increase nominal Ti-in-quartz values, but would not decrease them. As a result, Figure 7 should be viewed with the perspective that low values are more trustworthy and high-Ti outliers are likely contaminated.

In the high temperature samples, there is a decrease in Ti-in-quartz values in smaller grains to values of 1-3 ppm (Figure 7). The smallest grains in the two lowest temperature samples (78171 and 78169) also have Ti values in the range 1-2 ppm but generally lower Ti values in coarser grains. We hypothesize that these samples experienced an increase in Ti-in- quartz during recrystallization, however this cannot be confirmed without a larger data set. In any case, Ti-in-quartz concentrations of fine-grained quartz (<15 µm) are generally less-

21 dispersed than coarse-grained quartz, with values of 1-2 ppm being most common. The data in sample 78176 are an exception to this pattern, but this pseudotachylite sample was strongly affected by brittle-ductile deformation and the extremely fine-grained quartz (~6 µm) analyzed is considerably smaller than the beam size (10 µm) such that results from these grains are more likely to be contaminated by non-quartz material occurring along grain boundaries. We note, however, that the lowest-Ti values for this fine grained material are ~1 ppm, consistent with the results from the other samples. The limited range of Ti values of 1-2 ppm in fine grained quartz coincides with observed values in ultramylonites in the central Alpine Fault (Kidder et al., 2018) which also show similar scatter that was interpreted as resulting from incomplete re-equilibration during deformation. In summary, despite a fair amount of scatter, the Ti-in-quartz data appears to indicate a convergence from both higher and lower Ti values towards the range 1-2 ppm.

Quartz in slightly-recrystallized cross cutting veins (e.g. Figure 3) has Ti values of 0.19 ppm – 0.46 ppm in sample 78171 and 0.39 ppm – 1.09 ppm in sample 78176 (Table 5), much lower on average than non-vein quartz. Since the formation of these veins occurred after most mylonitization had occurred, we interpret that at the time the veins formed, temperature and pressure had decreased enough to stabilize these lower Ti values. The presence of recrystallized grains indicates that quartz deformation at temperatures of 300 – 350 °C (Dresen et al., 1997;

Dunlap et al., 1997; Stipp et al., 2002; Stöckhert et al., 1999; van Daalen et al., 1999), thus by comparison with values of 1-2 ppm in the mylonites, they can be used to constrain mylonitization temperature and Ti activity (see below).

6.4 Significance of CL zoning

As described above, sample 78177 contains quartz bands with brighter internal CL coinciding with higher Ti concentrations and darker CL edges containing lower Ti

22 concentrations (Figure 8). The relatively smooth variation resembles zoning due to diffusion, which is fairly common in Ti-in-quartz (e.g. Bestmann and Pennacchioni, 2015). The diffuse zoning does not appear to relate to the position of quartz grain boundaries, so we interpret that the zoning likely developed as a result of lattice diffusion of Ti (rather than diffusion localized along grain boundaries). The equation of Cherniak et al. (2007) can thus be used to roughly constrain a time period associated with this gradient in Ti given that diffusion occurred over a lengthscale of ~80 �m (Figures 8, 9, 10). This timescale can then be used to determine whether diffusion was associated with Alpine Fault deformation or cooling from a metamorphic peak in the late Cretaceous.

The profiles in Figures 8, 9, and 10 suggest diffusive loss of Ti from concentrations of ~3 to 2 ppm, i.e. Ti diffusion likely occurred at a temperature somewhat less than the peak temperature of 621 °C recorded by RSCM in this sample. If concentrations of ~3 ppm in the center of the band represent peak conditions, then the diffusion occurred ~40 °C lower than peak temperature, or ~580 °C, given the relative temperature difference between 2 and 3 ppm from Ti- in-quartz equilibration experiments by Thomas et al. (2010; e.g. Figure 11 of Kidder et al.

(2018)). Figure 10 shows that Ti diffusion at observed length scales ~80 µm (Figure 9) at such temperatures requires time scales on the order of ~100 m.y., i.e. at a time scale similar to the age of peak metamorphism. We interpret then that most of the diffusion had occurred prior to

Neogene Alpine Fault deformation, and as a result that cooling of at least ~40° below peak temperatures had occurred by the Neogene.

As described above, despite the considerable range of initial Ti-in-quartz values of coarse grains, small grains (10-20 um) in the analyzed samples have Ti-in-quartz have values restricted to generally 1-2 ppm. We interpret that the deformation associated with the recrystallization of

23 10-20 µm grains occurred at similar pressure-temperature conditions in the different samples, despite their different peak metamorphic grades. This is consistent with the interpretation of

Cooper and Norris (2011) that the inverted gradient observed in Mosquito Creek was assembled structurally rather than representing a true thermal gradient that existed at the time of Alpine

Fault deformation.

6.5 Ti Activity

Ti activity (aTiO2) represents the availability of Ti present in the sample and can vary from 0.0 to 1.0 and is a major uncertainty in TitaniQ thermobarometry. aTiO2 has been observed, for unknown reasons, to apparently vary by an order of magnitude in rutile-bearing rocks from different localities with otherwise similar temperatures and mineral assemblages (Kidder et al.

(2013, 2018). We consider aTiO2 an unknown, calibration-dependent parameter to be fit using independent constraints. Mineral assemblage is not a reliable indicator of aTiO2 as often assumed

(e.g. Ashley and Law, 2015; Behr and Platt, 2011), at least in low-grade metamorphic rocks. A further complication is the existence of two TitaniQ calibrations. This study uses the Thomas et al. (2010) calibration, but we note that use of the Huang and Audétat (2012) calibration would result in similar temperatures but larger values of aTiO2 (e.g. Cross et al., 2015; Kidder et al.,

2018).

Given the compositional and metamorphic similarities between the central and southern

Alpine Fault, we consider it likely that the aTiO2 of 0.1 (using the Thomas et al. (2010) calibration) interpreted from a large dataset of central Alpine Fault Ti-in-quartz values (Cross et al., 2015; Kidder et al., 2018) is also appropriate for the southern Alpine Fault despite the fact that most samples from both areas contain rutile (Table 5). We can test this hypothesis in several ways, first by making use of peak temperature constraints from RSCM analyses in the southern

24 Alpine Fault (Table 4, Figures 5, 11) assuming that there has been minimal loss of Ti in these grains since peak metamorphism. We interpret that at least some large grains maintain Ti levels from peak metamorphism because of the overall correlation of Ti values with metamorphic grade

(Figure 6), the sluggishness of diffusion of Ti (e.g. Figure 10), and because this would be consistent with the findings of the larger Ti-in-quartz dataset in the central Alpine Fault (Kidder et al., 2018). If we assume a paleo-geothermal gradient of 25 °km-1, consistent with thermobarometric data from the Haast area (Grapes and Otsuki, 1983; Grapes, 1995; Scott, 2015;

Wallace, 1974) and aTiO2 of 0.1, then Ti-in-quartz temperatures are ~50-70 °C below the peak

RSCM temperatures for these samples (Figure 11), i.e. the RSCM and “initial” TitaniQ temperatures are equivalent within likely uncertainties of +/-50° C for the two techniques.

