STRAIN LOCALIZATION IN THE UPPER CRUST AND STRESS FIELD EVOLUTION ADJACENT TO THE ALPINE FAULT

IN NORTHERN FIORDLAND, NEW ZEALAND

A Thesis Presented

by

Phoebe A. Judge

to

The Faculty of the Graduate College

of

The University of Vermont

In Partial Fulfillment of the Requirements for the Degree of Master of Science

Specializing in Geology

October, 2006

Accepted by the Faculty of the Graduate College, The University of Vermont, in partial fulfillment of the requirements for the degree of Master of Science, specializing in Geology. Thesis Examination Committee: ________________________________ Advisor Keith A. Klepeis, Ph.D.

________________________________ Thomas Neumann, Ph.D.

________________________________ Chairperson Mandar Dewoolkar, Ph.D. ________________________________ Vice President for Frances E. Carr, Ph.D. Research and Dean Of Graduate Studies Date: August 28, 2006

ii

Abstract In this thesis, I present structural, kinematic, and stress inversion data from the Darran Range in northern Fiordland, New Zealand. The ~800 km2 Darran Range has excellent exposure of faults within the upper crust, and provides an opportunity to study strain localization and fluid-induced weakening processes adjacent to the obliquely convergent Australia-Pacific plate boundary. Fault-slip data collected in the region outline several distinct zones of deformation, as well as the presence of elevated pore fluid pressure in the upper crust. Strain localization and weakening mechanisms, including elevated fluid pressure, may assist in explaining differences in deformation, such as the width of the deformation zone and the continued reactivation of major faults, near collisional plate boundaries. Kinematic solutions and stress inversions reveal spatial variations in the degree of strain localization and strike-slip partitioning adjacent to the Alpine Fault. Stress inversions of fault-slip data from major fault segments within ~10 km of the plate boundary show evidence of elevated pore fluid pressure. Geometrical and frictional constrains on the analysis of stress tensors calculated from stress inversions indicate that the coefficient of friction is extremely low (µ = 0.10) near the Alpine Fault. Compression axes are oriented ~60º from the dominantly northeasterly strike of the plate boundary, and this orientation, combined with the low coefficient of friction, suggest a weakening of the crust around the Alpine Fault Zone. Deformation within 10 km of the plate boundary is characterized by reverse, oblique-reverse, and strike-slip fault populations. At distances greater than 10 km to the southeast of the plate boundary zone, deformation is characterized by oblique-reverse and strike-slip motion on reactivated steep, brittle faults; vertical motion is predominantly localized at lithologic boundaries. This deformation results in the extrusion of wedge-shaped blocks in the Darran Range. Cross-cutting relationships and kinematic analysis indicate the superposition of stress fields in northern Fiordland, including an older phase of normal faulting from the Late Cretaceous – Early Tertiary. Descriptions of strain localization mechanisms, such as fluid-induced weakening and the partitioning of strain in different regions, improves our understanding of deformation processes at collisional plate boundaries, and how the upper crust responds to tectonic stresses. The dominant strain localization mechanism in northern Fiordland is elevated pore fluid pressure in the near-boundary deformation zone. The inherited structure of the region assists in the localization of strain along rheological and lithologic boundaries, but it is subordinate to fluid-induced weakening.

iii

Acknowledgements

I owe heart-felt thanks to many people for their help, knowledge, and support during

this project. Above all, I am extremely grateful to Keith Klepeis for being an outstanding

advisor and mentor, and for having such a thoroughly engaging project. And for

hesitantly eating silverbeets in the Milford Lodge. Many thanks also to Tom Neumann

and Mandar Dewoolkar for being part of my thesis committee.

I am indebted to the all wonderful faculty, staff, and students in the Geology

department. Tom Neumann and Adam Schoonmaker provided me with entertaining

comments and very helpful suggestions during the crucial periods when I thought neither

were available. Dan King and Rob Zimmermann were superior field “assistants.”

(‘Assistant’ does not properly evoke the help they provided, but it shall have to suffice.)

Certainly I owe thanks to Bernard Célérier, Rick Allmendinger, and Arnaud

Etchecopar for willfully serving as my adopted structure community, and answering

endless questions about stress and strain models. I am also grateful for the education and

support I received from the Geology and Physics departments at Mount Holyoke College

in preparation for graduate work, especially from Michelle Markley and Janice Hudgings.

The National Science Foundation, the UVM Geology department, and the New

Zealand Institute of Geological and Nuclear Sciences in Dunedin have given me financial

support, a job, and technical support (or some combination thereof) for the past two

years, and deserve proper recognition.

I owe thanks to all of my dear friends, in and out of Vermont, for their incredible

support when I needed it most, and for knowing when that was.

Last but not least, thanks again to my parents for just about everything else.

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Table of Contents Acknowledgements.......................................................................................................iii List of Tables ...............................................................................................................vii List of Figures.............................................................................................................viii Chapter I: Introduction ...............................................................................................9

1. Overview of project .................................................................................................9

2. Research methodology...........................................................................................12

2.1 Field methods ..................................................................................................12

2.2 Laboratory methods.........................................................................................13

3. Thesis outline ........................................................................................................14

Chapter II: Literature Review...................................................................................15

1. Overview of the geology and the tectonic history of Fiordland, New Zealand ........15

1.1 Tectonic & geologic history of Fiordland and the Australia-Pacific plate

boundary ...............................................................................................................15

1.2 The modern Australia-Pacific plate boundary..................................................19

1.3 The modern structure of Northern Fiordland ...................................................21

2. Transpressional tectonics ......................................................................................22

3. The role of fluid in upper crustal faulting...............................................................24

3.1 Pore fluid pressure ..........................................................................................25

3.2 Fluid infiltration and frictional weakening of crustal-scale faults ....................26

4. The application of fault-slip data to tectonic settings .............................................28

Chapter III: Strain localization in the upper crust adjacent to the tectonically active Alpine Fault Zone in Fiordland, New Zealand...........................................................34

1. Introduction..........................................................................................................35

2. Geologic history and evolution of the Australia-Pacific plate boundary.................38

v

2.1 Geology and tectonic history of Fiordland & the Australia-Pacific plate

boundary ...............................................................................................................38

2.2 Modern tectonic setting of the Alpine Fault......................................................39

3. Structure of field regions in Fiordland..................................................................41

3.1 Structure of the Darran Range and Northern Fiordland..................................41

3.1a The Hollyford Valley Fault Zone...................................................................42

3.1b The interior of the Darran Range..................................................................46

3.1c The northern margin of the Darran Range ....................................................47

3.1d The Harrison-Kaipo Fault Zone....................................................................48

3.2 Structure of the Skippers Range ......................................................................49

3.3 Structure of Doubtful Sound............................................................................52

4. Kinematics of the Darran Range...........................................................................52

4.1 Kinematic methods..........................................................................................52

4.2 Kinematic results ................................................................................................54

4.2a The Hollyford Fault Valley and the interior of the Darran Range..................54

4.2b The northern margin of the Darran Range and the Harrison-Kaipo Fault

Valley ....................................................................................................................57

4.2c Doubtful Sound .............................................................................................59

5. Stress inversion.....................................................................................................60

5.1 Methods...........................................................................................................60

5.2 Results using the geometric constraint ............................................................67

5.3 Results using the friction constraint ................................................................69

6. Discussion .............................................................................................................70

6.1 Normal faults in Fiordland .............................................................................70

6.2 The evolution of the stress field in the Darran Range .......................................72

6.3 Fluid infiltration and strain-induced weakening of the crust ...........................75

6.4 Strain localization and deformation partitioning.............................................77

7. Conclusions...........................................................................................................81

vi

Chapter IV: Discussion ..............................................................................................83

1. Overview and conclusions .....................................................................................83

2. Future work...........................................................................................................85

Bibliography ................................................................................................................87 Appendices: ............................................................................................................... 101

vii

List of Tables Table 3.1: Fault populations used for geometric constraint experiments and the principle stress axis orientations..................................................................................62

viii

List of Figures Figure 2.1: a. Tectonic setting of New Zealand and the Australia/Pacific plate boundary.......................................................................................................................................16 Figure 3.1: a. Tectonic setting of New Zealand and the Australia/Pacific plate boundary.......................................................................................................................................37 Figure 3.2: Topography, sitemap, and cross-section of the Darran Range northern Fiordland. ..................................................................................................43 Figure 3.3: Photographs, sketches, and photomicrographs from sites in the Darran Range.........................................................................................................45 Figure 3.4: a. Map of Skippers Range showing location of cross-section. .....................51 Figure 3.5: Detailed map of southern region showing normal fault slip data .................55 Figure 3.6: Detailed map of southern region showing dextral fault slip data .................57 Figure 3.7: a. Detailed map of northern region showing fault slip data (b - i) and fault- 58 Figure 3.8: Simplified geologic map of Doubtful Sound region ....................................58 Figure 3.9: Results of frictional constraint method........................................................66 Figure 3.10: Results of geometric constraint method. ...................................................68 Figure 3.11: A DEM showing the central and northern Fiordland.................................71 Figure 3.12: Equal-area lower-hemisphere stereoplot of stress and strain data from.....73 Figure 3.13: Cartoon cross-section of the plate boundary..............................................78

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Chapter I: Introduction

1. Overview of project

Deformation at obliquely convergent plate boundaries often includes uplift,

shortening, and lateral strike-slip motion, but the dominant mechanisms that control these

features is unresolved. Strain localization processes such as elevated pore fluid pressure

or the preferential reactivation of a weak fault are important in controlling deformation

patterns (Stewart et al., 2000; Bunds, 2001; Rutter et al., 2001), but it may be difficult to

identify features related to inherited structures versus those signals related to deformation

mechanisms (Koons et al., 1998; Little et al., 2002a). Studying the development of

structures in the upper crust near obliquely convergent plate boundaries assists in the

description of large-scale processes that control deformation in these areas. The

Fiordland region of the Alpine Fault Zone in New Zealand is a zone of transpression

because it accommodates both shearing and stretching related to oblique convergence.

Additionally, the region contains excellent exposure of structures in the upper crust. A

more complete understanding of the nature of processes controlling deformation

partitioning and strain localization in regions such as Fiordland can help to refine models

that describe continental tectonics. Describing the structural evolution and crustal

strength of the Alpine Fault may generate new models and predictions for the behavior of

other plate boundary systems.

For this project, I use field-based structural observations combined with quantitative

modeling of strain and strain fields to study deformation processes in the upper crust.

Fault-slip data collected in the field allowed me to make spatial and temporal

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observations about structures present in northern Fiordland. I then used these data to

model the kinematic compatibility of fault populations in different deformation zones,

and to model the stress fields in which fault populations can be activated. By collecting

these data and using quantitative models, I was able to answer several specific questions

about deformation and strain localization processes in the upper crust: 1) In a region of

transpression with several superposed fault populations, is there evidence of strain

partitioning within the populations? 2) Is there a specific style of structure that controls

strain localization at obliquely collisional plate boundaries? 3) What role do strain-

induced weakening processes, such as elevated pore fluid pressure or the presence of

inherited structures, have in reducing the work required to strain the crust near a plate

boundary? 4) Do fault populations in Fiordland preserve evidence of superimposed stress

fields?

There are several primary results of this study. First, I document the presence of a 10

km-wide zone of deformation adjacent to the Alpine Fault that contains fault populations

consistent with the modern stress field. Fluid infiltration and frictionally weak faults are

also characteristic of this zone. Outside of the near-boundary zone, strain is localized

along lithologic boundaries, and on steep brittle faults that are reactivated. Secondly, this

work shows the frictionally weak characteristsics of the Alpine Fault in Fiordland. The

frictionally weak faults and the moderately high angles between the plate boundary and

compression angles confirm previous descriptions of the Alpine Fault as a weak plate

boundary fault. Finally, this work describes several superimposed phases of faults

preserved in central and northern Fiordland, including phases of older extension.

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Results of this study have both local and regional tectonic significance. Locally, this

study is significant because of its new description of the interior of the Darran Range, and

the documentation and interpretation of normal faults in Fiordland. The Darran Range is

characterized by widely-spaced (500 – 1000 m) steep dextral faults that record a

component of vertical motion. These dextral faults everywhere cut across older normal

faults in the Darran Range and elsewhere in Fiordland. Normal faults in the region

represent two phases of extension that likely occurred during the Late Cretaceous and the

Oligocene.

The regional significance of this study is its description of fluid-induced strain

localization and partitioning processes, and of the relatively weak nature of the Alpine

Fault. The existence of the near-boundary deformation zone is consistent with other

studies of strain localization and GPS surveys from the South Island of New Zealand

(Norris & Cooper, 2001; Sutherland et al., 2006). I show that it is likely that the elevated

pore fluid pressure near the plate boundary is partially responsible for the localization of

strain in the near-boundary zone. The enhanced fluid pressure in the region contributes

to the frictionally weak nature of faults in this zone, and suggests that the Alpine Fault in

northern Fiordland may be generally weak. This weakness may allow the Alpine Fault to

continue slipping for several hundreds of kilometers within a narrow fault zone. Several

studies of faults and earthquake data near the plate boundary in other sections of the

South Island (Liu & Bird, 2001; Balfour et al., 2005), which suggests that the Alpine

Fault may be a weak plate boundary fault.

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2. Research methodology

2.1 Field methods

I completed the field-based component of this work over two field seasons in

northern Fiordland. I collected data from 8 field sites near Mount Thunder, Lake Truth,

in the central Skippers Range, from the southeast of Mount Madeline between Madeline

Creek and Catch Creek, Homer’s Tunnel, Gertrude’s Saddle, and from a prominent

roadcut between Key Summit and the Lower Hollyford Road on the Milford – Te Anau

Road (Fig. 3.2 B). Sites were selected on the basis of their proximity to mapped faults,

lithology, exposure of bedrock, and accessibility (for helicopter access). Helicopter

access to remote regions such as Mount Thunder, Lake Truth, and the Skippers Range

allowed me to spend the majority of each field season collecting data instead of traveling

to field sites.

Fieldwork at each site included the identification of lithologies and mineral

assemblages, measurement of the orientation of fault planes and fractures, and the

orientation of slip indicators such as slickenlines, grooves, mineral lineations,

chattermarks, and offset markers; and the collection of oriented samples to make thin-

sections to determine composition and kinematics on a microstructural scale. Field notes

included sketches and maps, and observations of ductile fabrics, larger patterns of

dominant fault populations, and other features that may have been site-specific.

Field sites are located on Figure 3.2 B. Fault-slip data are presented in Appendices A

and B, and are plotted on equal-area lower-hemisphere stereographic projections on

detailed maps of geographic regions in Chapter III. Descriptions and interpretations of

structural and kinematic data are discussed in detail in Chapter III.

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2.2 Laboratory methods

I used two computer-modeling techniques for this study that invert the fault-slip data

collected in the field to determine average stress and strain axes. The first technique

involved the calculation of fault-plane solutions for fault populations using the program

FaultKin v. 4.3.5, created by R. W. Allmendinger, R. A. Marrett, & T. Cladouhos (1994;

modified 2006). This program allowed me to determine the kinematic compatibility of

populations from each field site and from field sites within a specific region. I also used

the results to compare the orientation of contraction and extension axes to compressional

(P) axes from regional earthquakes and to principle stress axes determined from stress

inversions. The method and results of this technique are fully described in Chapter III.

The second quantitative technique involved the inversion of fault-slip data using Fault

Slip Analysis (FSA) v. 28.5 created by B. Célérier (1988; modified 2006). I used this

program to model the orientation of paleostress fields using idealized fault populations

for different deformation zones. As part of this inversion method, I tested the frictional

strength of the modeled fault populations to determine the relative role of fluid pressure

or other strain-induced weakness. I describe the methods and results of this technique in

Chapter III.

I also analyzed approximately 20 thin sections of leucogabbros, mylonites, and

cataclasites using a petrographic microscope made from hand samples collected from

field sites to supplement my structural observations. The mineral assemblages in thin

sections helped to determine the composition of fluids, if present, and the conditions of

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deformation at each field site. The primary results of the microscopy were mineralogy

and microstructural observations.

3. Thesis outline

I have organized this thesis into four chapters. Chapter I (this section) serves as an

introduction to the significance of the project and describes the main results of the thesis.