Figure 11 permits a number of other tests for consistency with an aTiO2 of 0.1 and an indication of the range of aTiO2 values that would be consistent with the obtained data. In addition to the constraint (1) evident above, that Ti-in-quartz values in coarse grains should fall near or below

RSCM temperatures, we identify the following additional constraints: 2) dynamic crystallization of quartz does not occur below temperatures of 300-350 °C (Daalen et al., 1999; Dresen et al.,

1997; Dunlap et al., 1997; Stipp et al., 2002; Stöckhert et al., 1999), thus recrystallized veins that cross cut foliation and have Ti-in-quartz concentrations of ~0.2 ppm (Figure 3) should have temperatures >300 °C; 3) since Alpine Fault deformation is associated with exhumation, TitaniQ pressures indicated by dashed lines in Figure 11 should be similar to or less than peak pressures of 4-8 kbar determined by thermobarometry in the Mataketake Range (Grapes, 1995; Scott,

2015; Wallace, 1974).

Examination of Figure 11 demonstrates that aTiO2 > 0.2 would violates criteria 1, placing coarse-grained quartz in high-grade rocks at temperatures less than 100°C lower than RSCM

25 -1 temperatures. An aTiO2 of 0.2 is permissible if the geotherm were lowered to 20 °C km (Figure

11), however at an TiO2 ≥ 0.3 both criteria 1 and 2 are violated for plausible geotherms of 20-30

-1 °C km . The tradeoff between geotherm choice and aTiO2 is further shown in Figure 12. A minimum value for aTiO2 can also be determined, since aTiO2 < 0.04 violates criteria 1 by putting TitaniQ temperatures above RSCM temperatures. To summarize, aTiO2 is dependent on choice of geotherm, but most likely falls between values of 0.04 and 0.2, consistent with the value of 0.1 determined for the central Alpine Fault (Kidder et al., 2018). We emphasize that these activity values are for rocks that generally contain rutile, which according to conventional understanding should indicate an aTiO2 of 1.0. This result suggests that estimates of aTiO2 that rely heavily on mineral assemblage, at least in mylonites, (e.g. Ashley and Law, 2015; Behr and

Platt, 2011; Kohn and Northrup, 2009) may be less well constrained than they appear to be, at least in low temperature rocks.

6.6 Temperature and Pressure of mylonitization

As demonstrated in the previous section, independent constraints on temperature in the interiors of large quartz grains and cross cutting veins, allows aTiO2 for quartz to be constrained to a fairly narrow range 0.04 and 0.2. Using this range of aTiO2 together with RSCM constraints, an estimate of mylonitization temperature of 390-495 °C associated with recrystallized grains of size 10-20 �m (Ti-in-quartz values ~1-2 ppm) can be confidently made. The lower limit of 390

-1 °C assumes equilibrium at 1 ppm Ti, aTiO2 = 0.2 and geothermal gradient of 30 °C km . The upper limit of 495 °C is based on the lowest peak RSCM temperature measured in a mylonite sample (Robinson creek sample NZ75). This temperature range includes lower values than

Alpine Fault mylonitization temperatures estimated from further north, e.g. Cross et al. (2015) estimated deformation temperatures of quartz in two samples of >450–500 °C based on the c-

26 axis opening angle thermometer (Kruhl, 1996; Law, 2014) and microstructural evidence of subgrain rotation and grain boundary migration recrystallization in quartz. Kidder et al. (2018) estimated temperatures of mylonitization from equilibrium Ti-in-quartz values of ~3 ppm to be

>475 °C. Despite slightly higher deformation temperature in the central Alpine Fault relative to the southern Alpine Fault, the extremely high geotherm in the upper crust of the central Alpine

Fault (Cross et al., 2015; Kidder et al., 2018; Sutherland et al., 2017a; Toy et al., 2010), results in deformation at much shallower crustal levels (5-10 km) than values of ~15-20 km that can be estimated from the TitaniQ thermobarometer (Figure 12).

6.7 Recrystallized Grain Size and Stress

The size of recrystallized grains in pure materials deforming by dislocation creep are determined by the differential stress associated with deformation (Twiss, 1977). We interpret, based on the consistency of the recrystallized grain size within the mylonite zone (10-12 �m;

Figure 5), that stress was fairly constant within the mylonites, at least during the last stages of penetrative mylonitization. We infer that the steady decrease in average grain size toward the fault is due to a strain gradient (as previously interpreted in the central Alpine Fault region). We hypothesize that the variations in grain size within the fault-distal rocks are related to local stress or strain perturbations (Figure 5). At least in one case, this perturbation is associated with a seismic event as indicated by the presence of pseudotachylite in the examined sample (Toy et al.,

2011). Such events were apparently either uncommon in the mylonite zone or evidence of their occurrence was overprinted due to greater strain.

We estimate stresses in both Mosquito Creek and Robinson Creek (Table 2, 3) using the

Stipp et. al. (2010) piezometer. The recrystallized grain size of ~10-12 �m within Mosquito and

Robinson Creek suggest stress values of 93-108 MPa. For comparison, the average recrystallized

27 grain size in the central Alpine Fault mylonites is in the range of 25–70 �m (Cross et al., 2015;

Lindroos, 2013; Little et al., 2015; Toy, 2007). This range corresponds to stresses of 23-52 MPa

(Stipp et al., 2010), roughly half the value observed in our study area (Figure 13).

6.8 Strength Evolution of the Alpine Fault

Due to the differences in exhumation rate between the central and southern Alpine Fault, and based on Barth et al.’s (2013) Ar40-Ar39 hornblende age near Martyr River (18 ± 4 Ma), our study area records rocks deformed ~10-15 m.y. earlier than the well-studied mylonites found in the central Alpine Fault. These ages indicate that our rocks were deformed prior to the onset of rapid uplift beginning about 9.24 ± 0.04 Ma (Batt et al., 2000), which presently brings rocks in the central Alpine Fault up to the subsurface at extreme rates. By comparing stress and temperature data from the two areas (central and southern Alpine Fault zone), we can thus provide quantitative constraints on the strength and thermal evolution of the fault, i.e. the southern Alpine Fault can serve as a proxy for rocks that previously were exposed in the north, but are long since eroded away.

The Alpine Fault is one of the very few places that is sufficiently well-constrained to allow for the comparison of independent constraints on rock rheology (Figure 13). We plot peak stress and deformation temperature from our study and the central Alpine Fault in Figure 13.

These values are compared with independent estimates in Figure 13 using the well-constrained strain rate from Alpine Fault mylonites (Norris and Cooper, 2003), the Kidder et al. (2016) flow law which is based on Hirth et al. (2001), a coefficient of friction of 0.17 from Liu and Bird

(2002) numerical model, and estimates of the geothermal gradient for the central Alpine Fault from Kidder et al. (2018). These parameters allow for the plotted strength-depth profile estimates and give an inferred location of the brittle-ductile transition zone (BDTZ). The

28 predicted BDTZ are ~8 km for the central area and ~16 km in southern area. The shallower

BDTZ locality can be attributed to the central region’s rapid uplift rate bringing up hot rock from depth and shifting the ductile regime closer to the surface. The BDTZ in the southern region is interpreted to be in its original location prior to the rapid uplifting occurrence in the central area.

Figure 13 shows consistency between the two estimates of peak stress for the central

Alpine Fault, with a peak differential stress from central area of 51 MPa. Stresses in the

Southern Alpine Fault are predicted to be higher by both approaches (~100 MPa). Figure 13 demonstrates a rough consistency between the various approaches, suggesting that assumptions typically built into strength analyses of the crust (e.g. that the crust can be represented simply by a pure quartz rheology based on laboratory flow laws) may be reasonable to first order.