This chapter also provides an overview of the methodology I used for this study. Chapter

II is a literature review of the geology and tectonic history of the Fiordland region of New

Zealand. Chapter II also contains information on the role of fluids in deformation

processes, and on the development of fault-slip analysis techniques described in Chapter

I. Chapter III serves as the main body of this thesis, and I have structured this chapter in

the form of a manuscript intended for submission to the Geophysical Journal

International on the nature of processes controlling deformation and strain localization in

the upper crust adjacent to a tectonically active transpressional plate boundary. Because

Chapter III is intended for submission and must function independently from the

remainder of the thesis, there is some overlap in content with other chapters in the thesis.

Chapter IV is a synthesis of the main results from Chapter III and places the results in a

broader context. This section also outlines possible future work that could provide

additional depth to the findings of this study.

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Chapter II: Literature Review

1. Overview of the geology and the tectonic history of Fiordland, New Zealand

1.1 Tectonic & geologic history of Fiordland and the Australia-Pacific plate boundary

The Fiordland region of New Zealand is located on the southwest coast of the South

Island, and is composed of a variety of accreted terranes that record Paleozoic and

Mesozoic convergence (Fig. 2.1). The terranes in Fiordland are part of two major

provinces, the Western and Eastern Provinces (Bishop et al., 1985), that are on either side

of the Median Batholith (Mortimer et al., 1999a; Mortimer et al., 1999b). In the Eastern

Province, the Brook Street and Caples terranes are composed of metamorphosed

sandstones and mudstones that represent sediments and arc material from the Pacific

margin of Gondwana (Bishop et al., 1985; Mortimer et al., 1999a). The Western

Province terranes contain Early Paleozoic metasedimentary rocks intruded by Devonian-

Carboniferous plutons (Bishop et al., 1985; Mortimer et al., 1999a; Mortimer, 1999b).

In central and northern Fiordland, the Median Batholith includes the Darran Complex

(Mortimer et al., 1999a). The Median Batholith is composed largely of Middle Triassic

to Early Cretaceous intrusive units with calc-alkaline compositions that are consistent

with subduction-related arc magmatism (Mattinson et al., 1986; Gibson, 1990). The

Darran Complex in northern Fiordland has yielded Early Cretaceous ages (142 – 137 Ma)

determined from U-Pb dating of zircon (Kimbrough et al., 1994). The dominant lithology

in the Darran Range is a medium-grained biotite-rich leucogabbro that contains local rafts

of coarse-grained diorite from the Triassic Mistake Suite (Williams & Harper, 1978;

Mortimer et al., 1999a; Turnbull, 2000). The western margin of the Darran Suite is

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Figure 2.1: a. Tectonic setting of New Zealand and the Australia/Pacific plate boundary. Relative plate motion vector of the Australian plate indicated (arrow) calculated from the NUVEL 1A (DeMets et al., 1994). Location of the Euler pole (filled circles) as it rotated over time (ages and magnetic anomalies in parentheses) shown (Walcott (1998)), as well as the location of the modern instantaneous Euler pole from NUVEL-1A model (open circle). Inset shows location of the Fiordland region. b. Geologic and tectonic map of the Fiordland region of the South Island, New Zealand (after Bradshaw (1990), Norris & Turnbull (1993), and Claypool et al., (2002)).

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altered to hornblende diorite (Turnbull, 2000). Along the northwestern margin of the

Darran Complex, the leucogabbro has been deformed and metamorphosed (Bradshaw,

1990; Muir et al., 1995; Claypool et al., 2002; Marcotte et al., 2005).

In central and western Fiordland, the Cretaceous phase of magmatism is the Western

Fiordland Orthogniess (WFO). The WFO forms part of the intrusive units of the Median

Batholith (Mattinson et al., 1986; Gibson, 1988; Hollis et al., 2004), and was emplaced

between 126 and 116 Ma (Mattinson et al., 1986; Gibson, 1988). In the northern

Fiordland region, WFO at Mt. Daniel yielded dates between 121 and 115 Ma from

zircons (Hollis et al., 2004); several zircons yielded much older ages (366 to 239 Ma),

indicating a possible Paleozoic protolith for the WFO (Hollis et al., 2004). Deformation

of the WFO began soon after it was emplaced (123 – 121 Ma determined from U-Pb

dating of zircons; Hollis et al., 2004). Cretaceous deformation included ductile

deformation and recrystallization between 750º and 850º C at pressures of 10 – 13 kbar.

(Hollis et al., 2004).

Collision between Gondwana and outboard terranes at the Pacific margin changed to

extension in the Late Cretaceous, possibly due to the arrival of a spreading center at the

subduction zone between the continent and the terranes (Weissel et al., 1977; Muir et al.,

1994; Walcott, 1998). Sea-floor spreading between Gondwana and the outboard

continental fragments of Zealandia initiated by 85 – 83 Ma (Wood et al., 2000).

Zealandia is the composite continental fragment that is composed of New Zealand, as

well as the submerged Campbell Plateau, Challenger Plateau, Chatham Rise, and Lord

Howe Rise (Fig. 2.1 A) (Mortimer et al., 2006). Spreading between Australia and New

Zealand was linked to extension between New Zealand and Antarctica by a transform

18

fault along the margin of the Campbell Plateau (Wood et al., 2000). Extension along the

Tasman Sea ridge initiated the opening of the Tasman Sea by 80 Ma (Wood et al., 2000).

Extension associated with rifting led to wide-spread normal faulting in New Zealand in

the Late Cretaceous, and the creation of large sedimentary basins on the continental

margins and within Zealandia (Bishop & Laird, 1976; Tulloch & Kimbrough, 1989;

Bishop, 1992; Laird, 1993; Norris & Turnbull, 1993; Barnes et al., 2005). Spreading

along the Tasman Sea ridge ceased by ~75 Ma (Walcott, 1998; Wood et al., 2000), and

this termination marked a change in the relative motion between the Australia and Pacific

plates (Gaina et al., 1998). Sea-floor spreading continued along the ridge in the South

Tasman Sea, leading to the separation of New Zealand from Antarctica by 45 – 40 Ma

(Weissel et al., 1977; Wood et al., 2000).

The Euler pole of rotation between the Pacific and Australian plates began to migrate

southward at approximately 30 Ma, leading to oblique right-lateral motion along the

Pacific-Australia plate boundary (Lamarche et al., 1997; Wood et al., 2000). As the

Euler pole continued to migrate, motion across the plate boundary became increasingly

oblique until ~20 Ma, when movement along the plate boundary became entirely right-

lateral (Sutherland, 1995a; Walcott, 1998). Further rotation (Walcott, 1998) of the Euler

pole (Fig. 2.1 A) continued as subduction of the Australian plate initiated beneath the

southern margin of New Zealand by approximately 10 Ma (Lamarche et al., 1997).

Oblique convergence along the Australia-Pacific plate boundary began within the past 5

m.y. due to an increase in the obliquity of motion across the plate boundary (Walcott,

1998). The increase in oblique movement led to the onset of uplift and transpressional

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deformation along the plate boundary (Walcott, 1998). The Euler pole has not shifted

significantly since transpression stabilized at ~5 Ma (Sutherland, 1995a; Walcott, 1998).

1.2 The modern Australia-Pacific plate boundary

New Zealand is an exposed fragment of Zealandia, which is the micro-continent that

has been sutured together and uplifted by transpressional movement across the Australia-

Pacific plate boundary. In the south and central regions of the South Island, the plate

boundary is represented by the Alpine Fault Zone, where it is a linear feature that strikes

050º - 055º (Fig. 2.1 A). The current relative plate motion velocity vector at Milford

Sound is 36 ± 3 mm/yr toward a bearing of 067 ±2º (for latitude, longitude: 44º 30’ S,

168º E from NUVEL 1A model of DeMets et al., 1994). This motion may be resolved

into 23 ± 2 mm/yr of Alpine Fault-parallel dextral strike-slip motion, 12 ± 4 mm/yr of

horizontal motion accommodated by clockwise rotation of crustal blocks and oblique

motion, and 5 ± 3 mm/yr of throw on reverse faults at the margins of the plate boundary

(Sutherland et al., 2006).

GPS surveys indicate that 65% - 75% of the fault-parallel plate motion is

accommodated on the Alpine Fault, and more than 60% of the strain along the fault is

concentrated within ~20 km of the fault zone (Norris & Cooper, 2001). Metamorphic

mineral assemblages indicate that amphibolite facies deformation occurred at depths of

15 – 25 km (Grapes & Watanbe, 1992). Ages determined from K-Ar dating of micas in

Alpine schists (Adams, 1981) and from fission track dates of zircon (Tippett & Kamp,

1993) are very young (1 – 5 Ma). The very young ages of the Alpine schists combined

with the mid-crustal depths of metamorphism indicate rapid exhumation of Alpine fault-

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related rocks associated with transpression (Koons et al., 2003). Sutherland et al. (2006)

suggest that uplift on reverse faults may be associated with a crustal detachment fault or

distributed shortening of the lithosphere at depth.

The orientation and sense of motion of the Alpine Fault changes across the South

Island as the geometry of the plate boundary changes. Near Milford Sound on the west

coast, the Alpine Fault is subvertical to very steeply southeast-dipping with strike-slip

striae (Norris & Cooper, 1995; Norris & Cooper, 2001). In the center of the South Island,

the Alpine Fault is composed of dominantly strike-slip faults that are linked by segments

of oblique-thrust movement (Norris et al., 1990; Norris & Cooper, 1995; Little et al.,

2002a). In the northeastern section of the South Island, the surface expression of the

plate boundary is the 100 km wide Marlborough fault zone that contains many fault

traces with predominantly strike-slip motion (Anderson et al., 1993). The Alpine Fault

links the westward-verging subduction of the Pacific plate in the Hikurangi Trench north

of the South Island (Reyners & McGinty, 1999; McGinty et al., 2000; Eberhart-Phillips

& Chadwick, 2002) to the eastward-verging subduction of the Australian plate beneath

southern Fiordland and in the Puysegur Trench off the southwestern coast of the South

Island (Eberhart-Phillips & Reyners, 2001; Reyners et al., 2002).

In the central and southern regions of the South Island, the Alpine Fault is interpreted

to have a narrow and subvertical Benioff zone (Reyners, 1989), and has had little seismic

activity during the past 150 years greater than Mw ≥ 5 (Anderson et al., 1993; Eberhart-

Phillips & Reyners, 2001). The Benioff zone is the region in the Earth’s crust at plate

boundaries that is seismically active and interpreted to represent the geometry of the

boundary region. In the central Fiordland region, ML ≈ 4 have been recorded deeper than

21

~130 km (Reyners et al., 2002). South of Milford Sound, the Alpine Fault continues off-

shore, and seismicity occurs in a broad zone across the branching plate boundary (Moore

et al., 2000; Eberhart-Phillips & Reyners, 2001). Earthquakes along the Australia-Pacific

plate boundary and the Alpine Fault represent the modern stress field in the region (e.g.

Ghisetti, 2000). P-axis azimuths calculated from the focal mechanisms of oblique- and

reverse-sense earthquakes from the region are oriented 50º - 65º from the plate boundary

(Anderson et al., 1993; Moore et al., 2000). These P-axes suggest that the current stress

field in the Fiordland region is relatively homogenous, and that the compressional

component is oriented at a moderately high angle to the Alpine Fault.

1.3 The modern structure of Northern Fiordland

The Milford Sound region of northern Fiordland (Fig. 2.1 B) is composed primarily

of high-grade gneisses, the Median Batholith (shown on Fig. 2.1 B), and

metasedimentary and volcaniclastic rocks within the Brook Street Terrane (Mortimer et

al, 1999b). These units are separated by several large terrane boundary faults and shear

zones (Blattner, 1991; Hill, 1995; Mortimer et al., 1999b; Turnbull, 2000). The Glade-

Darran Fault forms the eastern boundary between the Darran Complex and the Brook

Street Terrane (Blattner, 1991; Sutherland, 1995b; Mortimer, 1999b). The Hollyford

Fault separates the Brook Street Terrane from the Maitai Terrane to the east (Turnbull,

2000). Both of these faults merge in the Hollyford Valley fault zone to the east of the

Darran Complex, and may be active due to aseismic slip (Sutherland, 1995b). The

Harrison-Kaipo Fault zone forms the northwestern margin of the Darran Range to the

22

north of Milford Sound (Fig. 2.1 B), and has Tertiary dextral oblique-reverse motion

(Claypool et al., 2002).

The Harrison-Kaipo Fault zone forms the boundary between the Darran Complex and

the Milford and Harrison gneisses (Claypool et al., 2002). This fault zone contains

several generations of lineations and foliations related to Early Cretaceous shortening

(Daczko et al., 2001) and Tertiary shear zone deformation (Claypool et al., 2002). Upper

greenschist facies Tertiary deformation consists of mylonitic foliations and mineral

lineations, as well as upright isoclinal folds (Claypool et al., 2002).

2. Transpressional tectonics

The term ‘transpression’ describes deformation in zones of oblique convergence

(Harland, 1971). Harland (1971) originally characterized transpression in Norway as a

zone of shearing and stretching between steep, parallel faults that do not show vertical

displacement. Sanderson & Marchini (1984) expanded the definition to include

deformation zones of oblique convergence with lateral confinement and constant volume.

This generalized description allowed Sanderson & Marchini (1984) to model the strain

ellipsoid for both transpression and transtension, as well as make predictions for the

physical effects due to fault bends and terminations. The surface expression of

transpression includes thickening and uplift, and flower structures in the deformation

zone.

Fossen & Tikoff (1993) modified this model to include progressive deformation for

both transtension and transpression. The new model combined pure and simple shearing

to generate different strain paths, as well as a description of the shape of the strain

23

ellipsoid. Fossen & Tikoff (1993) also make predictions for the physical deformation of

passive markers in zones of transpression, which include flattening and the development

of planar fabrics. Tikoff & Teyssier (1994) further refined the transpression model to

consider the role of frictional strength of plate boundary faults and the relative orientation

of the plate motion vector. By applying their model to Sumatra and central California,

Tikoff & Teyssier (1994) were able to test the role of strain partitioning and the

orientation of the principle stress axes. Tikoff & Teyssier (1994) characterized distinct

plate boundary conditions and related the transcurrent and contractional components of

motion to deformation style.

Teyssier et al. (1995) again expanded the previous model to include parameters such

as the degree of kinematic partitioning, the orientation of instantaneous strain axes, and

the relative direction of plate motion along the San Andreas Fault in California and the

Alpine Fault in New Zealand. Teyssier et al. (1995) were able to quantify a relationship

between the degree of strike-slip partitioning and the orientation of the instantaneous

strain axes relative to the direction of plate motion. In transpressional regimes, none of

the instantaneous strain axes is parallel to the direction of plate motion, but the axes are

instead rotated toward the translation component of motion (Sanderson & Marchini,

1984; Teyssier et al., 1995; Fossen & Tikoff, 1998).

Recent expansion of the transpressional model has worked to overcome limitations

due to the original specific boundary conditions (Sanderson & Marchini, 1984).

Transpression has been expanded to include not only vertical extrusion, but also the

lateral extrusion of material (Jones et al., 1997). This expansion accounts transpression

zones that are not strictly wrench- or pure shear-dominated (Fossen & Tikoff, 1993), and

24

also allows for a wider variety of orientations in fabrics in a transpressional zone (Czeck

& Hudleston, 2004). New models also include inclined transpressional zones (Jones et

al., 2004) that describe such features as triclinic strain and asymmetrical structures, the

development of foliation that is not parallel to the transpression boundaries, and the

partitioning of strain into strike-slip and down-dip components (Jones et al., 2004).

Many plate boundaries are now described as transpressional (Bunds, 2001; Paterson

et al., 2002; Malservisi et al., 2003; West & Roden-Tice, 2003; Cunningham, 2005), but

the phrase is commonly used to describe obliquely convergent plate motion rather than

strictly a deformation style. However, the Alpine Fault in New Zealand is an example of

a plate boundary that is used to typify transpression (Norris & Cooper, 1995; Teyssier et

al., 1995; Koons et al., 2003) due to the obliquely convergent plate motion between the

Australian and Pacific plates, the steep plate boundary (Reyners, 1989), and uplift

(Gerbault et al., 2002; Malservisi et al., 2003).

3. The role of fluid in upper crustal faulting

Pore fluid pressure at various levels in the crust is the result of water released

from pore spaces in sediments buried and lithified, or from dehydration during

metamorphism (Hubbert & Rubey, 1959; Stern et al., 2001). As fluids move through the

crust, they tend to weaken the response of the host material to imposed stresses (Hubbert

& Rubey, 1959; Zoback et al., 1987; Byerlee, 1990; Byerlee, 1992; Rutter et al., 2001).