Acknowledgements

I am incredibly grateful for the mentoring and oversight of this study from my advisor

Steven Kidder. Hamidreza Soleymani provided helpful conversations and analyses in understanding microstructural significance and helped construct Figure 11. Chih-Tung Chen completed the RSCM work in the Mosquito Creek and Robinson Creek localities. Virginia Toy,

Marco Billia and Alan Cooper supplied some of the 29 samples that were used to execute the study.

29 Figure 1. (A) Map of the southern region of the South Island of New Zealand showing metamorphic zones, structural features, and RSCM peak temperature data. (B) Detail of Robinson Creek displaying previously-mapped step-over structures. (C) Map of South Island, New Zealand showing metamorphic zones and location of study area.

30

NW SE NW SE

Oligoclase Zone Oligoclase Zone PeristeritePeristerite Amphibolite Facies Oligoclase TransitionalMylonite Zone Oligoclase 621Mylonite °C Zone 78177 78179 78176 ≥642 °C 78177 78180 500500 mm 78175 78181 78182 Albite-Garnet 78176 78174 Quartzofeldspathic Schist 78175 78182 7817378180 78181 Albite-GarnetZone 534 °C78173 78174 Mylonite 78171 7817078171 Protomylonite Schist 78170 250250 mm 51378169 °C Protomylonite 78169 Mylonite Quaternary Mylonite Quaternary 00 GravelGravel AlpineCataclasite FaultUltramylonite

Amphibolite Facies Greenschist Facies 250250 mm

500500 mm

Figure 2. Schematic cross section of the mylonite zone present at Mosquito Creek. Zoning and widths of sections are from Cooper and Norris (2011). Colored dots represent samples that have Ti-in-quartz data and RSCM peak temperature (values given above). Five-digit numbers are sample numbers. The position of the amphibolite facies and greenschist facies transition relative to the structural boundaries is approximate.

31

a

250

b

100

Figure 3. Cross polarized photomicrographs of a vein in sample 78171 (~108 m southeast of the Alpine Fault) containing evidence of dislocation creep such as undulose extinction (large vein grain at base of image B), subgrains (yellow arrows), recrystallized grains (red arrow) and grain boundary bulges (blue arrows) with similar sizes as seen in the matrix. B is a magnified view of the central part of the vein shown in A. The dark layer to the right of the coarse vein grains is crack filled with epoxy.

32

Figure 4. Cross polarized image with analyzer inserted showing evidence of high-temperature dynamic recrystallization in sample a greenschist facies sample (78170). Foliation is parallel to the mica bands. CPO in the quartz is strong.

33 A. Grain Size (total of all recrystallized fractions)

100 Grain Size (um) 10 1 10 100 1000 Distance to Alpine Fault (m) Robinson Creek A Mosquito Creek A Mosquito Creek B

B. Recrystallized Grain Size

100

10 Grain Size (um) 1 1 10 100 1000 Distance to Alpine Fault (m) Robinson Creek B Mosquito Creek B

C. Finest Fraction Recrystallized Grain Size

100

10 Grain Size (um) 1 1 10 100 1000 Distance to Alpine Fault (m)

Mosquito Creek B Robinson Creek B Mosquito Creek A Robinson Creek A

Figure 5. Grain size data with respect to distance from the Alpine Fault. (A) Figure showing average quartz grain size for Mosquito and Robinson creek. Average sizes include both relict and recrystallized grains. Average grain size decreases towards the fault within the mylonite and ultramylonite zones (distances <300 m). At greater distances from the fault, there is a greater diversity of grain size, including very fine-grained samples that are in some cases are associated with pseudotachylites. (B) Recrystallized grain size is constant except at large distances from the fault. (C) The smallest grain size measured in a line intercept for each sample. Due to the overall increasing stress experienced in exhumational shear zones, these grain sizes may represent grains formed during the last phase of ductile deformation of quartz. Samples in the mylonite and ultramylonite zone show a slight increase in size towards the Alpine Fault. Finest-fraction grain sizes for pseudotachylite-adjacent samples are similar to typical sizes in the mylonite zone. Data are provided in Tables 2, 3.

34

mylonite cataclasite ultramylonite protomylonite

Figure 6. Plot showing the positive correlation between Ti concentration and structural distance from the Alpine Fault. Vein data has been excluded, all other data was used and is illustrated above. The sizes of analyzed grains are represented by the size of the plotted circles. The relationship between increasing grain size and increasing Ti concentrations is not seen clearly in this figure. Data are slightly jittered along the x-axis to increase visibility. The boundaries of the metamorphic zones are approximate. The Ti data are consistent with the overall internal increase in peak metamorphic grade away from the fault (Cooper and Norris, 2011). The increase in Ti within the oligoclase zone is not evident from mineralogical observations alone, but is consistent with the continuation of the increasing temperature to the southeast also evident from RSCM data (e.g. Figure 2).

35

Grain size vs. Titanium concentration

10.0

5.0

2.0 Ti (ppm) 1.0

0.5 OU78169 OU78171 OU78176 OU78177 OU78182 OU78171 Vein OU78176 Vein 0.2

500 200 100 50 20 10 5

Grain Size (um)

Figure 7. Plot of grain size vs. Ti for Ti-in-quartz analyses. In the two oligoclase-zone samples (78177 and 78182, blue and green, respectively), the data show a slight negative trend with larger grain sizes correlating with higher Ti concentrations. With the exception of pseudotachylite sample 78176, most grains of size <20 �m have values between 1 and 2 ppm. The open triangles are analyses from late, cross-cutting veins.

36 5 4 a b

3

2 T3 T1 T2

1 Ti (ppm) Ti .7 .6 .5 .4

.3 T4

200 μm .2 200 μm 200 μm Figure 8. (a) Cross-polarized photomicrograph and (b) Hyperspectral blue-wavelength CL image of sample 78177. Both images show results of Ti-in-quartz analyses. CL and Ti-in-quartz analyses reveals higher Ti concentrations towards CL-bright region of two quartz-rich layers, and lower Ti concentrations along the edges of the layers. T1 – T4 represent transects seen in Figure 9.

37

Ti,Diffusion, Paths 4.50

4.00

3.50

3.00

2.50

2.00 Ti,(ppm)

1.50

1.00

0.50

0.00 0 20 40 60 80 100 120 140 160 180 200 Distance,(µm)

Transect,1 Transect,2 Transect,3 Transect,4

Figure 9. Plot showing Ti-in-quartz concentrations versus distance from the edgeo of quartz-rich bands. Locations of transects are shown Figure 8. The gradients in Ti are interpreted to result from diffusion with a characteristic lengthscale of 60-100 µm.

38 Titanium:in:quartz1Diffusion1 Rate 800

750

700

650 C) ° 600

550

500

Temperature1(450

400

350

300 10 100 1000 Time1(m.y.)

Lengthscale1:160 Lengthscale1:180 Lengthscale1:1100

Figure 10. Plot of the characteristic diffusion time scale of Ti-in-quartz as a function of temperature with length scales of 60, 80, and 100 µm. This study uses a lengthscale of 80 µm seen in bolded purple. Dashed lines represent the peak temperature result from RSCM for this sample (621 °C; upper) and 40 °C less than the peak temperature when diffusion is thought to occur (581 °C; lower). That lines have an uncertainty of +/- 50 °C relative to the RSCM technique (Beyssac et al., 2002). This plot allows for the analysis of time scales relative to the Alpine Fault.

39

Figure 11. Plots showing predicted temperatures with the incorporated geothermal gradients (colored contours) and pressure isobars (dashed black lines) from the used TitaniQ calibration. Boxes show constraints for individual samples based on independent temperature constraints from RSCM (right edge) and the minimum possible temperature of quartz recrystallization (left edge); and mean Ti values +/- standard error for different recrystallized populations. Ti values plotted are 10-20 µm recrystallized grains interpreted to have equilibrated with respect to Ti during late deformation (samples 78169 and 78171); in sample 78177: 90-200 µm grains interpreted to have equilibrated at high temperatures; in sample 78182, 50-200 µm grains that are interpreted to have equilibrated at high temperatures. The plots express how differing aTiO2 and geotherm options impact resulting temperature and pressure estimates and graphically illustrates how aTiO2 values of <0.04 or >0.3 contradict independent temperature constraints (see text for details).

40 a

Ti concentrations of 1 ppm

Figure 12 (a & b). Plots showing the effect of different possible Ti activities and geothermal gradients on predicted TitaniQ pressures and temperatures for the endmembers (1 ppm (a) and 2 ppm (b)) of the interpreted “stable” Ti concentration range of 1 – 2 ppm. Temperatures associated with deformation at between 1 – 2 ppm most likely plot at a temperature <495° (to the left of the grey shaded area), since similar mylonites and recrystallized grain sizes were found to occur at RSMC temperatures less than this value in Robinson creek (Table 4).

41 b

42 Figure 13. Comparison of observed paleopiezometric constraints on strength based on grain size and temperature from Ti-in-quartz measurements for the central and southern Alpine Fault (boxes). The values plotted on the y axis (depth) are converted from temperature estimates using the geotherm from Kidder et al. (2018) for the central Alpine Fault. A geotherm of 25 °C km-1 is assumed for the southern Alpine Fault. Also plotted are theoretical strength depth profiles using the 0.17 effective coefficient of friction from (Liu and Bird, 2002), and the (Kidder et al., 2016) quartz flow law assuming a strain rate of 10-12 s-1. These allow for the interpreted location of the brittle-ductile transition zone (BTDZ). There are considerable uncertainties in depths of the plotted data and relationships, however the figure shows that crustal strength in the central Alpine Fault is roughly half the value in the south, and 2) that the interpreted strength and depths are consistent with the predictions of laboratory-constrained friction and flow laws resulting from a plausible change in geothermal gradient over time.

43 Sample ID Latitude Longitude Distance to AF (m) Composition

78169 35 qfs -43.87018 169.14176 78170 -43.86982 169.14427 86 qfs 78171 - - 108 qfs 78173 -43.8731 169.14843 430 qfs 78174 -43.87403 169.14899 628 qfs 78175 - - 628 qfs 78176 - - 628 qfs 78177 - - 628 chert 78180 -43.87449 169.14927 690 qfs 78181 -43.87495 169.14956 730 qfs 78182 - - 800 qfs NZ 62 -43.8349 169.2180 46 qfs NZ 64 -43.8366 169.2173 223 qfs NZ 65 -43.8378 169.2182 367 qfs NZ 66 -43.8345 169.2184 30 qfs

NZ 68 -43.8350 169.2180 67 qfs

NZ 69A -43.8352 169.2179 89 qfs

NZ 71 -43.8362 169.2172 196 qfs

qfs NZ 73D -43.8376 169.2172 342 NZ 73E -43.8376 169.21724 330 chert NZ 74 - - 321 qfs NZ 75 - - 421 - NZ 76A -43.83806 169.2182 383 chert NZ 76B -43.8381 169.2182 391 qfs NZ 76C -43.8381 169.2182 391 qfs NZ 77 - - 690 qfs NZ 78 -43.83965 169.2193 475 qfs NZ 79 - - 505 - NZ 80 -43.8386 169.21843 461 chert

Table 1. Data on individual samples including latitude and longitude positions, sample distances from the Alpine Fault, and composition of quartzofeldspathic schists (qfs) and metacherts (chert).

Sample Average 1� N Stress +/- SE RXD GS 1� N Stress +/- SE ID GS (�m) (�m) 78169 46.6 3.0 106 31.7 0.2/0.2 10.5 2.2 104 103.5 1.7/1.7 78170 139.9 20.4 140 13.3 0.1/0.1 11.3 2.5 102 97.6 1.7/1.7 78171 179.2 72.1 158 10.9 0.3/0.3 12.2 2.8 101 91.9 1.7/1.6 78173 161.7 12.0 127 11.8 0.1/0.1 19.6 3.9 103 63.1 1/1 78174 63.7 8.2 151 24.8 0.2/0.2 9.8 1.1 114 109.3 0.9/0.9 78175 31.4 4.5 141 43.4 0.4/0.4 7.7 1.2 102 132.4 1.6/1.6 78176 34.6 3.9 161 40.2 0.3/0.3 7.5 1.4 106 135.2 2/1.9 78177 299.8 50.1 206 7.2 0.1/0.1 13.4 3.4 104 85.3 1.7/1.6 78180 284.5 27.6 152 7.6 <0.1/<0.1 20.9 2.6 107 59.9 0.6/0.6 78181 299.7 23.2 106 7.2 <0.1/<0.1 19.6 6.0 102 63.1 1.6/1.5 78182 417.1 34.5 203 5.6 <0.1/<0.1 23.5 5.5 101 54.6 1/1

Table 2. Grain size data from Mosquito Creek of average quartz grains (Average GS), and fine grain recrystallized quartz grains, (RXD GS), with their distance to the Alpine Fault. Average GS represents an average of all quartz grains with no attempt made to distinguish recrystallized grains from non-recrystallized grains. N is the grain count in relation to each sample and type of quartz grain. SE distinguishes the standard error of the different grain types with respect to the standard deviation (1�).

44 Sample ID RXD GS (µm) 1� N Stress +/- SE NZ 62 10.4 1.6 120 104.3 1.2/1.1 NZ 64 10.5 2.3 102 103.5 1.8/1.7 NZ 65 10.2 1.5 104 105.9 1.2/1.2 NZ 66 11.3 1.1 101 97.6 0.8/0.7 NZ 68 10.9 1.7 102 100.5 1.2/1.2 NZ 69A 10.7 2.3 102 102.0 1.8/1.7 NZ 71 11.4 1.2 100 97.0 0.8/0.8 NZ 73D 11.1 1.5 109 99.0 1/1 NZ 73E 10.9 1.1 101 100.5 0.8/0.8 NZ 76A 11.9 1.7 101 93.7 1.1/1 NZ 76B 11.1 1.5 105 99.0 1/1 NZ 76C 11.4 1.7 120 97.0 1.1/1 NZ 78 10.1 1.2 101 106.7 1/1 NZ 80 12.1 2.3 104 92.5 1.4/1.3

Table 3. Recrystallized quartz grain sizes from Robinson Creek and distance to the Alpine Fault. N is the grain count in relation to each sample and SE distinguishes the standard error of the grains with respect to the standard deviation (1�).