Consequently, fluids play an important role in localizing strain in deformation zones, and

in allowing faults to slip that are otherwise not optimally oriented to accommodate stress

(Hubbert & Rubey, 1959; Sibson, 1985; Byerlee, 1992). In this section, I review of the

25

role of fluids in frictional weakening, and then describe several significant examples of

crustal-scale transform faults weakened due to elevated pore fluid pressure.

3.1 Pore fluid pressure

Elevated pore fluid pressure influences the amount of shear stress required for the

formation and failure of faults in the crust. Coulomb’s (1776) law of failure describes the

relationship between the critical stress required for brittle failure (σc), the cohesive

strength of a material (σo), the angle of internal friction (ϕ), and the normal stress (σN) on

a fault as:

σc = σo + tan ϕ • (σN). (2.1)

Failure occurs according to Coulomb’s rule as long as the normal force is the only

compressive force acting on the material. Hubbert & Rubey (1959) first modified this

equation to quantify the effects of pore fluid pressure in a system by changing the

compressive force to include fluid pressure, (Pf):

σc = σo + tan ϕ • (σN - Pf). (2.2)

Subtracting pore fluid pressure from the normal stress (σN - Pf) yields a new term, σ∗,

which is the effective stress on the fault (Hubbert & Rubey, 1959):

σc = σo + tan ϕ • (σ∗N), (2.3)

where σ∗N is the effective normal stress. The effective stress term allows for significant

variation in the critical stress a particular fault may require for failure. Depending on the

fluid pressure, a fault could experience an effective confining stress as high as lithostatic

(i.e. no fluid pressure) to almost no confining stress (i.e. fluid pressure is close to

lithostatic pressure).

26

Hubbert & Rubey (1959) also introduced the fluid pressure ratio (λ) to describe the

relationship between pore fluid pressure and lithostatic pressure (Pl):

λ = (Pf) / (Pl) = (Pf) / (ρr g h), (2.4)

where ρr is the density of the overlying rock type, g is the acceleration due to gravity, and

h is the height of the column of rock above the fault. The fluid pressure ratio ranges from

λ = 0.37 to 0.47 in hydrostatic conditions with non-elevated pore fluid pressure (Suppe,

1985), up to λ = 0.50 to 0.90 for abnormal or elevated pore fluid pressure (Suppe, 1985).

At abnormal values of pore fluid pressure, the effective confining stress for faults can be

extremely low.

In most intracontinental and nonorogenic upper crustal conditions, pore fluid pressure

is thought to be hydrostatic (Hubbert & Rubey, 1959; Byerlee, 1990). Using the KTB

deep drill hole in Germany, Grawinkel & Stöchkert (1997) measured the expected

hydrostatic pore fluid conditions in the crust up to 9 km in depth in a nonorogenic setting.

However, Simpson (2001) used numerical models to demonstrate that elevated pore fluid

pressure can reduce the strength of rocks in the upper crust by as much as 60% in

compressional environments. Byerlee (1992) showed that the pore fluid pressure may be

as high as 85% of lithostatic pressure on some segments of the San Andreas Fault in

central California. Additionally, Stanislavsky & Garven (2002) modeled the failure of

thrust faults due to pore fluid pressure elevated to near lithostatic values at depths >3 km.

3.2 Fluid infiltration and frictional weakening of crustal-scale faults

Weak crustal-scale strike-slip faults may play an important role in focusing strain in

an otherwise strong upper crust (Byerlee, 1990; Grawinkel & Stöckhert, 1997).

27

Intraplate upper crustal rocks are strong and slip on optimally oriented faults with a

coefficient of friction (µ) of µ = 0.6 – 0.7 (Brudy et al., 1997; Grawinkel & Stöckhert,

1997). The coefficient of friction is equal to the tangent of the angle of internal friction

(ϕ) (Eq. 2.1). If interplate strike-slip faults are weak (µ = 0.1 – 0.2), movement may be

preferentially focused on these large faults (Provost & Chéry, 2006). Weak crustal-scale

faults may facilitate continued slip on interplate transform faults, even if they are not

optimally oriented for slip in a stress field.

The San Andreas Fault in central California is an example of a crustal-scale strike-

slip fault that is described as extremely weak with fluid pressure ratio values of λ = 0.85

or higher (Zoback et al., 1987; Byerlee, 1992; Zoback & Healy, 1992). The orientation

of maximum horizontal compression is at very high angles (~85º) to the main fault

(Mount & Suppe, 1987; Zoback et al., 1987; Provost & Houston, 2001; Townend &

Zoback, 2004), and there is low heat flow associated with the fault (Lachenbruch & Sass,

1980; Lachenbruch & McGarr, 1990; Lachenbruch & Sass, 1992). Both of these features

are interpreted to indicate the weak nature of the fault (Zoback et al., 1987; Byerlee,

1992; Provost & Houston, 2001; Hickman & Zoback, 2004; Townend & Zoback, 2004).

Both frictional weakening due to fault gouge minerals (Townend & Zoback, 2001;

Holdsworth, 2004) and increased pore fluid pressure (Zoback et al., 1987; Byerlee, 1990;

Byerlee, 1992; Rice, 1992) have been invoked to explain the weakness of the fault.

While there has been some reinterpretation of the data used to describe the San Andreas

Fault as weak (Scholz, 2000a; Scholz, 2000b), the fault is still widely interpreted to be

weak at least partially due to elevated pore fluid pressure (Zoback, 2000; Holdsworth,

2004; Provost & Chéry, 2006).

28

There are several other crustal-scale faults that are described as “weakened” due to

elevated pore fluid pressure. The Great Glen fault in Scotland is interpreted as a fluid-

weakened fault (Stewart et al., 2000) due to foliated cataclastic rocks and hydrous

mineral phases in the fault zone. The Castle Mountain strike-slip fault in Alaska is also

an example of a crustal-scale fault that has evidence of progressive weakening due to

elevated pore fluid pressure and the growth of a clay-rich fault gouge (Bunds, 2001).

Additionally, Srivastava & Sahay (2003) recently identified the Great Boundary fault in

northwestern India as a fault that has likely been reactivated as a thrust fault several times

due to elevated pore fluid pressure. The fault zone shows pervasive fluid inclusions,

which Srivastava & Sahay (2003) interpret as fluid-assisted weakening during

reactivation.

Recent work on pore fluids and seismic-wave behavior beneath the Southern Alps in

New Zealand (Koons et al., 1998; Stern et al., 2001) indicates that the Alpine Fault may

be another active weak plate boundary fault. Liu & Bird (2002) used modeling of faults

in central New Zealand to calculate an extremely low coefficient of friction of µ = 0.17.

Paleostress and shear-wave splitting work (Balfour et al., 2005) indicate a fluid pressure

ratio of λ = 0.7 for regions near the Alpine Fault. These values are similar to those from

the central region of the San Andreas Fault, and suggest that pore fluid pressure may have

an effect on the frictional strength of the Alpine Fault Zone.

4. The application of fault-slip data to tectonic settings

The use of fault-slip data to describe a stress tensor circumvents a common problem

in structural geology: it is not possible to directly measure the orientation and magnitude

29

of a stress field acting upon the Earth’s crust (Twiss & Unruh, 1998). Instead, one may

measure deformation features, such as fractures and faults or borehole break-outs created

within a stress field, and then calculate the orientation of the stress axes from the fault

data. Stress inversion of fault-slip data was first proposed by Wallace (1951), and has

since been modified and expanded by numerous workers including, Bott (1959), Angelier

(1979) and many coworkers (see Angelier (1994) for a detailed summary), Michael

(1984), Célérier (1988), Gephart (1990), and Zoback (1989; 1992). Gephart & Forsyth

(1984) and Michael (1987) have also developed stress inversion methods using

earthquake focal mechanisms, and the mechanism of Abers & Gephart (2001) relies on

first motion data from earthquakes. I do not address earthquake data specifically because

the methods do not rely on fault-slip data, and are therefore not directly comparable to the

kinematic analysis of fault-slip data. Focal mechanism inversion is a useful method (e.g.

Balfour, 2005), but is not applicable to this work due to its focus on fault-slip data.

The direct inversion method described by Angelier (1979) is a common technique and

is based on the graphical inversion method of Angelier & Mechler (1977). Graphical

inversion relies on the assumption that all faults in a population moved independently due

to one maximum stress direction. Graphical inversion involves calculating and plotting

the orientation of the maximum and minimum stress axes for each fault and averages

their orientation (Angelier & Mechler, 1977). However, this method does not uniquely

constrain the orientation or magnitude of the stress axes. Angelier & Goguel (1978)

modified the graphical inversion method to calculate the orientation of the stress axes via

direct inversion. Direct inversion relies on a least-squares minimization of the tangential

stress perpendicular to the measured slickenline (Angelier, 1979) to determine the

30

orientation of the principal stress axes. Célérier (1988) used the modified Monte Carlo

search technique of Etchecopar et al. (1981) to further refine the stress inversion method

described by McKenzie (1969) and Angelier (1979). Célérier (1988) also added a

frictional constraint to this technique to restrict possible stress tensors by considering the

effect of sliding friction.

Direct stress inversion methods yield a reduced stress tensor composed of 4 values:

the orientation of the three principle stresses (σ1, σ2, σ3), and the ratio of their

magnitudes, δ (Angelier, 1975). The ratio has values of 0 ≤ δ ≤ 1, and is described by the

following:

δ = (σ1 – σ2) / (σ1 – σ3). (2.5)

For this term, δ = 0 represents an oblate stress ellipsoid, and δ = 1 represents a prolate

stress ellipsoid. It is not possible to calculate the absolute magnitude of the principle

stresses based on direction inversion methods, only their relative ratio. However,

Célérier (1988) indicates that a wide range of fault planes requires a high relative value of

σ1 to activate the variety of orientations than do clustered fault-slip data.

Each stress inversion method relies on independent mathematical techniques, but all

of the inversion methods rely on several similar assumptions about the nature of the

material and the stress tensor being modeled. The first assumption is that the material in

question is homogenous, and that the material responds homogenously to the applied

stress. Secondly, these methods assume that the direction of resolved shear stress on a

fault plane is parallel to the direction of applied stress. Stress inversion methods also

assume that all faults used correspond to a single tectonic event. Finally, stress inversion

methods require the assumption that only faults in a range of optimal angles (βopt) will be

31

activated (βopt = 22.5º - 30º) (Wallace, 1951; Angelier, 1979; Sibson, 1985; Célérier,

1988). The validity of these assumptions is widely invoked, and is not frequently

addressed. Twiss & Unruh (1998) and Gapais et al (2000) summarize and discuss these

assumptions, and express significant concerns about their reliability.

Lisle & Srivastava (2004) attempted to test two assumptions made by stress inversion

methods. The first assumption is that slip occurs parallel to the direction of resolved

shear stress on a plane of preëxisting weakness (Wallace, 1951; Bott, 1959). The second

assumption is that only faults in the range of optimal orientations with respect to the

stress field will be activated. By comparing a survey of published fault-slip data to the

predicted orientations of fault-slip data, Lisle & Srivastava (2004) showed that fault striae

from natural data are consistent with predicted striae for a stress tensor, in good

agreement with the first assumption. The study also showed that the magnitude of

friction on a fault plane controls the activity of a fault, also in good agreement with the

second assumption.

The kinematic analysis of fault-slip data requires several assumptions about strain.

Similar to the assumption made by stress inversion methods, kinematic analysis involves

the assumption that the direction of motion preserved on a fault plane is parallel to the

slip vector. Several kinematic analysis techniques, such as the method used by the

program FaultKin (Allmendinger et al., 1994; modified 2006), also require that the

maximum and minimum instantaneous strain axes lie in a plane that is perpendicular to

the fault plane, and that these axes are both 45º from the fault plane. However, the

second requirement does not involve any interpretation of the data (Marrett &

32

Allmendinger, 1990), and is essentially the calculation of a fault plane solution from the

fault-slip data.

Kinematic analysis uses sense of displacement on a fault surface to calculate the

orientation of instantaneous strain axes (Marrett & Allmendinger, 1990; Twiss & Unruh,

1998). Displacement on a fault surface is one way a material accommodates incremental

deformation. This accommodation represents the straining of a material. Therefore,

kinematic analysis uses incremental strain markers (fault-slip data) to model

instantaneous strain. Using incremental strain to describe instantaneous strain requires

fewer assumptions than using incremental strain to describe stress. For instance, the

assumption of stress inversion methods that a fault will not fail if it is not in the optimal

orientation is only valid for hydrostatic conditions (Byerlee, 1992). This is not a useful

assumption for faults in regions of elevated pore fluid pressure (see Sections 3.1 and 3.2).

The models and solutions provided by stress inversion and kinematic analysis contain

different data, and are useful for describing stress and strain on different scales.

Kinematic analysis of fault-slip data provides strain models that rely on fewer

assumptions than stress inversion methods, but the results also provide simpler models.

Kinematic analysis is best applied to modeling the local strain rate, and is therefore most

useful for describing local features and processes (Twiss & Unruh, 1998). Using fault-

slip data to model stress through direct inversion methods relies on the additional

assumption that the modeled material responds to stress in an isotropic or homogenous

manner. Stress inversion results have the potential to describe regional stress fields, but

are hampered by assumptions about the Earth’s crust (Twiss & Unruh, 1998). While

Lisle & Srivastava (2004) have shown that assumptions about reactivation potential and

33

the orientation of resolved shear stress are robust, they do not test assumptions about the

strength and isotropy of the crust.

For this study, using the stress inversion program, Fault Slip Analysis (FSA) provided

by Célérier (2006) was most useful for describing the frictional conditions required to

activate segments of major faults. These frictional experiments provided results that are

distinct from any results possible using kinematic analysis, and were useful for describing

the relatively weak nature of the crust near the Alpine Fault. Results provided by FSA

experiments using the geometric constraint alone provide models similar to those

generated by the kinematic analysis, but are subject to crustal isotropy assumptions.

Stress inversion results derived from earthquake focal mechanisms such as those of

Gephart & Forsyth (1984) and Michael (1987), or from earthquake first motion data

(Abers & Gephart, 2001), may provide more robust results because they do notrely on the

assumptions related to using fault-slip data.

The results of kinematic analysis using FaultKin v. 4.3.5 (Allmendinger et al., 1994;

modified, 2006) provide a local result, and therefore, a more limited model than those of

stress inversion methods. However, this method of kinematic analysis relies on fewer

assumptions than direct inversion, and uses strain measurements to model average

displacement patterns in the upper crust. Additionally, the technical benefits of this

method include the ability to weight the significance of each fault-slip datum, and grade

the reliability of the fault-slip data. Many direct inversion methods using fault-slip data,

including FSA, do not have this flexibility.

34

Chapter III: Strain localization in the upper crust adjacent to the tectonically active Alpine Fault Zone in Fiordland, New Zealand

Abstract Structural observations and analysis of fault-slip data from a ~800 km2 region of the Darran Range in northern Fiordland, New Zealand, reveal spatial variations in strain localization and the occurrence of strike-slip partitioning adjacent to the Alpine Fault. Geometrical and frictional constraints on the analysis of stress inversion results from fault-slip data indicate that the coefficient of friction is extremely low (µ = 0.10) for major fault segments within ~10 km of the Australia-Pacific plate boundary. Compression axes are oriented ~60º from the dominantly northeasterly strike of the Alpine Fault. The large angle between the compression axes and the plate boundary, combined with the low coefficient of friction, suggest a weakening of the crust around the Alpine Fault Zone. Deformation within 10 km of the plate boundary is characterized by reverse, oblique-reverse, and strike-slip fault populations. Away from the plate boundary, deformation is characterized by oblique-reverse and strike-slip motion on reactivated steep faults; vertical motion is predominantly localized along lithologic boundaries. This deformation results in the extrusion of wedge-shaped blocks that make up the Darran Range. Cross-cutting relationships and kinematic analysis indicate a superposition of distinct stress fields in northern Fiordland, including an older phase of normal faulting from the Late Cretaceous – Early Tertiary. We suggest that the dominant strain localization mechanism in the upper crust in northern Fiordland is elevated pore fluid pressure in the near-boundary deformation zone. Manuscript in preparation for submission to Geophysical Journal International

35

1. Introduction

The nature of processes controlling the degree and style of strain localization at

obliquely convergent plate boundaries is an unresolved problem in continental tectonics.