Dist. To Metamorphic N R2 1�a Peak T 1� +/- SE Sample ID AF (m) Zone (°C) 78169 35 Garnet 11 0.29 0.04 513 17 5 78171 108 Garnet 13 0.24 0.02 534 10 2.9 78176 628 Oligoclase 11 0.04 0.09 621 39 12 78182 800 Oligoclase 18 N/A N/A ≥642 N/A N/A NZ 68 67 Garnet 10 0.19 0.03 557 15 4.7 NZ 69 89 Garnet 14 0.20 0.03 553 15 4.7 NZ 71 196 Garnet 10 0.28 0.03 518 14 4.3 NZ 74 321 Garnet 10 0.24 0.07 533 29 9.2 NZ 75 421 Garnet 13 0.33 0.04 495 17 4.7 NZ 77 690 Garnet 10 0.24 0.05 533 20 6.4 NZ 78 475 Garnet 10 0.27 0.03 520 14 4.5 NZ 79 505 Garnet 10 0.24 0.03 536 15 4.8

Table 4. T able showing RSCM peak temperatures from Mosquito Creek and Robinson Creek along with the number of spectra (N), R2 ratio and associated standard deviation (1�), calculated peak temperatures and standard deviation (1�) and standard error (SE).

45 Mean Ti N +/- SE Mean Vein Ti N +/- SE Rutile Chlorite Sample Concentration TitaniQ T Concentration Vein T ID (ppm) (°C) (ppm) (°C) 78169 0.9 10 0.2 450 N/A N/A N/A N/A yes no 78171 1.3 23 0.2 465 0.3 5 0.1 350 yes yes 78176 N/A N/A N/A N/A 0.7 3 0.2 425 yes yes 78177 2.3 33 0.1 520 N/A N/A N/A N/A yes no 78182 3.8 19 0.2 575 N/A N/A N/A N/A no yes

Table 5. Ti concentrations and errors of the analyzed samples and associated TitaniQ temperature as discussed in the text (assuming a 25 °/km geotherm and aTiO2 = 0.1) with their respective standard error. Rutile and chlorite columns indicate if the specified minerals were visually distinguished in each sample.

46 References

Ashley, K. T., and Law, R. D., 2015, Modeling prograde TiO2 activity and its signicance for Ti-

in-quartz thermobarometry of pelitic metamorphic rocks: Contributions to Mineralogy

and Petrology, v. 169.

Barth, N. C., Boulton, C., Carpenter, B. M., Batt, G. E., and Toy, V. G., 2013, Slip localization

on the southern Alpine Fault, New Zealand: Tectonics, v. 32, p. 620–640.

Batt, G. E., and Braun, J., 1999, The tectonic evolution of the Southern Alps, New Zealand:

insights from fully thermally coupled dynamical modelling: Geophysical Journal

International, v. 136, no. 2, p. 403-420.

Batt, G. E., Braun, J., Kohn, B. P., and McDougall, I., 2000, Thermochronological analysis of

the dynamics of the Southern Alps, New Zealand: Geological Society of America

Bulletin, v. 112, no. 2, p. 250-266.

Behr, W., Thomas, J., and Hervig, R., 2010, Calibrating Ti concentrations in quartz on the SIMS

using NIST silicate glasses with applications to the TitaniQ geothermobarometer:

Goldschmidt.

Behr, W. M., and Platt, J. P., 2011, A naturally constrained stress profile through the middle

crust in an extensional terrane: Earth and Planetary Science Letters, v. 303, no. 3-4, p.

181-192.

Berryman, K., Beanland, S., Cooper, A. F., Cutten, H. N., norris, R., and Wood, P. R., 1992, The

Alpine Fault, New Zealand: Variation in Quarternary structural style and geomorphic

expression: Annales Tectonicae, Supplement to Vol., v. 6, p. 126-163.

47 Bestmann, M., and Pennacchioni, G., 2015, Ti distribution in quartz across a heterogeneous

shear zone within a granodiorite: The effect of deformation mechanism and strain on Ti

resetting: Lithos, v. 227, p. 37-56.

Bestmann, M., Pennacchioni, G., Mostefaoui, S., Göken, M., and de Wall, H., 2016,

Instantaneous healing of micro-fractures during coseismic slip: Evidence from

microstructure and Ti in quartz geochemistry within an exhumed pseudotachylyte-

bearing fault mylonite: Lithos.

Beyssac, O., Cox, S. C., Vry, J., and Herman, F., 2016, Peak metamorphic temperature and

thermal history of the Southern Alps (New Zealand): Tectonophysics, v. 676, p. 229-249.

Beyssac, O., Goffé, B., Chopin, C., and Rouzaud, J. N., 2002, Raman spectra of carbonaceous

material in metasediments: a new geothermometer: Journal of Metamorphic Geology, v.

20, no. 859-871.

Beyssac, O., Goffé, B., Petitet, J. P., Froigneux, E., and Rouzaud, J. N., 2003, On the

characterization of disordered and heterogeneous carbonaceous material using Raman

spectroscopy.

: Spectrochim. Acta A Mol. Biomol. Spectrosc., v. 59, p. 2267-2276.

Briggs, S. I., and Cottle, J. M., 2018, Accessory mineral petrochronology reveals 30 m.y. of

partial melting during the separation of Zealandia from eastern Gondwana: Lithosphere,

v. 11, p. 169-189.

Briggs, S. I., Cottle, J. M., and Smit M.A, 2018, Record of plate boundary metamorphism during

Gondwana breakup from Lu-HF garnet geochronology of the Alpine Schist: Journal of

Metamorphic Geology.

48 Bull, W. B., and Cooper, A. F., 1986, Uplifted marine terraces along the Alpine Fault, New

Zealand: Science, v. 234, p. 1225-1228.

Cande, and Stock, J., 2004, Pacific-Antarctic-Australia motion and the formation of the

Macquarie Plate: Geophysical Journal International, v. 157, p. 399-414.

Carter, R. M., and Norris, R., 1976, Cenozoic history of southern New Zealand: an accord

between geological observations and plate-tectonic predictions: Earth and Planetary

Science Letters, v. 31, p. 85-94.

Chamberlain, C., Zeitler, P. K., and Cooper, A. F., 1995, Geochronologic constraints of the uplift

and metamorphism along the Alpine Fault, South Island, New Zealand: New Zealand

Journal of Geology and Geophysics.

Cherniak, D. J., Watson, E. B., and Wark, D. A., 2007, Ti diffusion in quartz: Chemical

Geology, v. 236, no. 1-2, p. 65-74.

Cooper, A. F., 1972, Progressive Metamorphism of Metabasic Rocks from the Haast Schist

Group of Southern New Zealand: Journal of Petrology, v. 13, p. 457-492.

Cooper, A. F., Barreiro, B. A., Kimbrough, D. L., and Mattinson, J. M., 1987, Lamprophyre dike

intrusion and the age of the Alpine fault, New Zealand: Geology, v. 15, p. 941-944.

Cooper, A. F., and Ireland, T. R., 2013, Cretaceous sedimentation and metamorphism of the

western Alpine Schist protoliths associated with the Pounamu Ultramafic Belt, Westland,

New Zealand: New Zealand Journal of Geology and Geophysics.

Cooper, A. F., and Norris, R. J., 2011, Inverted metamorphic sequences in Alpine fault mylonites

produced by oblique shear within a plate boundary fault zone, New Zealand: Geology, v.

39, no. 11, p. 1023-1026.

49 Cross, A. J., Kidder, S., and Prior, D., 2015, Using microstructures and TitaniQ thermobarometry

of quartz sheared around garnet porphyroclasts to evaluate microstructural evolution and

constrain an Alpine Fault Zone geotherm: Journal of Structural Geology, v. 75, p. 17-31.

Daalen, M. v., Heilbronner, R., and Kunze, K., 1999, Orientation analysis of localized shear

deformation in quartz fibres at the brittle–ductile transition: Tectonophysics, v. 303, p.

83-107.

Dresen, G., Duyster, J., Stöckhert, B., Wirth, R., and Zulauf, G., 1997, Quartz dislocation

microstructure between 7000 m and 9100 m depth from the Continental Deep Drilling

Program KTB: Journal of Geophysical Research, v. 102, no. B8, p. 18,443-418,452.

Dunlap, W., Hirth, G., and Teyssier, C., 1997, Thermomechanical evolution of a ductile duplex:

Tectonics, v. 16, no. 6, p. 983-1000.

England, P. C., and Molnar, P., 1993, The interpretation of inverted metamorphic isograds using

simple physical calculations: Tectonics, v. 12, no. 1, p. 145-157.

Exner, H. E., 1972, Analysis of grain- and particle-size distributions in metallic materials:

International Metallurgical Reviews, v. 17, p. 25-42.

Gillam, B. G., Little, T. A., Smith, E., and Toy, V. G., 2013, Extensional shear band

development on the outer margin of the Alpine mylonite zone, Tatare Stream, Southern

Alps, New Zealand: Journal of Structural Geology, v. 54, p. 1-20.

Gleason, G. C., and Tullis, J., 1995, A flow law for dislocation creep of quartz aggregates

determined with the molten salt cell: Tectonophysics, v. 247, p. 1-23.

Grapes, R., and Otsuki, M., 1983, Peristerite compositions in quartzofeldspathic schists, Franz

Josef-Fox Glacier Area, New Zealand: Journal of Metamorphic Geology, p. 47-61.

50 Grapes, R., and Watanabe, T., 1992, Metamorphism and uplift of Alpine schist in the Franz

Josef-Fox Glacier area of the Southern Alps, New Zealand: metamophic Geology, v. 10,

p. 171-180.

-, 1994, Mineral composition variation in Alpine Schist, Southern Alps, New Zealand:

Implications for recrystallization and exhumation: The Island Arc, v. 3, p. 163-181.

Grapes, R. H., 1995, Uplift and exhumation of Alpine Schist, Southern Alps, New Zealand:

thermobarometric constraints: New Zealand Journal of Geology and Geophysics, v. 38, p.

525-533.

Grujic, D., Stipp, M., and Wooden, J. L., 2011, Thermometry of quartz mylonites: Importance of

dynamic recrystallization on Ti-in-quartz reequilibration: Geochemistry Geophysics

Geosystems, v. 12, no. 6, p. Q06012.

Haertel, M., Herwegh, M., and Pettke, T., 2013, Titanium-in-quartz thermometry on

synkinematic quartz veins in a retrograde crustal-scale normal fault zone:

Tectonophysics.

Hirth, G., Teyssier, C., and Dunlap, W., 2001, An evaluation of quartzite flow laws based on

comparisons between experimentally and naturally deformed rocks: International Journal

of Earth Sciences (Geol Rundsch), v. 90, p. 77-87.

Hirth, G., and Tullis, J., 1992, Dislocation creep regimes in quartz aggregates: Journal of

Structural Geology, v. 14, no. 2, p. 145 - 159.

Huang, R., and Audétat, A., 2012, The titanium-in-quartz (TitaniQ) thermobarometer: A critical

examination and re-calibration: Geochimica Et Cosmochimica Acta, v. 84, p. 75-89.

Jessel, M. W., 1987, Grain-boundary migration microstructures in a naturally deformed

quartzite: Journal of Structural Geology, v. 9, no. 8, p. 1007-1014.

51 Jochum, K. P., Weis, U., Stoll, B., Kuzmin, D., Yang, Q., Raczek, I., Jacob, D. E., Stracke, A.,

Birbaum, K., Frick, D. A., Günther, D., and Enzweiler, J., 2011, Determination of

Reference Values for NIST SRM 610--617 Glasses Following ISO Guidelines:

Geostandards and Geoanalytical Research, v. 35, no. 4, p. 397-429.

Kalceff, M. A. S., and Phillips, M. R., 1995, Cathodoluminescence microcharacterisation of the

defect structure of quartz: Physical Review B, v. 52, no. 5, p. 3122-3144.

Kamp, P. J. J., 1997, Paleogeothermal gradient and deformation style, Pacific front of the

Southern Alps Orogen: Constraints from fission track thermochronology:

Tectonophysics, v. 271, p. 37-58.

Kidder, S. B., Avouac, J. P., and Chan, Y. C., 2012, Constraints from rocks in the Taiwan orogen

on crustal stress levels and rheology: Journal of Geophysical Research, v. 117.

-, 2013, Application of titanium-in-quartz thermobarometry to greenschist facies veins and

recrystallized quartzites in the Hsüehshan range, Taiwan: Solid Earth, v. 4, p. 1-21.

Kidder, S. B., Hirth, G., Avouac, J. P., and Behr, W. M., 2016, The influence of stress history on

the grain size and microstructure of experimentally deformed quartzite: Journal of

Structural Geology, v. 83, p. 194-206.

Kidder, S. B., Toy, V. G., Prior, D. J., Little, T. A., Khan, A., and MacRae, C., 2018, Constraints

on Alpine Fault (New Zealand) mylonitization temperatures and the geothermal gradient

from Ti-in-quartz thermobarometry: Solid Earth, v. 9, no. 5, p. 1123-1139.

Kohlstedt, D. L., Evans, B., and Mackwell, S. J., 1995, Strength of the lithosphere–constraints

imposed by laboratory experiments: Journal of Geophysical Research-Solid Earth, v. 100,

no. B9, p. 17587-17602.

52 Kohn, M. J., and Northrup, C. J., 2009, Taking mylonites' temperatures: Geology, v. 37, no. 1, p.

47-50.

Koons, P. O., 1987, Some thermal and mechanical consequences of rapid uplift: an example

from the Southern Alps, New Zealand: Earth and Planetary Science Letters, v. 86, p. 307-

319.

Kruhl, J. H., 1996, Prism- and basal-plane parallel subgrain boundaries in quartz: a

microstructural geothermobarometer: Journal of Metamorphic Geology, v. 14, p. 581-

589.

Lamb, S., Mortimer, N., Smith, E., and Turner, G., 2016, Focusing of relative plate motion at a

continental transform fault: Cenozoic dextral displacement >700 km on New Zealand’s

Alpine Fault, reversing >225 km of Late Cretaceous sinistral motion: Geochemistry,

Geophysics, Geosystems, v. 17, p. 1197-1213.

Law, R. D., 2014, Deformation thermometry based on quartz c-axis fabrics and recrystallization

microstructures: A review.: Journal of Structural Geology, v. 66, p. 129-161.