Strain localization due to strain-induced weakening mechanisms may control elements of

upper crustal deformation such as fault strength (Rutter et al., 2001; Buck & Lavier,

2001; Liu & Bird, 2002; Balfour et al., 2005). It may also contribute to a partitioning of

displacements within the crust and influence the mechanisms by which deeply buried

rocks are uplifted and exhumed (Norris & Cooper, 1995; Claypool et al., 2002; Little et

al., 2002a; Little et al., 2002b; Koons et al., 2003). In this chapter, I present an analysis

of fault patterns combined with strain modeling of fault-slip data from zones of

continental collision that provide important insights into the processes that influence

strain localization, strain partitioning, and the evolution of stress fields in New Zealand.

Transpression associated with the Australia-Pacific plate boundary on the South

Island of New Zealand (Fig. 3.1 A) has resulted in the uplift of the Southern Alps

(Koons, 1987; Norris et al., 1990; Simpson et al., 1994; House et al., 2002; Koons et al.,

2003; Little et al., 2005) and a region of active faulting up to 100 km wide (Sutherland,

1994; Sutherland & Norris, 1995; Norris & Cooper, 2001; Sutherland et al., 2006).

Along the southern end of the Alpine Fault in northern Fiordland, the fault zone is

characterized by strike-slip motion and uplift (Sutherland & Norris, 1995; Norris &

Cooper, 2001; Sutherland, et al., 2006), and earthquake activity (Anderson et al., 1993;

Doser et al., 1999; Moore et al., 2000; Eberhart-Phillips & Reyners, 2001; Leitner et al.,

2001). Near Milford Sound in Fiordland, there are many terrane boundary faults

inherited from Mesozoic convergence (Blattner, 1991; Mortimer et al., 1999a; Mortimer

36

et al., 1999b) that may be reactivated as splays off of the Alpine Fault. The inherited

lithologic boundaries and deformation structures (Norris et al., 1990; Mortimer et al.,

1999b; Sutherland et al., 2000; Claypool et al., 2002; Marcotte et al., 2005) may localize

strain in northern Fiordland.

In this paper, we describe kinematic and field-based structural data that indicate the

extent of strain partitioning in the shallow crust from the ~20 x 40 km Darran Range in

northern Fiordland (Fig. 3.1 B). Fault-slip data from 11 sites in the region, combined

with kinematic data and stress inversions, show different deformation styles at increasing

distance from the plate boundary. We document a near-boundary deformation zone that

is confined with within ~10 km of the plate boundary, and is structurally and

kinematically distinct from regions farther to the southeast of the Alpine Fault. The uplift

and deformation styles that are present in northern Fiordland indicate that the Darran

Range is dissected by numerous strike-slip and oblique-slip faults that result in the

vertical extrusion of fault-bound wedges. Finally, consistent cross-cutting relationships

between fault populations indicate the superposition of paleostress fields in the Darran

Range, including the presence of an early extensional regime that most likely reflects

both Late Cretaceous and early Tertiary normal faulting.

37

Figure 3.1: a. Tectonic setting of New Zealand and the Australia/Pacific plate boundary. Relative plate motion vector of the Australian plate indicated (arrow) calculated from the NUVEL 1A (DeMets et al., 1994). The location of the Euler pole (filled circles) as it migrated over time (ages and magnetic anomalies in parentheses) shown (after Walcott (1998)), as well as the location of the modern instantaneous Euler pole from NUVEL-1A model (open circle). Inset shows location of the Fiordland region. b. Geologic and tectonic map of the Fiordland region of the South Island, New Zealand (after Bradshaw (1990), Norris & Turnbull (1993), and Claypool et al., (2002)).

38

2. Geologic history and evolution of the Australia-Pacific plate boundary

2.1 Geology and tectonic history of Fiordland & the Australia-Pacific plate boundary

The Mesozoic and Tertiary tectonic history of northern Fiordland includes the Early

Cretaceous collision of successive terranes onto the southeastern margin of Gondwana

(Howell, 1980; Mackinnon, 1983; Bradshaw, 1989; Gibson, 1990; Mortimer et al.,

1999b). Associated with terrane accretion was the intrusion of the Median Batholith,

including the Darran complex, with U-Pb determined ages from zircons of 142 – 137 Ma

(Kimbrough et al., 1994; Mortimer et al., 1999a). Collision at the Gondwana margin

changed to extension in the Late Cretaceous, possibly due to the arrival of a spreading

center at the subduction zone between the continent and outboard terranes (Muir et al.,

1994), and sea-floor spreading began 85 – 83 Ma (Weissel et al., 1977; Walcott, 1998).

Extension led to the separation of the continental fragments of Zealandia from Gondwana

by approximately 80 Ma, and to the opening of the Tasman Sea (Wood et al., 2000).

Extension associated with rifting led to wide-spread normal faulting in New Zealand

in the Late Cretaceous (Bishop & Laird, 1976; Tulloch & Kimbrough, 1989; Bishop,

1992; Norris & Turnbull, 1993). Spreading along the Tasman Sea ridge ceased by ~75

Ma, and this termination marked the onset of a change in plate motion (Gaina et al.,

1998). Seafloor spreading led to separation between New Zealand and Antarctica by 45

– 40 Ma (Weissel et al., 1977; Wood et al., 2000). The Euler pole of rotation between the

Pacific and Australian plates began to migrate to the south at approximately 30 Ma,

leading to oblique right-lateral motion across the Pacific-Australia plate boundary

(Lamarche et al., 1997; Wood et al., 2000). As the pole continued to migrate, motion

across the plate boundary became more oblique until ~20 Ma, when motion along the

39

plate boundary became entirely right-lateral (Sutherland, 1995; Walcott, 1998). Further

southward rotation of the Euler pole (Fig. 3.1 A) continued as subduction initiated

beneath the southern margin of New Zealand by approximately 10 Ma (Lamarche et al.,

1997). Oblique convergence initiated along the Australia-Pacific plate boundary within

the past 5 m.y. due to an increase in obliquity of motion across the boundary, leading to

the onset of transpression and uplift (Walcott, 1998); plate motion has not shifted

significantly since transpressional motion stabilized at ~ 5 Ma (Sutherland, 1995a;

Walcott, 1998).

2.2 Modern tectonic setting of the Alpine Fault

The modern relative plate motion velocity vector at Milford Sound is 36 ± 3 mm/yr

toward a bearing of 067 ± 2º (from NUVEL 1A model of DeMets et al., 1994). This

motion may be resolved into 23 ± 2 mm/yr of Alpine Fault-parallel dextral strike-slip

movement, 12 ± 4 mm/yr of horizontal motion accommodated by clockwise rotation of

crustal blocks and oblique motion, and 5 ± 3 mm/yr of throw on reverse faults at the

margins of the plate boundary (Sutherland et al., 2006). The uplift on reverse faults may

be associated with crustal detachment or distributed shortening of the lithosphere

(Sutherland et al., 2006). These motion and uplift data are consistent with rates reported

by Bishop (1991), Sutherland & Norris (1995), and Norris & Cooper (2001).

Metamorphic mineral assemblages indicate that amphibolite facies deformation occurred

at depths of 15 – 25 km (Grapes & Watanbe, 1992). Ages determined from K-Ar dating

of micas in Alpine schists (Adams, 1981) and from fission track dates of zircon (Tippett

& Kamp, 1993) are very young (c. 1 – 5 Ma). The very young ages of the Alpine schists

40

combined with the mid-crustal depths of metamorphism indicate rapid exhumation of

Alpine fault-related rocks associated with transpression (Koons et al., 2003).

The plate boundary comes on shore to the south of Milford Sound where it is the

Alpine Fault. The fault is very linear in this region and strikes 050º - 055º, is subvertical

to very steeply southeastward dipping, and has primarily strike-slip striations (Norris &

Cooper, 1995; Norris & Cooper, 2001). In the center of the South Island, the Alpine

Fault appears to be a linear feature, but is actually composed of dominantly strike-slip

faults linked by segments of oblique thrust motion (Norris et al., 1990; Norris & Cooper,

1995; Little et al., 2002b). The fault links the westward-verging subduction of the Pacific

plate in the Hikurangi Trench to the north to the eastward-verging subduction of the

Australian plate in the Puysegur Trench to the south (Fig. 3.1 A). There have been at

least 460 km of offset along the Alpine Fault in the past 45 m.y. (Wellman, 1953;

Sutherland, 1999), and approximately 70 km of shortening near Milford Sound

(Sutherland et al., 2000). More than 60% of the strain along the Alpine Fault is

concentrated within 20 km of the fault (Norris & Cooper, 2001).

Seismic activity in the South Island of New Zealand reflects the different geometries

of the plate boundary. In the north, the Benioff zone dips to the west, parallel to the

subduction direction of the Pacific plate. The plate boundary is the 100 km wide

Marlborough fault zone at the northern end of the South Island, and has had several large

(Mw ≥ 6) earthquakes in the past 150 years (Gledhill et al., 2000; Leitner et al., 2001).

The central region of the South Island along the Alpine Fault is interpreted to have a

narrow and subvertical Benioff zone (Reyners, 1989), and has had relatively little seismic

activity during the past 150 years greater than Mw ≥ 5 (Anderson et al., 1993; Eberhart-

41

Phillips, 1995). South of Milford Sound, the Alpine Fault continues off-shore and

subduction of the Australian plate begins by the Puysegur Trench. Seismicity occurs in a

broad zone across the branching plate boundary to the south of Milford Sound (Moore et

al., 2000; Eberhart-Phillips et al., 2001).

Recent earthquakes in Fiordland and along the Australia-Pacific plate boundary can

be interpreted to represent the modern stress field (e.g. Ghisetti, 2000). P-axis azimuths

calculated from focal mechanisms of oblique and thrust earthquakes from Fiordland and

the northern Puysegur trench are oriented at approximately 50º - 65º to the plate boundary

(Anderson et al., 1993; Moore et al., 2000). These P-axes suggest that the modern stress

field is relatively homogenous in the region near Milford Sound, and that the contraction

component is at a moderately high angle to the Alpine Fault.

3. Structure of field regions in Fiordland

3.1 Structure of the Darran Range and Northern Fiordland

Northern Fiordland is composed primarily of high-grade gneisses, the Median

Batholith (partially composed of the Mistake and Darran Suite gabbros, diorites, and

granites), and volcaniclastic sediments of the Brook Street Terrane (Mortimer et al.,

1999b). These lithologies are the metasedimentary rocks associated with the collision of

terranes with Gondwana during the Jurassic and Early Cretaceous, and the corresponding

intrusive and volcanic units.

42

3.1a The Hollyford Valley Fault Zone

The Hollyford Valley forms the eastern margin of the Darran Range, and is a major

topographic low with an average elevation of 20 – 40 m (relief is 2500 m between the

valley floor and Mt. Madeline, 5 km away) that contains three large faults (Fig. 3.2 A).

The Glade-Darran Fault is the western-most fault in the Hollyford Fault Valley, is

approximately 45 km long, and is cut by the Alpine Fault on the northern end, and by the

Hollyford Fault at the southern end (Fig. 3.2 B) (Turnbull, 2000). The Hollyford Fault is

at least ~55 km long, depending on the location of the intersection of the Hollyford Fault

with the Te Anau Fault to its south; the Hollyford Fault and the Glade-Darran Fault are

possibly active due to aseismic slip in northern Fiordland (Sutherland, 1995b).

Geomorphologic evidence such as prominent topographic lineaments and sag ponds in

the Hollyford Valley fault zone suggest Quaternary movement. Finally, the eastern-most

Livingstone Fault is over 100 km long and has evidence for Middle Miocene

displacement along the fault (Turnbull, 2000). Sense of motion on faults in the Hollyford

Valley is uncertain; early work suggested a small sinistral component of motion due to

fault-fold relationships (Sutherland, 1995b), but we present new fault-slip evidence for

dextral and reverse motion on faults within the valley.

Until recently, there has been little exposure of these faults within the Hollyford

Valley. However, there is now excellent exposure of a large splay of the fault zone along

a roadcut on the Milford-Te Anau Road. The Hollyford Fault roadcut (HF-05; Fig. 3.2

B) is within the Eglinton subgroup of the Mesozoic Brook Street Terrane; the outcrop is

in the Consolation Formation, and is composed of bedded volcaniclastic sandstones with

extensive Kaka siltstone horizons (Turnbull, 2000).

43

Fig. 3.2: Topography, sitemap, and cross-section of the Darran Range northern Fiordland. a. DEM of Milford Sound region (derived from NZMS 260 data) shows topographic differences in the Milford Sound area. Area shown similar to that in part b). b. Geologic map of Milford Sound area of northern Fiordland. Geology is based on mapping for this study and on Turnbull (2000), Claypool et al. (2002), and Marcotte et al (2005). c. Cross-section between A - B (no vertical exaggeration) showing the orientation of faults and contacts from data collected for this study and from Claypool et al. (2002).

44

The majority of the faults present at HF-05 are within two dominantly dextral strike-

slip populations that are distinguishable based primarily on fault plane orientation.

Figure 3.3 A shows a photograph and schematic profile of the main structures from the

roadcut. The dominant fault within the outcrop (Fig. 3.3 A) is characteristic of the first of

the two strike-slip fault populations. This population strikes north-south and dips steeply

to the west (e.g. 004:81W); motion along these faults is dextral strike-slip with a

component of shortening across the faults, as evidenced by folds and small thrust faults

adjacent to the main fault (Fig. 3.3 A). The second major fault population strikes

northwest-southeast and dips predominantly to the southwest (e.g. 315:88 S). Faults in

this population have nearly pure dextral strike-slip motion, and are subordinate to the

north-south striking population in terms of trace length.

A third, minor fault population at HF-05 is composed of faults that strike dominantly

east-west and dip moderately to the north and south. These faults all have reverse sense

of motion and often a small component of dextral horizontal offset. This population is

smaller in terms of both trace length and number than either of the strike-slip populations.

Fig. 3.3: Photographs, sketches, and photomicrographs from sites in the Darran Range. a. Photograph of Hollyford Fault (HF-05) road cut, facing southeast. b. Schematic profile of major fault at HF-05. c. Photograph of large normal fault from Lake Truth (LT-06) on the eastern wall of the site, facing southeast. d. Photograph from Gertrude’s Saddle (GS-06) of one of the large steep dextral faults in the southern region, facing southwest. Relief is ~800 m. e. Photomicrograph of thin section from Mount Thunder (MT-05) (crossed polars) showing hydrous phases such as epidote and pumpellyite, and veins. pl = plagioclase; pum = pumpellyite; ep = epidote. Scale bar = 2 mm.

45

46

3.1b The interior of the Darran Range

The interior and eastern margin of the Darran Range are topographically elevated

(Fig. 3.2 A) and have excellent exposure of faults and kinematic indicators. The

dominant lithology in the region is medium-grained biotite-rich leucogabbro within the

Darran Suite in the Mesozoic Median Batholith (Blattner, 1978; Tulloch et al., 1999;

Turnbull, 2000); locally, there are large rafts of coarse-grained Mistake Suite diorite.

Towards the western margin of the Darran Range, the leucogabbro is altered to

hornblende diorite (Turnbull, 2000). There are steep felsic dikes in the interior of the

Darran Range, and several sites, such as Lake Truth (LT-06) and Gertrude’s Saddle (GS-

06), show primary igneous layering.

There are two distinct types of fault populations in the interior of the Darran Range.

The first group of fault populations strikes dominantly northeast and northwest, and dips

both to the northeast and to the southeast (e.g. 055:50S and 320:64N). This population is

composed of normal faults, and there is typically a component of sinistral motion

associated with the down-dip movement on the faults. The normal faults have trace

lengths of up to several hundred meters, and often form in conjugate sets (Fig. 3.3 B),

with most fault surfaces showing epidote staining.

The second type of fault population strikes north-south and dips steeply to the east

and west (e.g. 354:88E); these faults are dextral strike-slip faults, some with a vertical

component of motion. The majority of the major fault surfaces have epidote staining.

These faults are widely spaced (500 – 1000 m apart) and often cut across large valleys in

the interior of the Darran Range (Fig. 3.3 C). The dextral faults consistently cross-cut the

normal fault populations at all the field sites in the interior and at the margins. Several of

47

the steep faults utilize the margins of felsic dikes in the region. The dikes are typically

between 5 and 10 cm wide, and strike north-south and dip moderately to the west.

At the eastern margin of the Darran Range, the site between Madeline Creek and

Catch Creek (MC-06; Fig. 3.2 B) contains an additional, small fault population that is

likely synchronous with the dextral strike-slip faults at the site. The subordinate faults

are part of a population of thrust faults that strike northeast-southwest and dip both

northwest and southeast (e.g. 313:64N). All fault populations cut the primary igneous

layering.

3.1c The northern margin of the Darran Range

Within ~10 km of the Alpine Fault zone, the leucogabbros of the Darran Suite have

been altered, and the orientation and type of fault population differs from those in the

interior and along the eastern margin. The rocks within this zone are highly fractured,

sheared, have pervasive ductile and brittle faults, and there are many pegmatitic felsic

dikes. Mylonitic ductile fabrics are over-printed and reactivated by brittle deformation.

The rocks in the northern region show evidence of pervasive fluid infiltration, and

minerals at these sites include epidote and pumpellyite in the host rock (Fig. 3.3 D).

The dominant fault populations in this region are reverse and dextral oblique-reverse

faults that strike north and northeast and dip moderately to the northwest and southeast

(e.g. 058:57SE). These faults have long traces (50 – 100 m), and well-developed fault

gouge and epidote staining. The secondary fault population present at Mount Thunder

(MT-05; Fig. 3.2 B) is northeast striking and dips steeply in both directions, with a

primarily dextral strike-slip sense of motion.

48

3.1d The Harrison-Kaipo Fault Zone

The Harrison-Kaipo Fault zone forms part of the western margin of the Darran

Complex and the contact zone with Milford and Harrison Gneisses. The region was first

described by Claypool et al. (2002). Field sites within the Harrison-Kaipo Fault zone are

within 10 km of the Alpine Fault, primarily from west of the Darran Range, and within

the Milford and Harrison gneisses in the Arthur River Complex (ARC) (Domain 2; Fig.

3.2 B). There are several generations of lineations and foliations related to Early

Cretaceous shortening (Daczko et al., 2001) and Tertiary deformation in the Milford and

Harrison gneisses (Claypool et al., 2002). Upper greenschist facies Tertiary shear zone

deformation consists of mylonitic foliations and mineral lineations, as well as upright

isoclinal folds. Data from outside of the ARC are from the northwestern part of the

Darran Suite that has been deformed and metamorphosed (Muir et al., 1995) or from part

of the Indecision Creek Complex of Bradshaw (1990) (Domain 1; Fig. 3.2 B). This unit

has also been described as the Selwyn Creek gneiss (Marcotte et al., 2005) due to its

highly deformed nature, distinguishing it from the Darran Suite.

We are reinterpreting some of the structural and kinematic data from the Harrison-

Kaipo Fault zone originally described by Claypool et al. (2002); data from this region are

part of two structural domains within the fault zone. Domain 1 is between the Darran

Complex and Pembroke Fault (Fig. 3.2 B), and is characterized by two steep fault

populations that have compatible, primarily horizontal dextral (e.g. 058:71S) and sinistral

(e.g. 000:80W) strike-slip motion. Domain 2 is located within the ARC and contains

faults that strike north-south and dip variably to the east with dextral oblique-thrust

motion.

49

Mount Daniel is to the south and west of the Harrison-Kaipo Fault zone (Fig. 3.1 B),

and exposes the contact between the Milford Gneiss and the Western Fiordland

Orthogneiss (see section 3.3). There are several normal faults that strike to the northeast

and dip moderately to the west. This population is similar in orientation and sense of

motion to the normal fault populations in the Darran Range and in Doubtful Sound.

3.2 Structure of the Skippers Range

The Skippers Range is located between the Glade-Darran Fault and the Hollyford

Fault (Fig. 3.2 B) at the northeastern margin of the Fiordland region. The dominant

lithology in the central Skippers Range (SR-05; Fig. 3.2 B) is part of the Lone Stag

Formation, composed of serpentinized pyroxenites and gabbros, with remnants of

primary igneous layering (S0). The lithology widely is sheared and altered to greenschist

facies. The igneous layering has been folded into several large, open folds with axes

plunging to the northeast (Fig. 3.4 B, E). In addition to igneous layering, there is a strong

ductile fabric composed of foliation planes, regions of localized shear, and large-scale

folds (Fig. 3.4 B). F1 is a crenulation foliation that contains pressure shadows of chlorite

and epidote surrounding tapered feldspar clasts. F1 is folded (Fig. 3.4 E) and creates an

intersection lineation (L1) with S0. Major shear zones have been mapped by Ballard

(1988) and Turnbull (2000), and represent boundaries with adjoining intrusive units, but

there are smaller shear zones, 30 – 60 m in width, that are within one lithology. Shear

zones strike northeast and dip steeply in both directions (e.g. 050:85S) (Fig. 3.4 B, C).

Shear planes in these zones have pervasive epidote staining, and kinematic indicators

50

such as s-c fabrics and asymmetric tails on porphyroclasts indicate dextral motion (Fig.

3.4 G).

Similar to the northern margin of the Darran Range, the ductile fabrics in the Skippers

Range have been over-printed by younger, brittle deformation. The small shear zones

have been reactivated by brittle faulting, and microstructures such as fractured grains,

cataclasite development, and irregular grain boundaries indicate both ductile and brittle

deformation (Fig. 3.4 F). The dominant brittle fault population strikes northeast and dips

steeply predominantly to the south (e.g. 065:75S; Fig. 3.4 D). There are very few

kinematic indicators for the brittle fault populations in the Skippers Range, but those that

do exist indicate dextral motion on the dominant fault population.

Fig. 3.4: a. Map of Skippers Range showing location of cross-section. Units and symbols same as for 3.2. b. Cross-section of the Skippers Range (no vertical exaggeration). c. Equal-area lower-hemisphere stereoplot of shear planes and mineral lineations from shear zones at center of valley. d. Equal-area lower-hemisphere stereoplot of fault planes and striae for major fault population in region. e. Equal-area lower-hemisphere stereoplot of poles to bedding planes (open triangles) and orientation of the fold axis (open square), and of poles to S1 (closed circles) and orientation of the fold axis (closed square). f. Photomicrograph of sheared quartz vein showing both ductile and brittle deformation. Scale bar is 1 mm. g. Photomicrograph of asymmetric tails composed of pumpellyite and epidote on a clockwise-rotated amphibole grain forming s-type geometry, indicating dextral shearing. Scale bar is 1 mm.

51

52

3.3 Structure of Doubtful Sound

The dominant lithology in Doubtful Sound is the Western Fiordland Orthogneiss

(WFO), which is associated with intrusive units in the Median Batholith (Fig. 3.2 B)

(Gibson, 1988; Hollis et al., 2004). The WFO was emplaced within the Median Batholith

between 126 and 116 Ma (Mattinson et al., 1986; Gibson, 1988). Deformation of the

WFO began soon after emplacement (123 – 121 Ma), determined from U-Pb dating of

zircons (Hollis et al., 2004). Cretaceous deformation of the WFO included ductile

deformation and recrystallization at 750º - 850º C at pressures of 10 – 13 kbar (Hollis et

al., 2004). WFO is in fault contact with ortho- and paragneisses in Doubtful Sound that

also show ductile deformation and recrystallization (Fig. 3.1 B). Brittle fault populations

in Doubtful Sound are primarily normal faults measured along the shoreline of the sound.

The normal faults are moderately dipping and strike to the northeast and the northwest.

4. Kinematics of the Darran Range

4.1 Kinematic methods

The kinematic data that we present in this section are the average fault plane solution

and instantaneous strain axes calculated from fault-slip data collected from the Darran

Range. The fault-slip data consist of fault plane orientation, slip direction, sense of

motion, and, ideally, trace length, and the quality of the sense of motion and striation.

The kinematic technique we use is based on incremental strain summations, similar to the

technique used to calculate infinitesimal strain on fault planes ruptured during seismic

activity. This approach relies on several assumptions about motion on the fault planes.

The first assumption is that the direction of motion on the fault is parallel to the slip

53

vector, represented by the slickenline or other striation; similarly, the technique assumes

that the striae record only the most recent phase of movement. The second principle

assumption is that the instantaneous strain axes lie in a plane parallel to the slip direction

and perpendicular to the fault plane. This requirement also confines the two principle

strain axes to be oriented 45º from the fault plane. Finally, this method assumes that the

strain due to faulting is less than ~60% of the total strain, allowing for the calculation of

the instantaneous strain axes to have error limits consistent with other field data

(Cladouhos & Allmendinger, 1993). This final requirement is reasonable considering

faults used for this work are typically between 1-100 m in trace-length, which is minor

compared to ~70 km of shortening across the Alpine Fault and 460 km of fault-parallel

displacement along the Alpine Fault (Walcott, 1998; Norris & Cooper, 2001).

We calculate instantaneous strain axes, including the principle axes of contraction (Z)

and extension (X), and the average fault plane solution for each fault population at each

field site from the collected fault slip data using the program FaultKin v. 4.3.5 created by

R. W. Allmendinger, R. A. Marrett, & T. Cladouhos (1994; modified 2006). We

determine sense of motion using offset markers, s-c fabric, and chattermarks when

necessary, and we group faults into populations based on fault plane orientation, trend

and plunge of mineral striae (typically epidote in the Darran Range), and sense of motion.

Consistent cross-cutting relationships from observed field data allow us to separate faults

into distinct fault populations that we infer to be related to distinct events. We attempted

to use all fault data collected at each site in order to incorporate the maximum orientation

of faults and therefore justify our assumption that all fault populations are represented.

Faults were only excluded from solutions if the instantaneous strain axes did not group

54

well with any of the fault populations at a site, and were small, had poor control on sense

of motion, or were graded “poor” in terms of certainty of orientation of striation.

Average kinematic solutions from individual instantaneous strain axes are determined

using linked Bingham distribution statistics to calculate the directional maxima of the X

and Z axes (Marrett & Allmendinger, 1990).

The average fault plane solution shows the bulk orientation of fault planes and striae

measured for each fault population (Fig. 3.5). Figure 3.5 C shows the average fault plane

solution for each fault population, and the average instantaneous strain axes provide a

bulk representation for each kinematic solution. These figures show the clustering of the

instantaneous strain axes even though there is heterogeneity in the orientation of the fault

slip data for each population.

4.2 Kinematic results

4.2a The Hollyford Fault Valley and the interior of the Darran Range

The southern region has two distinct fault populations. The first type of fault

population shows dominantly normal motion with a small component of sinistral

movement on moderately dipping fault planes (Fig. 3.5). These fault populations are cut

by dextral and reverse faults. The major fault populations strike northeast and northwest

and form conjugate fault sets that dip to the northeast and the southeast; smaller

populations strike east-west and dip to the north and south. Instantaneous axes of

contraction (Z-axes) are oriented vertically in these populations, and the instantaneous

axes of extension (X-axes) plunge gently towards the west-northwest for the major

55

Fig. 3.5: Detailed map of southern region showing fault slip data and fault-plane solutions for normal fault populations. Solid lines represent lineaments related to normal faults; dashed lines represent strike-slip lineaments. Lithologic units same as for Fig. 3.2.

56

populations, and to the north and south for the smaller populations. Normal faults from

Mount Daniel are north-east striking, northwest-dipping faults, and has an eastern-

trending extension direction, consistent with the large fault populations from the Darran

Range. Normal fault are not present at the Hollyford Fault site (HF-05).

The second variety of fault populations are mutually cross-cutting reverse and steep

dextral strike-slip populations. The dextral strike-slip faults show horizontal motion on

steep faults that generally strike north-south (Fig. 3.6). Some populations show vertical

motion on steep dextral faults (e.g. MC-06 & GS-06; Fig. 3.6). Several steep faults that

show vertical motion have multiple sets of slickenlines or other striae; one orientation is

typically sub-horizontal, whereas the other shows down-dip motion. With the exception

of two subordinate populations at HF-05 (Fig. 3.6), the Z-axes of these fault populations

plunge gently toward the northeast or southwest. The X-axes plunge gently toward the

northwest or southeast for dextral strike-slip populations, with the same exceptions at

HF-05 (Fig. 3.6). It is possible that these subordinate fault populations are related to

interaction between the Glade-Darran and the Hollyford fault, both of which are located

in the Hollyford Fault valley within 500-1000m of each other (Fig. 3.2).

Thrust populations in the Darran Range also show variation in the orientations of the

their contraction axes. The reverse faults from HF-05 show Z-axes in a similar

orientation to those of the subordinate dextral population, possible for the same reason.

The reverse fault populations are from the eastern margin of the Darran Range and

indicate that the Darran Complex may be uplifted.

57

Fig. 3.6: Detailed map of southern region showing fault slip data and fault-plane solutions for dextral strike-slip and reverse fault populations. Units and symbols same as for Fig. 3.2 and 3.5.

4.2b The northern margin of the Darran Range and the Harrison-Kaipo Fault Valley

In the northern region of the Darran Range, fault populations from sites within ~10km

of the plate boundary are dominantly reverse or dextral-oblique reverse faults. Mount

Thunder (MT-05; Fig. 3.7) represents the dominant fault populations: dextral oblique-

reverse, sinistral and dextral strike-slip faults, and sinistral oblique-reverse. Z-axes trend

east-west or northwest-southeast and are nearly horizontal or gently inclined for these

58

populations. These fault populations show dextral strike-slip motion and reverse motion,

which is predicted for a dextral transpressional zone.

Fig. 3.7: a. Detailed map of northern region showing fault slip data (b - i) and fault- plane solutions for strike-slip and reverse fault populations (j - q). Units shown same as in Fig. 3.2 and 3.5.

(following page)

Fig. 3.8: Simplified geologic map of Doubtful Sound region showing fault-slip data and fault-plane solutions for normal fault populations. Map after King (2006).

59

4.2c Doubtful Sound

The dominant brittle fault populations at Doubtful Sound are within the WFO and

other ortho- and paragneisses in the sound. The fault populations show normal and

oblique-normal motion on moderately dipping fault planes (Fig. 3.8). Z-axes are nearly

vertical or inclined to the northeast, and X-axes plunge shallowly to the east or south-

southeast. The variation in the orientation of the X-axes from Doubtful Sound is similar

to the range of variation seen in the normal fault populations near Milford Sound (Fig.

3.8).

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5. Stress inversion

5.1 Methods

We model paleostress fields in the Darran Range using the Fault Slip Analysis (FSA)

program (v. 28.5), developed by Célérier (1988; 2006) after the Monte Carlo search

technique, modified by Etchecopar et al. (1981). The FSA technique, like the majority of

stress inversion programs (e.g. Michael, 1984; Angelier, 1990; Gephart, 1990) relies on

several assumptions. First, that the stress tensor is homogenous for the entire body of

modeled material and that the modeled material behaves isotropically. The second

assumption is that all of the slip on a fault occurs in the same direction as the resolved

shear stress on the fault plane; this direction is given by the orientation of the striation.

Finally, the orientation of the fault planes in each modeled population must be diverse

enough to constrain the stress tensor. Célérier (1988) indicates that the minimum number

of fault slip data per population needed to constrain the stress tensor is 4 faults. In this

study, we use between 5 and 28 faults in each population to constrain the stress tensor.

We use stress inversion solutions in this study to model paleostress fields for

idealized fault populations in each region using geometrical and frictional constraints to

analyze the stress tensors. The geometrical constraint relies on the compatibility between

the orientation of the observed slip on a fault plane and the orientation of the predicted

slip direction for a given generated stress tensor (Angelier, 1979; Etchecopar et al., 1981;

Angelier, 1984; Michael, 1984; Célérier, 1988). The frictional constraint analyses the

ability for a fault population to be activated at given values of the coefficient of friction

(Célérier, 1988; Burg et al., 2005). We use the results of the geometric constraint

analysis to compare to the orientation of the instantaneous strain axes generated by

61

kinematic analysis. Because the kinematic analysis and geometrically constrained

inversion rely on different assumptions, comparing the orientation of the results also

serves as a test of the techniques. The frictional constraint experiment is useful to

compare the relative role of pore fluid pressure, or other strain-induced weakening

mechanisms, between different fault populations, as described below.