Leeman, W. P., MacRae, C. M., Wilson, N. C., Torpy, A., Lee, C.-T. A., Student, J. J., Thomas,

J. B., and Vicenzi, E. P., 2012, A Study of Cathodoluminescence and Trace Element

Compositional Zoning in Natural Quartz from Volcanic Rocks: Mapping Titanium

Content in Quartz.: Microscopy and Microanalysis, v. 18, p. 1322-1341.

Lindroos, Z. R., 2013, Microstructures Developed during Creep Deformation of Quartz: an

Investigation of Shape and Crystallographic Characteristics [M. Sc.: University of Otago,

119 p.

53 Little, T., and Mortimer, N., 2001, Rotation of ductile fabrics across the Alpine Fault and

Cenozoic bending fo the New Zealand orocline: Journal of the Geological Society,

London, v. 158, p. 745-756.

Little, T., Wightman, R., Holcombe, R. J., and Hill, M. P., 2007, Transpression Models and

Ductile Deformation of the Lower Crust of the Pacific Plate in the Central Southern Alps,

a Perspective From Structural Geology: American Geophysical Union.

Little, T. A., Cox, S., Vry, J. K., and Batt, G., 2005, Variations in exhumation level and uplift

rate along the obliqu-slip Alpine fault, central Southern Alps, New Zealand: Geological

Society of America Bulletin, v. 117, no. 5-6, p. 707-723.

Little, T. A., Holcombe, R. J., and Ilg, B. R., 2002, Kinematics of oblique collision and ramping

inferred from microstructures and strain in middle crustal rocks, central Southern Alps,

New Zealand: Journal of Structural Geology, v. 24, p. 219-239.

Little, T. A., Prior, D. J., Toy, V. G., and Lindroos, Z. R., 2015, The link between strength of

lattice preferred orientation, second phase content and grain boundary migration: A case

study from the Alpine Fault zone, New Zealand: Journal of Structural Geology, v. 81, p.

59-77.

Liu, Z., and Bird, P., 2002, Finite element modeling of neotectonics in New Zealand: Journal of

Geophysical Research, v. 107(B12).

MacRae, C. M., Wilson, N. C., and Torpy, A., 2013, Hyperspectral cathodoluminescence:

Mineralogy and Petrology, v. 107, no. 3, p. 429-440.

Mason, B., 1962, Metamorphism in the Southern Alps of New Zealand: Bull. Am. Mus. Nat.

Hist., v. 123, p. 211-248.

54 Mortimer, N., 2018, Evidence for a pre-Eocene proto-Alpine Fault through Zealandia: New

Zealand Journal of Geology and Geophysics.

Mortimer, N., and Cooper, A. F., 2004, U-Pb and Sm-Nd ages from the Alpine Schist, New

Zealand: New Zealand Journal of Geology and Geophysics.

Mortimer, N., McLaren, S., and Dunlap, W., 2012, Ar-Ar dating of K-feldspar in low grade

metamorphic rocks: Example of an exhumed Mesozoic accretionary wedge and forearc,

South Island, New Zealand: Tectonics, v. 31.

Nachlas, W., and Hirth, G., 2015, Experimental constraints on the role of dynamic

recrystallization on resetting the Ti-in-quartz thermobarometer: Journal of Geophysical

Research, v. 120.

Nachlas, W., Thomas, J., and Hirth, G., 2018, TitaniQ deformed: Experimental deformation of

out-of-equilibrium quartz porphyroclasts: Journal of Structural Geology.

Nevitt, J. M., Warren, J. M., Kidder, S., and Pollard, D. D., 2017, Comparison of thermal

modeling, microstructural analysis, and Ti-in-quartz thermobarometry to constrain the

thermal history of a cooling pluton during deformation in the Mount Abbot Quadrangle,

CA: Geochemistry Geophysics Geosystems, v. 18, p. 1270-1297.

Norris, R., and Cooper, A., 2007, The Alpine Fault, New Zealand: Surface Geology and Field

Relationships: American Geophysical Union, p. 157-175.

Norris, R., and Toy, V., 2014, Continental transforms: A view from the Alpine Fault: Journal of

Structural Geology, v. 64, p. 3-31.

Norris, R. J., and Cooper, A. F., 2000, Late Quaternary slip rates and slip partitioning on the

Alpine Fault, New Zealand: Journal of Structural Geology, v. 23, p. 507-520.

55 Norris, R. J., and Cooper, A. F., 2003, Very high strains recorded in mylonites along the Alpine

Fault, New Zealand: implications for the deep structure of plate boundary faults: Journal

of Structural Geology, v. 25, p. 2141-2157.

Passchier, C. W., and Trouw, R. A. J., 2005, Microtectonics, 2nd edition, Berlin Heidelberg,

Springer-Verlag.

Rattenbury, M. S., Jongens, R., and Cox, S. C., 2010, Geology of the Haast area. Institute of

Geological & Nuclear Sciences 1:250 000 geological map 14., Lower Hutt, New Zealand,

GNS Science, 1 sheet + 58 p.:

Reed, J. J., 1958, Regional metamorphism in Southeast Nelson: Bull. N.Z. Geolo. Surv., v. 60.

Ring, U., Mortimer, N., and Deckert, H., 2019, Critical wedge theory and the Mesozoic

accretionary wedge of New Zealand: Journal of Structural Geology.

Scott, J. M., 2015, New P-T and U-Pb constraints on Alpine Schist metamorphism in south

Westland, New Zealand: New Zealand Journal of Geology and Geophysics.

Shi, Y. L., Allis, R., and Davey, F., 1996, Thermal modeling of the Southern Alps, New Zealand:

Pure and Applied Geophysics, v. 146, no. 3-4, p. 469-501.

Sibson, R. H., White, S. H., and Atkinson, B. K., 1981, Structure and distribution of fault rocks

in the Alpine Fault Zone, New Zealand: Geological Society of London Special

Publication, v. 9, no. 197-210, p. 197-210.

Stern, T., Okaya, D., Kleffmann, S., Scherwath, M., Henrys, S., and Davey, F., 2007,

Geophysical Exploration and Dynamics of the Alpine Fault Zone, in Okaya, D., Stern, T.,

and Davey, F., eds., A Continental Plate Boundary: Tectonics at South Island, New

Zealand Washington D.C.

56 Stipp, M., Stünitz, H., Heilbronner, R., and Schmid, S. M., 2002, Dynamic recrystallization of

quartz: correlation between natural and experimental conditions: Deformation

Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives, v. 200, p.

171-190.

Stipp, M., Tullis, J., and Behrens, H., 2006, Effect of water on the dislocation creep

microstructure and flow stress of quartz and implications for the recrystallized grain size

piezometer: Journal of Geophysical Research, v. 111, no. B4, p. 1-19.

Stipp, M., Tullis, J., Scherwath, M., and Behrmann, J. H., 2010, A new perspective on

paleopiezometry: Dynamically recrystallized grain size distributions indicate mechanism

changes: Geology, v. 38, no. 8, p. 759-762.

Stöckhert, B., Brix, M. R., Kleinschrodt, R., Hurford, A. J., and Wirth, R., 1999,

Thermochronometry and microstructures of quartz - a comparison with experimental

flow laws and predictions on the temperature of the brittle-plastic transition: Journal of

Structural Geology, v. 21, no. 3, p. 351-369.