FSA generates a reduced stress tensor composed of the orientations of the principle

stresses, and of the relative ratio of the stresses, δ (Angelier, 1975). This ratio has values

of 0 ≤ δ ≤ 1, and represents the shape of the stress ellipsoid; it is defined using the

following relationship:

δ = (σ1-σ2)/( σ1-σ3) (3.1)

where δ=1 represents a prolate stress ellipsoid, and δ=0 represents an oblate stress

ellipsoid (Célérier, 1988). FSA then uses the geometric constraint technique to analyze

the compatibility between generated stress tensors and the fault population. Analysis of

the generated tensors yields an angular misfit value between the orientation of the

observed slip on the fault and the orientation of the predicted slip for each stress tensor

(Etchecopar et al., 1981; Célérier, 1988). The misfit angle between the observations and

the model serves as a measure of quality of the stress tensor with regard to its

compatibility with the measured fault-slip data.

Fault populations used with the geometric constraint in FSA are composed of faults

with the smallest angular misfit values. We have artificially improved the fault

populations by requiring that all modeled faults have angular misfit values of less than

15º (Table 1) and removing outliers from the populations. These misfit angles are larger

than the maximum suggested by Gillard & Wyss (1995) (6º), but are generally 4º - 6º

62

(Table 1). Fault populations in this study are smaller in size than other stress inversion

studies (e.g. Bergerat, 2000; Ghisetti, 2000; Saintot & Angelier, 2002; Vandycke, 2002;

Saintot et al., 2003), but because we use this technique to compare to the results of the

kinematic analysis method and not as stand-alone results, we feel confident using the

results from our small population size (5 ≤ n ≤ 28) (Table 1).

Table 3.1: Fault populations used for geometric constraint experiments and the principle stress axis orientations. Fault Population (n) Misfit Angle σ1 σ2 σ3 (º) (Trend/Plunge) (Trend/Plunge) (Trend/Plunge) Hollyford Fault (HF-005) 1 8 7.1 068:19 096:-69 161:09 2 11 7.2 345:17 016:-71 078:09 Gertrude’s Saddle (GS-06) 5 12.2 187:47 267:-10 348:41 Madeline Creek (MC-06) 7 8.9 206:07 289:-44 303:45 Lake Truth (LT-06) 7 6.2 046:19 168:57 307:26 South Mt. Thunder (SMT-06) 6 4.4 117:25 216:20 341:57 Mount Thunder (MT-05) 7 6.4 340:01 069:-52 071:38 Lake Never Never (LNN) 28 9.2 293:25 312:-63 026:08 Mount Ongaruanuku (MO) 6 4.4 268:08 252:-32 011:57

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The frictional constraint method in FSA uses, in addition to the geometric analysis, an

algorithm to investigate the effect of varying the value of the coefficient of friction (µ).

This method determines whether the reduced stress tensor enables an idealized fault

population to be activated at specific values of µ. Frictional analysis of fault populations

depends on several established friction and failure relationships for rocks. The first is

that Coulomb’s law of brittle failure (1776) for compressive stress states:

σc = σo + µ • (σN) (3.2)

where σc is the critical shear stress required for failure, σo is the cohesive strength of the

rock, and σN is the lithostatic normal stress. Additionally, frictional analysis assumes that

the cohesive strength of the rock and the internal friction angle (ϕ = arctan(µ)) are

independent of rock type (Byerlee, 1978). This assumption holds for undeformed

materials, but processes involved in faulting, such as the formation of fault gouge, the

infiltration of fluids, and mineral alteration are likely to change the value of the

coefficient of friction (Byerlee, 1978; Rice, 1992; Zoback & Healy, 1992; Lavier et al.,

2000; Townend & Zoback, 2001).

The frictional analysis by FSA provides a graphical estimate of the relationship

between the critical stress differences (sc'), the value of δ, and the location of each fault in

the fault population based on its shear stress (σs) and normal stress (σN) components.

The shear and normal stress components for each fault are calculated from eigenvalues of

the reduced stress tensor. The modified stress difference, s′, is defined as:

s′ = (σ1 - σ3)/(σ1 + τ0/tanϕ 0), (3.3)

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and has critical values of sc′ = 0.68 (for µ = 0.6, the standard frictional coefficient for

most rocks (Byerlee, 1978)), which is the stress difference required to initiate slip on a

plane; sc′ = 0.8, which is the stress difference when new faults are created; sc′ = 1.00, the

maximum value of the normalized stress difference (Célérier, 1988). The values of sc′

are given values against which the calculated modified stress difference (s′) are

compared. Sliding on a plane of weakness is likely to occur before enough stress is

accumulated to create new faults (Wallace, 1951; Jaeger, 1960; Donath, 1964; Handin,

1969; Célérier, 1988), and faults that are favorably oriented require a normalized stress

difference above a critical value to initiate sliding. Therefore, faults that, when plotted in

Mohr space and normalized in terms of the shape parameter (δ), plot between sc′ = 0.68

and sc′ = 0.80 are faults that can be activated by the stress tensor being tested (Fig. 3.9 C).

Faults that plot below sc = 1.00 may be considered over-pressured (Célérier, personal

communication, 2006), and do not fit with the assumed value of the coefficient of

friction, µ.

By modifying µ, it is possible to test the importance of pore fluid pressure, or other

strain-induced weakening mechanisms, during movement on the faults (Célérier, 1988;

Burg et al., 2005). FSA relies on equation (3.2) to model the effects of the normal stress

and the coefficient of friction. Because the coefficient of friction and the normal stress

are related through multiplication, the mathematical effect of changing one parameter

cannot be distinguished from the results of changing the other parameter. Decreasing the

input value of µ has the effect of lowering the value of the normal stress (equation (3.2));

this same effect can also be achieved by increasing the amount of pore fluid pressure (Pf).

65

Hubbert and Rubey (1959) modified equation (3.2) to illustrate this effect by including

Pf:

σc = σ0 + µ • (σN- Pf). (3.4)

Changing the value of µ allows us to test the relative value of pore fluid pressure for

different fault populations and different regions. Idealized fault populations for the

frictional constraint analysis contain the 5 largest faults from each region that have the

best constraint on slip direction and sense of motion.

66

Fig. 3.9: a, d, g: Equal-area lower-hemisphere stereoplot showing orientation of fault slip data used in frictional constrain analysis of stress tensors by FSA v. 28.5 (from Célérier (2006); see text for description of method) (diamonds: reverse motion; circles: normal motion). b, e, h: Equal-area lower-hemisphere stereoplot showing orientation of the stress tensor axes (closed pentagons: σ1; open triangles: σ3). c, f, i: Mohr circle diagrams of the stress states on each fault per population. σo is the value of the critical modified stress difference for failure along the fault, creation of new faults, and maximum stress difference criteria. σN = normal stress; σs = shear stress; σ'o = modified stress difference.

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5.2 Results using the geometric constraint

Stress solutions analyzed using the geometric constraint alone are shown in Figure

3.10. Stress solutions from the southern region show the maximum principle stress, σ1,

plots an average of 28º from the orientation of the Alpine Fault (stereoplot c; Fig. 3.10).

The exception to this pattern is the northwest striking fault population from HF-05 (Fig.

3.6), for which σ1 trends to the north. The instantaneous axis of shortening for this

population also trends towards the north. The minimum principle stress, σ3, plots an

average of 50º from the orientation of the Alpine Fault, and HF-05 presents the same

exceptions.

Stress solutions from the northern region show σ1 at much higher angles to the Alpine

Fault (stereoplot a; Fig. 3.10) compared to the stress solutions for the southern region.

The average angular distance from the Alpine Fault to σ1 is 53º in the northern region; σ3

trends at low angles to the Alpine Fault (average angle = 26º).

68

Fig. 3.10: a. Simplified site map showing major faults and lineaments. b. Equal-area lower-hemisphere stereoplot showing the orientation of σ1 and σ3 for fault populations from the northern region using the geometrical constraint analysis of stress tensors. c. Equal-area lower-hemisphere stereoplot showing the orientation of σ1 and σ3 for fault populations from the southern region using the geometrical constraint analysis of stress tensors.

69

5.3 Results using the friction constraint

Stress tensors analyzed using the frictional constraint are shown in Figure 3.9.

Results of the frictional constraint experiment for the fault population from the northern

region (Fig. 3.9 A) show σ1 plunging moderately to the west, and σ3 plunging moderately

to the northeast (Fig. 3.9 B). The orientation of the principle stresses is consistent with

the results from the geometric constraint analysis. However, results from the northern

region indicate that the frictional settings used for the southern region are not adequate to

activate faults in the northern population. The results shown in Figure 3.9 C are from

analysis using µ = 0.60, which is the same friction coefficient used for the southern

region. However, because all of the faults plot below sc′ = 1.00 on Figure 3.9 C, these

faults cannot be reactivated by the stress tensor at µ = 0.60. However, reducing the

friction coefficient to an extremely low value of µ = 0.10 allows all faults to be

reactivated by the stress tensor (Fig. 3.9 F). Using a lower coefficient of friction does not

significantly change the resulting orientation of the principle stresses (Fig. 3.9 B, E), and

suggests that the primary differences are in the deformation conditions in the northern

region, and not in the orientation of the stress fields.

The stress solution for the fault population (Fig. 3.9 G) for the southern region has σ1

in the northeast (Fig. 3.9 H), which is in good agreement with the majority of the stress

solutions determined from the geometric constraint analysis. The σ3 axis plunges gently

toward the northwest for the southern fault population, also in good agreement with

geometric results. All of the faults for the southern stress solution plot between sc′ = 0.68

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and sc′ = 0.80 (Fig. 3.9 I), indicating that all faults can be reactivated by the same stress

solution without requiring elevated pore fluid pressure.

6. Discussion

6.1 Normal faults in Fiordland

Normal faults populations from sites in the central Darran range and farther south in

central Fiordland preserve evidence for pervasive extensional and transtensional

displacements. The normal fault populations in the Darran Range are everywhere cross-

cut by the dextral strike-slip and reverse populations. On the basis of this relationship,

we conclude that the normal faults represent a phase of widespread extension that

predates all strike-slip activity related to Late Cenozoic movement on the Alpine Fault.

The populations display two dominant extension directions across central and northern

Fiordland. The largest fault populations are composed of northeast and northwest

striking faults (n = 89; n = 30; Fig. 3.11), and show X-axes trending approximately east-

west. The Mount Daniel fault population (n = 4; Fig. 3.11) is consistent with the large

fault population from the Darran Range. The smaller fault populations are composed of

primarily east-west striking faults (n = 29, n = 19; Fig. 3.11) that show X-axes trending

south-southeast.

The normal fault populations we measured are also consistent with two dominant

topographic lineament directions in central and northern Fiordland (Fig. 3.11). One set of

lineaments strikes east-northeast, and is associated with the smaller fault populations in

the region; the second group of lineaments strikes northwest and is associated with the

71

larger fault populations. There is no obvious cross-cutting relationship between the two

lineament groups, but it is unlikely that the two populations developed simultaneously

due to the poor agreement between extension axes of the associated fault populations. It

may be possible to distinguish the timing of the two populations with further mapping

and an analysis of fault populations at the outcrop-scale.

Fig. 3.11: A DEM (derived from NZMS 260 data) showing the central and northern Fiordland region with fault-plane solutions for normal fault populations. Solid lines are measured or known fault traces; dashed lines show major topographic lineaments.

72

A comparison of extensional directions indicated by the normal faults we measured

with pre-Late Cenozoic strain patterns suggest that the faults may be associated with

periods of Cretaceous and early Tertiary faulting. The south-southeast extension direction

indicated by the smaller group of faults is relatively consistent with north-south opening

directions of the Te Anau and Waiau basins (Fig. 3.1 B) during the Early Oligocene

(Norris & Turnbull, 1993). This phase of basin creation was the final stage of pure

extension in Fiordland, which would suggest that the east-northeast striking fault

population is the younger of the two groups. The northwest-striking fault populations

show east-west trending extension axes. These populations may be associated with latest

Cretaceous extension (Walcott, 1998; Wood et al., 2000). The clockwise rotation of the

plate motion vector after the Cretaceous (Walcott, 1998) may have led to a similar

clockwise rotation in the extension directions for preserved normal fault populations in

Fiordland from east-west to south-southeast.

6.2 The evolution of the stress field in the Darran Range

The average kinematic solutions we obtain from the Darran range show several

distinctive styles of displacements within northern Fiordland. Predominantly reverse and

oblique-slip fault populations occur within ~10 km of the plate boundary in the Darran

leucogabbro. The Z-axes associated with these faults plunge shallowly to the northwest

at high angles to the general northeasterly strike of the Alpine and Pembroke faults.

These Z-axes are in relatively good agreement with the orientation of both σ1 axes from

stress inversions and P-axes determined for local earthquakes (Fig. 3.12 A). All data sets

show reverse and oblique senses of slip (Moore et al., 2000). This compatibility does not

73

exist for faults from the southern region (Fig. 3.12 B). We interpret good agreement

among kinematic, paleostress and earthquake data in the northern region to indicate that

the dominant reverse- and oblique-slip fault populations in the northern Darran Range are

compatible with, and probably formed within, the modern stress field.

Fig. 3.12: Equal-area lower-hemisphere stereoplot of stress and strain data from the northern (a) and southern (b) regions. P-axes azimuths from earthquakes in the Fiordland region from Moore et al. (2000).

The relative timing of the formation and activity of the dextral faults in the Darran

Range is uncertain. It is possible that these dextral faults represent an intermediate phase

of faulting between the older normal faults and the oblique-reverse and reverse fault

populations in the north that we interpret to be related to the modern stress. The presence

of the normal faults throughout the Darran Range indicates that relict stress fields are

74

preserved in northern Fiordland, suggesting that the preservation of an intermediate stress

field is also possible. There was a phase of purely dextral strike-slip motion along the

plate boundary between ~20 and 10 Ma (Sutherland, 1995; Lamarche et al., 1997;

Walcott, 1998), and it is possible that these strike-slip faults are related to this phase of

motion. Additionally, Z-axes from strike-slip fault populations show clockwise rotation

from the south to the north, recording a rotation of the stress field, which is consistent

with a clockwise rotation of the plate motion vector (Walcott, 1998).

An alternative interpretation is that the dextral faults are concurrently active with the

oblique and reverse faults in the northern region, and that they represent further

partitioning of the stress field away from the Alpine Fault. There is deformation

partitioning within the northern region, which indicates that additional partitioning may

occur in the response to the modern stress field away from the plate boundary. P-axes

from earthquakes along the plate boundary show clockwise rotation from the Puysegur

Trench north to Fiordland (Moore et al., 2000), indicating that that modern stress field is

not homogeneous at all locations along the plate boundary.

Our preferred interpretation is that the formation of the steep dextral faults reflects an

intermediate stress field associated with transform motion along the plate boundary. We

also suggest that vertical motion on these steep faults is related to reactivation due to the

modern stress field. The clockwise rotation of Z-axes determined from the dextral fault

populations is consistent with the direction of rotation of the plate motion, which we

interpret to signify that the dextral faults formed in an older stress field. Several faults

that show vertical motion have more than one set of striae, suggesting reactivation of

these faults. Restricting the presence of the modern stress field to within ~10 km of the

75

plate boundary does not agree well with previous work documenting uplift and

shortening across the South Island (Norris et al., 1990; Pearson, 1998; Markley & Norris,

1999; Moore et al., 2000; Little et al., 2002a). Therefore we do not suggest that the

modern stress field is restricted to this zone in northern Fiordland. We propose that the

majority of strain due to the modern stress field is concentrated within ~10 km of the

plate boundary, but that the older dextral faults have been locally reactivated to

accommodate uplift related to transpression along the plate boundary. This is consistent

with the model suggested by Sutherland et al. (2006), who suggest that strain is contained

within several kilometers of the plate boundary, and that remaining strain and uplift may

be due to distributed shortening within the lithosphere.

6.3 Fluid infiltration and strain-induced weakening of the crust

There is evidence for fluid-assisted faulting in the northern deformation zone of the

Darran Range. Our analyses show that fault populations from the northern region cannot

be activated under the conditions characteristic for most rocks in the upper crust, which

include an internal friction angle of 31º and a coefficient of friction of µ = 0.60 (Byerlee,

1978; Burg et al., 2005; Célérier, personal communication, 2006). This result contrasts

with the fault populations in the southern region where hydrostatic values of these

parameters allow faulting to occur in the observed patterns. By changing the coefficient

of friction to µ = 0.10, we effectively increase the pore fluid pressure for the populations

(Eq. (3.4)). Increasing the role of pore fluid pressure in the northern region allows the

fault populations to move simultaneously, and suggests that fluid infiltration played a

76

significant role in deformation. This interpretation is supported by the presence of fluid

in veins and hydrous mineral phases in thin section from Mount Thunder (MT-05).