Sutherland, R., 1995, Displacement since the Pliocene Along the Southern Section of the Alpine

Fault, New Zealand - Reply: Geology, v. 23, p. 476 - 476.

Sutherland, R., 1999, Cenozoic Bending of New Zealand basement terranes and Alpine Fault

displacement: A brief review: New Zealand Journal of Geology and Geophysics, v. 42, p.

295-301.

Sutherland, R., Townend, J., Toy, V., Phaedra, U., Coussens, J., Allen, M., Baratin, L.-M., Barth,

N. C., Becroft, L., Boese, C. M., Boles, A., Boulton, C., Broderick, N. G. R., Janku-

Capova, L., Carpenter, B. M., Célérier, B., Chamberlain, C., Cooper, A. F., Coutts, A.,

Cox, S., Craw, L., Doan, M.-L., Eccles, J. D., Faulkner, D. R., Grieve, J., Grochowski, J.,

57 Gulley, A., Hartog, A., Howarth, J., Jacobs, K., Jeppson, T., Kato, N., Keys, S., Kirilova,

M., Kometani, Y., Langridge, R. M., Lin, W., Little, T., Lukacs, A., Mallyon, D.,

Mariani, E., Massiot, C. c., Mathewson, L., Melosh, B., Menzies, C. D., Moore, J.,

Morales, L., Morgan, C., Mori, H., Niemeijer, A. R., Nishikawa, O., Prior, D., Sauer, K.,

Savage, M., Schleicher, A., Schmitt, D. R., Shigematsu, N., Taylor-Offord, S., Teagle,

D., Tobin, H., Valdez, R., Weaver, K., Wiersberg, T., Williams, J., Woodman, N., and

Zimmer, M., 2017a, Extreme hydrothermal conditions at an active plate-boundary fault:

Nature, v. 000.

Sutherland, R., Townend, J., Toy, V., Upton, P., Coussens, J., Allen, M., Baratin, L.-M., Barth,

N., Becroft, L., Boese, C., Boles, A., Boulton, C., Broderick, N. G. R., Janku-Capova, L.,

Carpenter, B. M., Célérier, B., Chamberlain, C., Cooper, A., Coutts, A., Cox, S., Craw,

L., Doan, M.-L., Eccles, J., Faulkner, D., Grieve, J., Grochowski, J., Gulley, A., Hartog,

A., Howarth, J., Jacobs, K., Jeppson, T., Kato, N., Keys, S., Kirilova, M., Kometani, Y.,

Langridge, R., Lin, W., Little, T., Lukacs, A., Mallyon, D., Mariani, E., Massiot, C. c.,

Mathewson, L., Melosh, B., Menzies, C., Moore, J., Morales, L., Morgan, C., Mori, H.,

Niemeijer, A., Nishikawa, O., Prior, D., Sauer, K., Savage, M., Schleicher, A., Schmitt,

D. R., Shigematsu, N., Taylor-Offord, S., Teagle, D., Tobin, H., Valdez, R., Weaver, K.,

Wiersberg, T., Williams, J., Woodman, N., and Zimmer, M., 2017b, Extreme

hydrothermal conditions at an active plate-bounding fault: Nature, v. 546, p. 137-140.

Sutherland, R., Toy, V. G., Townend, J., Cox, S. C., Eccles, J. D., Faulkner, D. R., Prior, D. J.,

Norris, R. J., Mariani, E., Boulton, C., Carpenter, B. M., Menzies, C. D., Little, T. A.,

Hasting, M., Pascale, G. P. D., Langridge, R. M., Scott, H. R., Reid, Z., Lindroos,

58 Fleming, B., and Kopf, A. J., 2012, Drilling reveals fluid control on architecture and

rupture of the Alpine fault, New Zealand: Geology, v. 40, no. 12, p. 1143-1146.

Thomas, J., Spear, F., and Nayak, S., 2010, TitaniQ under pressure: the effect of pressure and

temperature on the solubility of Ti in quartz: Contributions to Mineralogy and Petrology.

Thomas, J., Watson, B., Spear, F., and Wark, D., 2015, TitaniQ recrystallized: experimental

confirmation of the original Ti-in-quartz calibrations: Contributions to Mineralogy and

Petrology.

Toy, V., Ritchie, S., and Sibson, R. H., 2011, Diverse habitats of pseudotachylytes in the Alpine

Fault Zone and relationships to current seismicity: The Geological Society of London, v.

359, p. 115-133.

Toy, V. G., 2007, Rheology of the Alpine Fault Mylonite Zone: deformation processes at and

below the base of the seismogenic zone in a major plate boundary structure [PhD:

University of Otago, 629 p.

Toy, V. G., Craw, D., Cooper., A. F., and Norris, R. J., 2010, Thermal regime in the central

Alpine Fault zone, New Zealand: Constraints from microstructures, biotite chemistry, and

fluid inclusion data: Tectonophysics, v. 485, p. 178-192.

Toy, V. G., Norris, R. J., Prior, D. J., Walrond, A. F., and Cooper, A. F., 2013, How do

lineations reflect the strain history of transpressive shear zones? The example of the

active Alpine Fault zone, New Zealand: Journal of Structural Geology, v. 50, p. 187-198.

Toy, V. G., Prior, D. J., Norris, R. J., Cooper, A. F., and Walrond, M., 2012, Relationships

between kinematic indicators and strain during syn-deformational exhumation of an

oblique slip, transpressive, plate boundary shear zone: The Alpine Fault, New Zealand:

Earth and Planetary Science Letters, v. 333-334, p. 282-292.

59 Trouw, R. A. J., Passchier, C. W., and Wiersma, D. J., 2010, Atlas of Mylonites- and related

microstructures, Springer, 322 p.:

Turner, F. J., 1933, The metamorphic and intrusive rocks of southern Westland: Trans. N.Z.

Inst., v. 63, p. 178-284.

Twiss, R. J., 1977, Theory and Applicability of a Recrystallized Grain-Size Paleopiezometer:

Pure and Applied Geophysics, v. 115, no. 1-2, p. 227-244. van Daalen, M., Heilbronner, R., and Kunze, K., 1999, Orientation analysis of localized shear

deformation in quartz fibres at the brittle-ductile transition: Tectonophysics, v. 303, no. 1-

4, p. 83-107.

Vry, J., Baker, J., Maas, R., Little, T. A., Grapes, R., and Dixon, M., 2004, Zoned (Cretaceous

and Cenozoic) garnet and the timing of high grade metamorphism, Southern Alps, New

Zealand: Journal of Metamorphic Geology, v. 22, p. 137-157.

Walcott, R. I., 1998, Present tectonics and late Cenozoic evolution of New Zealand: Reviews of

Geophysics, v. 36, p. 1-26.

Wallace, R. C., 1974, Metamorphism of the Alpine Schist, Mataketake Range, South Westland,

New Zealand: Journal of the Royal Society of New Zealand, p. 253-266.

Wark, D., and Spear, F., Ti in quartz: Cathodoluminescence and thermometry, in Proceedings

Goldschmidt 20052005, Volume 15, Geochemical Society, Moscow, Idaho, p. A592.

Wellman, H. W., 1979, An uplift map for the South Island of New Zealand, and a model for

uplift of the Southern Alps: Journal of the Royal Society of New Zealand, v. 4, p. 253 -

266.

60