We interpret this difference in activity to indicate that the fault populations in the

northern region experienced different deformation conditions than populations farther

south. Conditions in the northern region include alteration to greenschist facies, the

presence of earlier deformation fabrics, and fluid infiltration. Fluid infiltration in the

northern region is supported by evidence of elevated pore fluid pressure due to water

migration from metamorphism in the lower crust, which Stern et al. (2001) interpret from

a low seismic velocity zone under the Southern Alps. Although these results are on a

different spatial and temporal scale than this study, they support the evidence from this

work for elevated pore fluid pressure. Elevated fluid pressure in the region acts to reduce

the amount of work required to deform the crust near the Alpine Fault Zone. Because the

fault patterns present in the northern deformation zone are consistent with fault

populations predicted for dextral transpression, the faults do not require fluid-induced

weakening to explain their orientation. However, the fault populations do require a

weakening mechanism to induce movement on the dominant faults (Fig. 3.9 C).

The presence of fluid and the probability of high pore fluid pressure near the Alpine

Fault serves to localize strain due to the modern stress field to within ~10 km of the plate

boundary. This localization of strain due to fluid pressure weakens the strength of the

faults in the northern deformation zone, and suggests that the Alpine Fault near Milford

Sound is frictionally weak. Balfour et al. (2005) indicate that the Marlborough section of

the Australia-Pacific plate boundary is frictionally weak, with a coefficient of friction µ =

0.35, based on the orientation of calculated stress axes. This is consistent with the low

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frictional strength indicated by the fault populations in the northern region of this study,

and is also consistent with low frictional strength of other New Zealand faults (µ = 0.17)

described by Liu & Bird (2002).

6.4 Strain localization and deformation partitioning

The Darran Range contains large-scale fault patterns that are characteristic of a

northern and a southern region. The northern zone is dominated by reverse, dextral

oblique-reverse, and dextral strike-slip fault populations that are compatible with the

modern stress field. Northeast striking fault populations shows mostly horizontal motion

and dextral strike-slip kinematics (e.g. Fig. 3.7 H), whereas north-south striking faults

show dominantly down-dip motion with reverse offset (e.g. Fig. 3.7 I) within ~8 km2

around Mount Thunder (MT-05). Z-axes for these two fault populations plot ~8º from

each other, and indicate that the kinematic response to similar compression directions is

being partitioned between fault populations of different orientations and different

kinematics in the northern region.

78

Fig. 3.13: a. Cartoon cross-section of the plate boundary and the relationship between the mylonitic fabrics and folds in the intensely deformed Anita Shear Zone and the Milford gneiss (represented by the pattern on the diagram in the region) and the brittlely fault and uplifted Darran Range. No vertical exaggeration shown. b. Block diagram of the Milford Sound region showing uplift at the margins of the Darran Range and steep brittle faults in the center of the Darran complex.

79

The near-boundary deformation zone is consistent with GPS surveys that indicate that

approximately 65% - 75% of the fault parallel motion and a variable proportion of

reverse motion is accommodated on the Alpine Fault, and that ~60% of the strain within

the South Island is concentrated within 20 km of the plate boundary (Norris & Cooper,

2001). Within Fiordland, approximately half of the relative plate motion is

accommodated between the west coast and the Waiau basin (Fig. 3.1 B) (Pearson, 1998).

A significant proportion of the remaining plate motion remains unaccounted for (~25% -

35%), and it is possible that a large portion of this motion is accommodated on strike-slip

features away from the Alpine Fault. Sutherland et al. (2006) suggest that reverse motion

is most significant near the plate boundary, which is consistent with our findings that the

majority of reverse faults are within the near-boundary deformation zone, and that the

large faults away from the Alpine Fault Zone record mostly strike-slip motion. We

suggest that strain is localized in this deformation zone partially due to weakening related

to the presence of enhanced fluid pressure, as discussed in the previous section.

The central and southern region of the Darran Range, outside of the near-boundary

deformation zone, preserve predominantly dextral-strike slip faults of uncertain age, and

subordinate thrust fault populations. The thrust populations are present at sites on the

eastern margin of the Darran Complex, and we interpret these reverse faults to indicate

that the Darran Range may be uplifted at the margins. The eastern margin of the range is

the Glade-Darran Fault, which is also a lithologic contact between the Darran Complex

and the Mistake Suite with the Brook Street Terrane to the east. A significant proportion

of reverse motion is accommodated on the Alpine Fault, but approximately 25 – 30% of

the motion remains unaccounted for (Norris & Cooper, 2001). We interpret that a

80

proportion of the remaining reverse motion may be localized as uplift at lithologic

contacts or rheological boundaries.

The style and geometry of deformation in the Darran Range contains elements that

are similar to styles described by Little et al. (2002a; Little et al., 2002b) in the central

region of the Alpine Fault Zone near the Franz Josef Glacier. Deformation near Milford

Sound is characterized by a steeply-dipping Alpine Fault, a zone of intense deformation

to the southeast of the plate boundary, and vertical motion and uplift on steep brittle

faults farther southeast (Fig. 3.13 A). The Alpine Fault is almost vertical near Milford

Sound, dipping between ~85 – 90º to the southeast (Moore et al., 2000, Eberhart-Phillips

& Reyners, 2001). The Anita Shear Zone and the Milford and Harrison gneisses,

between the Alpine Fault Zone and the Harrison-Kaipo Fault zone, form the 10 km-wide

zone near-boundary deformation zone (Blattner, 1991; Hill, 1995; Claypool et al., 2002).

The uplifted and exhumed rocks in this zone show mylonitization, foliation, and large-

scale folding, consistent with deformation farther north on the Alpine Fault (Little et al.,

2002a; Little et al., 2002b). Kinematic solutions and fault slip data indicate vertical

motion on steep faults in the central Darran Range, and uplift along the margins (Fig.

3.13 A). These features are similar to those described by Little et al. (2002a) within the

brittlely faulted region of the Alpine Fault in the center of the South Island.

Whereas schists in the center of the South Island are being uplifted by an “escalator-

like” motion along a fault ramp (Little et al, 2002a), the uplift in northern Fiordland is

accommodated through processes related to lithologic boundaries and the splay-like

geometry of the terrane boundary faults, such as the Glade-Darran and Hollyford faults

(Fig. 3.13 B). Reverse motion is localized at the eastern margin of the Darran Complex,

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which is partially bounded by the Glade-Darran Fault, and results in the uplift of the

rheologically strong gabbros and diorites in the Darran Range. Additionally, the center

and southern Darran Range may be uplifted by vertical displacement on reactivated steep

faults that is kinematically linked to strike-slip motion due to the modern stress field. We

interpret the elevated topography of northern Fiordland and reverse motion to be related

to the extrusion and brittle wedge-like uplift of the Darran Range.

7. Conclusions

The majority of Tertiary reverse and oblique-reverse displacements in northern

Fiordland are localized with a ~10 km-wide near-boundary deformation zone to the

southeast of the Alpine Fault. This zone of proximal deformation is characterized by

fault populations of widely varying orientations, and there is evidence of fluid infiltration

and elevated pore fluid pressure. Outside of this zone, to within ~20 km, strike-slip and

vertical displacements are localized along lithologic boundaries and on reactivated steep

faults, leading to the uplift of the Darran Range. The majority of Tertiary fault

populations in northern Fiordland have dextral, dextral-oblique, and reverse sense of

motion, which is consistent with dextral transpression along the Alpine Fault.

Within the near-boundary deformation zone, contraction axes and σ1 axes trend

approximately 60º from the plate boundary, in relatively good agreement with P-axes

from earthquakes in the region. Fault populations in the deformation zone can only be

activated at a very low coefficient of friction (µ = 0.10), suggesting frictional weakening

due to elevated pore pressure. We interpret the elevated fluid pressure, combined with

the high angle between maximum compression directions and the northeasterly strike of

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the plate boundary, to indicate that the Alpine Fault is relatively weak in the Milford

Sound region. This weakening is consistent with other studies of the Alpine Fault and

New Zealand that describe low frictional strength of faults. A relatively weak Alpine

Fault would indicate that the San Andreas Fault is not unique in its status as a weak plate

boundary fault, and further work on the frictional strength of large transpressional and

transtensional plate boundary faults would help to determine which processes control the

strength of crustal scale strike-slip faults.

Finally, fault patterns and cross-cutting relationships indicate the superposition of

several stress states preserved in Fiordland. These phases include an older extensional

regime, an intermediate or modern dextral phase of faulting, and reverse faults that are

compatible with the modern stress field.

83

Chapter IV: Discussion

1. Overview and conclusions

The results of this thesis address strain localization and fault-zone weakening

processes in the upper crust in northern Fiordland. This thesis contributes to the study of

transpression along the Alpine Fault, and to the growing literature on the frictional

strength of crustal-scale and plate boundary faults. I use field-based structural

observations and fault-slip data to characterize deformation patterns in the Darran Range

in northern Fiordland. Kinematic analysis and stress inversion of fault-slip data allowed

me to determine spatial patterns and the compatibility of stress fields and slip on

dominant fault populations in different regions. By combining field and modeling

techniques, I determined the mechanisms and processes that governed the structural

evolution of the upper crust in northern Fiordland.

The main result of this thesis is the description of a partitioning of deformation into

different zones with increasing distance from the plate boundary. These zones are

characterized by their dominant fault populations, and by the relative frictional strength

of the faults themselves. In the near-boundary zone, motion is partitioned between

distinct fault populations, whereas reactivated brittle faults in the southern zone

accommodate both oblique-reverse and strike-slip motion. In addition to the

documentation of deformation zones, this thesis also presents evidence for the

superposition of several distinctive stress fields in central and northern Fiordland that

correlate to different phases of motion.

Brittle fault populations in the near-boundary deformation zone are compatible with

the modern stress field, and consist of dominantly reverse, oblique-reverse, and dextral

84

strike-slip populations. These fault populations are also consistent with patterns

predicted for transpression. I interpret the variety of orientations and slip on fault planes

to indicate that the kinematic response to compression is partitioned among different fault

populations. Fault patterns and the frictional strength of the faults indicate that strike-slip

partitioning and fluid-induced weakening is more pronounced in the near-boundary

deformation zone. The variety and the frictional weakness of fault populations in this

zone is not widely recognized more than ~10 km to the southeast of the plate boundary. I

suggest that the near-boundary deformation zone is partially due to strain-induced

weakening from the presence of elevated pore fluid pressure in the region. The enhanced

fluid pressure significantly reduces the frictional strength of the dominant faults in the

upper crust, and may reduce the bulk strength of the crust near the Alpine Fault. Away

from the plate boundary, vertical and oblique-reverse motion is localized on lithologic

boundaries, and strike-slip, oblique-reverse, and vertical motion reactivates steep, brittle

faults.

In addition to the results of this work that describe strain partitioning and crustal

strength, brittle faults in central and northern Fiordland preserve evidence for the

superposition of several stress fields. Large normal faults are everywhere cut by steep

dextral and reverse faults, and these normal faults correlate well with two major

topographic lineament orientations in the region. The two normal fault populations

predate recent Tertiary movement, and may be associated with faulting related to Late

Cretaceous and Early Tertiary extension.

The results of this work are consistent with other studies that describe distinct

deformation zones at regions farther north on the Alpine Fault. Additionally, several

85

other studies suggest that the crust around the Alpine Fault is relatively weak on the basis

of focal mechanism inversion and shear-wave splitting. The findings of this thesis

support these studies, and suggest that elevated pore fluid pressure may be partially

responsible for crustal weakening, in addition to inherited geologic structures.

This work suggests that elevated pore fluid pressure and frictional weakening of the

crust near the plate boundary are the dominant processes for strain localization in the

upper crust in northern Fiordland. The inherited sutured terrane structure of the plate

boundary region assists in the localization of strain along rheological and lithologic

boundaries, but it is not the most significant mechanism for deformation partitioning in

the upper crust. Inherited fabrics, such as the foliation and mylonitization seen in the

Anita Shear Zone and Harrison-Kaipo fault zone, may assist in localizing strain near the

plate boundary.

2. Future work

Future work on strain localization and crustal weakening in New Zealand should

focus on the relative extent of elevated pore fluid pressure, and on the influence of

inherited structures on strain partitioning. By comparing structures and deformation

mechanisms in Fiordland to those present at the northern end of the South Island, it may

be possible to test the relative importance of fluid infiltration. Because the Alpine Fault

off-sets the northern end of the Median Batholith and the sutured terranes, structural

analysis of the Nelson and Greymouth area of the South Island should produce similar

results to those of this study. Differences that may arise from such a study could have

86

significance regarding the important of fluid pressure, lithology, or the geometry of the

plate boundary.

Future work may also focus on the nature of the hydrous mineralogy in the near-

boundary deformation zone. Understanding the source of pore fluids in fault gouge,

veins, and in the host rock, and clarifying the deformation conditions of faulting may help

to describe the flow pattern of pore fluids through fault rocks. A better understanding of

the upper crustal flow paths of pore fluids could help us better understand the exhumation

of deep crustal rocks, as well as better understand the regional extent of elevated pore

fluid pressure.

Finally, a study of the frictional strength of other obliquely convergent continental

plate boundary faults such as the North Anatolian Fault or the Altyn Tagh Fault in the

Himalaya may help to determine the prevalence of weak plate boundary faults. If the

Alpine Fault and the San Andreas Faults are relatively weak, it is possible that the

majority of large continental plate boundary faults are weak due elevated pore fluid

pressure. The presence of relatively weak plate boundary faults may explain the

continued reactivation of crustal-scale faults, even if they are not optimally oriented for

slip due to plate motion relative to the stress field.

87

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Appendices:

Appendix A: Fault-slip data used for kinematic analysis Appendix B: Fault-slip data used for stress inversion with only geometric constraint

102

Appendix A: Fault-slip data used for kinematic analysis Fault-slip data that include question marks in the ‘Sense of motion’ column were included in populations of faults with similar fault plane and slickenline orientation. By plotting the fault plane solution for each population, I was able to test the appropriateness of the assigned slip-sense. Fault-slip data with kinematic axes that clustered with axes for faults with known kinematics were deemed to have an appropriate slip-sense assignment. Incompatible faults were tried with other fault populations for a better fit, and included in the most appropriate population to insure the inclusion of the maximum number of measured faults. Description of ‘Sense of motion:’ “d” = likely dextral motion “D” = described as dextral motion in field “s” = likely sinistral motion “S” = described as sinistral motion in field “n” = likely normal motion “N” = described as normal motion in field “t” = likely thrust motion “T” = described as thrust motion in field Description of ‘Notes:’ “Ques.” = Questionable “Slicks” = Slickenlines “Chat” = Chattermarks “Kin” = Kinematics Fault plane orientation Slickenline Sense of motion Notes Hollyford Fault (HF-05) Dextral faults: n = 21 Slickenline Sense of motion Notes 208:67 W 024:10 d 3 m long 031:74 W 028:12 d 012:88 W 011:24 d 020:85 N 201:08 Dextral chat 352:88 N 172:08 d 015:75 W 010:20 d 355:83 W 182:43 d 018:88 W 018:12 d A 002:82 W 358:26 d B 004:81 W 186:14 d B 006:74 W 351:42 d A 210:90 030:28 Dextral 185:79 W 003:11 d 030:79 W 025:26 Dextral 015:81 W 003:51 d 005:85 E 184:15 d 355:86 E 357:27 d A 009:73 E 026:44 d A

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215:68 E 040:12 Normal 025:70 S 201:12 d 215:68 E 040:12 d 079:76 S 132:72 normal B 055:79 W 042:50 n? 052:56 S 157:56 n? Dextral faults: n = 23 Slickenline Sense of motion Notes 313:65 S 313:02 d 312:89 N 132:20 d small B 312:73 N 121:32 d large B 296:86 S 119:36 d 326:82 N 143:21 d dominant set 310:75 N 127:10 Dextral 325:78 W 322:13 d A 289:76 S 278:08 Dextral A 321:87 S 320:09 d 315:88 S 136:10 d 302:85 N 120:21 d A 125:90 305:04 d 297:86 E 115:21 d 289:87 E 106:45 d 316:85 S 315:16 d 342:80 S 168:30 d questionable 336:82 S 335:04 d A 145:90 325:08 d 143:73 W 145:06 d 346:77 S 197:66 ?? 270:82 E 268:13 d Thrust faults: n = 15 Slickenline Sense of motion Notes 299:59 N 109:16 Dextral 296:58 N 036:58 t 289:44 N 318:25 t 286:66 N 296:22 t 290:21 N 017:22 t 278:75 N 008:75 t 273:64 N 007:64 t 082:60 N 069:21 d 081:16 N 346:16 Thrust B 035:48 N 280:45 t 016:19 N 331:14 t C 042:33 N 000:23 t B

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300:20 S 273:10 t A 347:26 W 317:14 Sinistral reverse A 347:41 W 330:14 Reverse A Gertrude’s Saddle (GS-06) Dextral Thrust: n = 4` Slickenline Sense of motion Notes 001:90 181:39 D? dominant set 001:87 W 182:20 ?d C 357:88 W 355:56 ?d B 006:82 W 359:39 ?d Large (50m) 354:84 W 177:25 D? Ques. chattermark C 354:88 E 173:14 D? Large 100m 359:64 E 044:55 ?d 348:84 E 000:50 D Large 100m 356:86 E 005:61 ?d 008:79 W 194:28 D Ques. chattermark C 349:81 N 168:08 D? Ques. s-c C 356:86 W 344:73 ?d 350:81 E 160:49 D major fault 100m 350:80 E 158:50 D? 100m 347:80 E 159:37 D? 100m 352:77 E 155:50 D 349:84 E 158:59 D 100m 015:87 E 195:39 D B Large 100m 015:67 E 171:44 d 1 m 185:79 N 001:19 d 000:85 E 005:47 d 019:79 N 013:29 d 355:84 S 351:36 dextral 352:88 N 353:25 dextral 350:86 W 172:22 d A 20-30m 025:77 E 198:30 d A 112:38 W 169:37 ?t "layering" Large 286:39 S 154:28 ?t "layering" 024:78 W 227:60 ?t cut by 070:85w 337:68 W 235:68 ?t large 50 m? 338:70 W 218:67 ?t 082:27 N 002:27 T? due to offset 356:60 E 098:60 N notes say N 085:24 N 004:24 T offset fractures 345:66 E 036:61 t 1 m 344:66 N 082:87 t 3 m 319:89 N 140:51 t 30 cm 348:86 W 186:77 t 3 m

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320:84 N 113:77 t B 4 m 322:31 N 046:38 t B 168:89 S 347:40 T 8 m Normal faults: n = 35 Slickenline Sense of motion Notes 062:83 N 249:45 ?s cut by dex. fault small 052:81 S 056:24 ?s 070:85 W 253:24 N ?s sinistral oblique 074:69 W 272:39 S Ques. chattermark C 292:75 N 110:05 ?s cut by below 085:70 N 264:00 ?s offset and s-c 072:63 S 090:30 ?s 329:71 E 145:12 ?D 100m 315:74 N 318:04 ?s Slickenlines ques. C 069:60 N 255:23 N ?s Ques. chattermark C 100:27 S 106:06 s? A 010:70 E 173:40 ?n 299:64 E 033:49 ?n 295:72 E 311:40 ?n Large 100m 315:43 S 181:29 ?n small fault 035:71S 073:51 ?n 311:50 S 213:50 ?n 016:65 E 168:45 ?n slicks not great C 015:74 W 141:70 ?n 293:55 N 087:35 N chat 011:80 W 222:70 N A small 011:80 W 205:54 N 15-m trace 332:34 W 244:15 ?n slicks not great C 278:62 N 296:40 ?n 069:60 N 255:23 N ?s 075:70 S 088:32 N (chat) 3.5 m 086:56 N 048:43 n small 271:50 S 107:19 N B 290:55 N 039:54 n 1 m 310:72 E 053:83 n 30 cm 301:78 S 197:87 n 6 m 296:68 S 241:64 n 133:67 S 235:70 n 294:74 N 325:61 n 296:49 N 105:13 n A 4m Homer's Tunnel (HT-05) Dextral faults: n = 9 Slickenline Sense of motion Notes 358:87 E 359:26 ?d

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005:78 E 166:57 ?d 013:43 S 155:31 ?d Questionable C 014:20 S 122:19 ?d C 028:75 E 034:20 ?d 029:65 S 041:25 ?d 030:88 S 032:46 ?d 035:46 E 211:04 ?d 204:10 N 330:09 ?d Normal faults: n = 19 Slickenline Sense of motion Notes 056:19 S 101:14 T? kin via chattermarks 062:52 S 114:45 n? 072:54 S 226:31 n? 076:84 N 074:24 n? 079:26 S 026:20 n kin via chat. B 081:44 S 150:42 n? Questionable C 088:52 S 251:21 n? 102:30 S 140:20 n kin via chattermarks 269:81 N 085:23 n? 280:90 280:16 n? 25 m 284:21 S 152:16 n? B 289:42 N 080:24 n? B 121:62 S 137:27 n? 128:70 N 110:40 sin? Questionable C 132:84 N 129:24 n? 135:76 N 123:40 n? 308:40 S 126:09 n? 310:80 N 112:61 n kin via offset 320:84 E 085:83 n? Madeleine creek (MC-06) Dextral faults: n = 14 Slickenlines Sense of motion Notes 324:81 W 147:17 d? 330:76 W 151:02 D s-c B 336:79 E 342:29 d? D/t 329:71 E 337:22 d? 313:64 E 319:12 d? C 312:83 E 319:45 D/t 305:83 E 309:30 d? 339:68 N 147:26 D Questionable C 340:66 W 163:06 d? Questionable C 176:83 E 171:35 D 130:84 S 138:53 D A 315:85 E 126:59 t?

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335:90 335:72 t? 133:58 N 089:47 T Reverse faults: n = 11 Slickenline Sense of motion Notes 152:36 S 152:00 d? 010:43 E 178:10 D 194:85 W 011:30 D 097:51 S 250:29 D chat. 200:62 N 212:23 D 359:51 W 230:45 d? Questionable C 070:65 N 021:58 t? 042:45 S 162:41 T 061:44 S 177:41 T 057:44 S 150:43 good slicks. No chat. 082:60 S 209:54 T older Sinistral oblique-normal: n=10 Slickenline Sense of motion Notes 281:83 S 177:83 S offset 060:62 S 225:25 s? 342:80 S 167:26 S C 045:54 S 076:35 s? 340:85 W 339:15 s? 084:57 S 103:27 S 082:60 S 102:26 S 125:64 S 094:47 S chat 074:73s 161:73 n? 070:65n 021:58 350:66e 066:64 N Lake Truth (LT-06) Dextral faults: n = 8 Slickenline Sense of motion Notes 341:87 W 340:38 ?d 176:67 S 342:36 ?d 331:84 S 153:24 ?d 140:85 E 327:29 ?d 330:52 N 356:30 ?d 334:67 S 328:15 D chat 004:80 E 013:40 T d? chat? 025:70 E 039:35 ?d Normal faults: n = 49 Slickenline Sense of motion Notes 076:60 S 099:34 ?n large 20 m 048:76 E 059:35 ?n

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060:77 E 073:44 ?n 044:54 S 088:43 T (?n) C 15 m 034:54 S 097:51 ?n 024:61 S 035:20 ?n dominant set 15 m 009:53 S 098:51 N minor fault 2m 055:50 S 100:40 ?n 010:55 S 140:40 ?n ? B .3m 035:53 S 105:53 ?n dominant set A 15m 007:60 E 205:12 ?n dominant set A 15m 065:46 S 100:30 ?n offsets below (275) 035:47 E 077:35 D ?n 005:50 E 070:40 n? 314:52 N 118:18 ?n epidote slicks 2 m 327:54 N 103:44 ?n 329:62 S 259:61 ?n 329:64 S 222:64 ?n smaller 2m 319:61 N 354:46 ?n 320:64 N 334:26 ?n 315:56 N 104:35 ?n 320:34 N 079:31 ?n 315:55 N 123:17 ?n 018:65 W 213:29 ?n 034:49 S 065:30 ?n slicks questionable 030:44 S 093:40 ?n B 045:69 E 060:34 N Chat B 349:50 E 160:10 ?n slicks not great C 114:34 N 061:17 N chat good A 019:45 S 185:14 DN? along dike 100m 020:42 S 177:19 ?n weigh same as above 047:50 S 077:33 ?n major fault 20 m 020:48 E 070:32 ?n dominant 25 m 114:34 N 061:17 N chat 040:52 E 345:24 ?n not great C 030:47 S 100:45 ?n slicks not great C 290:56 N 302:17 s small 2m 316:79 N 310:49 s 296:70 N 105:28 N chat 221:40 E 098:36 N chat 054:44 N 302:41 ?n smaller 2m 298:39 S 252:31 ?n smaller 2m 040:30 S 170:27 ?n 064:54 N 013:47 ?n 318:31 N 025:30 N 290:46 N 336:37 N master fault 200 m

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034:74 E 047:39 ?n 012:83 W 200:49 ?n 031:60 N 275:52 ?n smaller 2m S. of Mt. Thunder (SMT-06) Dextral/thrust fault: n = 12 Slickenline Sense of motion Notes 225:85s 050:48 D chat. A 002:67e 098:36 S-T Possibly B 185:41e 085:41 T C 75m 336:63n 057:68 D offset. Slicks ques. 350:40e 047:35 T dike offset A 037:31s 061:48 D-T s-c B 150:73e 058:73 S-T chat. 020:46e 064:35 D possibly? 000:62 e 325:56 t 005:65e 047:56 t? C 015:65e 079:63 T s-c B 005:54e 030:25 t Mt. Thunder (MT-05) Dextral faults: n = 12 Slickenline Sense of motion Notes 008:75 s 186:05 d small .2 010:60 n 003:12 T NW-up 014:64 n 009:11 dex kin via offset .2 018:66 e 019:03 sin kin via chattermarks 023:64 n 019:08 dex kin via offset A .2 031:66 n 218:16 dex kin via offset .2 032:45 n 351:34 ?T B 032:66 s 210:03 d Large 50m 037:68 s 039:05 d 039:76 n 343:14 ?T dex kin via offset .2 044:66 n 231:15 d 047:48 n 247:21 dex kin via chattermarks Dextral oblique: n = 20 Slickenline Sense of motion Notes 056:63 N 035:35 ?T 50m 057:47 N 253:16 ?d 057:55 N 039:24 ?T 057:61 N 240:05 d 058:57 E 157:57 ?T normal? 061:58 N 056:08 d 50m

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063:48 N 032:30 ?T 100m 065:75 S 226:50 ?T dex kin via s-c 067:55 N 056:16 dex? small 2m C 068:83 N 057:58 ?T 069:77 S 248:02 dex kin via s-c & offset A 070:81 S 097:70 ?T dex kin via s-c D 070:82 N 058:57 ?T 071:74 N 045:57 ?T 071:84 N 060:62 ?T 074:50 N 263:10 ?t 076:70 S 176:70 ?T dex B kin via s-c 078:79 N 070:20 ?d 084:33 N 066:11 ?t 084:64 S 248:30 ?t 50m Sinistral oblique: n = 16 089:76 N 087:07 ?s 275:72 N 090:16 ?s 276:73 N 093:12 sin kin via chattermarks 282:79 N 098:20 ?s 106:40 N 094:10 ?s 120:45 N 104:15 ?s 286:65 S 285:02 ?s 305:77 N 116:35 dex kin via visible offset 312:66 N 122:21 ?s 315:79 S 313:10 ?s 319:45 N 340:20 ?T 324:72 S 321:08 ?s 325:52 S 159:17 ?s 338:90 338:24 ?d based on s-c C 343:51 S 128:36 ?T 352:50 S 329:26 thrust kin via s-c

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Appendix B: Fault-slip data used for stress inversion with only geometric constraint Hollyford Fault (HF-05) dextral population: 'HF 000 dextral' n = 8 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

208 67 'W' 24 'A' 'D' 1 31 74 'W' 28 'A' 'D' 2 15 75 'W' 10 'A' 'D' 6 18 88 'W' 18 'A' 'D' 8 4 81 'W' 186 'A' 'D' 10

185 79 'W' 3 'A' 'D' 13 355 86 'E' 357 'A' 'D' 18 25 70 'S' 201 'A' 'D' 20

Hollyford Fault (HF-05) dextral population: 'HF 100 dextral' n = 11 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

313 65 'S' 313 'A' 'D' 1 312 89 'N' 132 'A' 'D' 2 312 73 'N' 121 'A' 'D' 3 326 82 'N' 143 'A' 'D' 5 310 75 'N' 127 'A' 'D' 6 325 78 'W' 322 'A' 'D' 7 125 89 'N' 305 'A' 'D' 11 336 82 'S' 335 'A' 'D' 15 145 89 'W' 325 'A' 'D' 16 302 85 'N' 120 'A' 'D' 18 90 82 'S' 268 'A' 'D' 21

Gertrude’s Saddle (GS-06) dextral population: 'GS dextral’ n = 5 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

357 88 'W' 355 'A' 'D' 3 6 82 'W' 359 'A' 'D' 4 0 85 'E' 5 'A' 'D' 19

355 84 'S' 351 'A' 'D' 25 348 89 'S' 347 'A' 'I' 41

Homer’s Tunnel (HT-05) dextral population: 'HT dextral' n = 7 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

358 87 'E' 359 'A' 'D' 1 13 43 'S' 155 'A' 'D' 2 28 75 'E' 34 'A' 'D' 4 29 65 'S' 41 'A' 'D' 5

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30 88 'S' 32 'A' 'D' 6 35 46 'E' 211 'A' 'D' 7 5 78 'E' 359 'A' 'D' 8

Madeline Creek (MC-06) dextral & reverse populations: MC dextral thrust n = 7 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

324 81 'W' 147 'A' 'D' 1 330 76 'W' 151 'A' 'D' 2 336 79 'E' 342 'A' 'D' 3 329 71 'E' 337 'A' 'D' 4 305 83 'E' 309 'A' 'D' 7 340 66 'W' 163 'A' 'D' 9 130 84 'S' 138 'A' 'D' 21

Lake Truth (LT-06) dextral population: LT dextral' n = 7 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

341 87 'W' 340 'A' 'D' 1 176 67 'S' 342 'A' 'D' 2 140 85 'E' 327 'A' 'D' 4 330 52 'N' 356 'A' 'D' 5 334 67 'S' 328 'A' 'D' 6

4 80 'E' 13 'A' 'D' 7 25 70 'E' 39 'A' 'D' 8

South of Mount Thunder (SMT-06) dextral & reverse populations: SMT dextral thrust n = 6 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

336 63 'N' 57 'A' 'I' 4 350 40 'E' 47 'A' 'I' 5 37 31 'S' 61 'A' 'I' 6 20 46 'E' 64 'A' 'I' 7 15 65 'E' 79 'A' 'I' 9 0 54 'E' 68 'A' 'I' 10

Mount Thunder (MT-05) strike-slip population: MT Sin. + Dex.' n = 7 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

56 63 'N' 35 'A' 'I' 1 58 57 'E' 157 'A' 'I' 5 68 83 'N' 57 'A' 'I' 10 70 81 'S' 97 'A' 'I' 12 70 82 'N' 58 'A' 'I' 13 71 74 'N' 45 'A' 'I' 14

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71 84 'N' 60 'A' 'I' 15 Lake Never-Never (LNN) strike-slip populations; ‘LNN' n = 28 az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

252 89 'W' 72 'A' 'D' 13 236 74 'W' 51 'A' 'D' 14 60 79 'S' 62 'A' 'D' 15 51 59 'S' 53 'A' 'D' 16 47 68 'E' 48 'A' 'D' 17 52 59 'E' 229 'A' 'D' 18 39 67 'E' 216 'A' 'D' 20 46 56 'E' 224 'A' 'D' 23

213 70 'W' 239 'A' 'D' 24 245 70 'N' 60 'A' 'D' 26 80 44 'S' 237 'A' 'D' 29 1 89 'E' 1 'A' 'S' 35

352 89 'E' 352 'A' 'S' 39 180 80 'W' 356 'A' 'S' 43 175 88 'W' 354 'A' 'S' 44 354 62 'E' 0 'A' 'S' 46

4 79 'E' 6 'A' 'S' 47 341 75 'E' 344 'A' 'S' 48 340 82 'E' 341 'A' 'S' 49 296 78 'E' 329 'A' 'S' 50 293 68 'N' 7 'A' 'S' 51 299 80 'N' 358 'A' 'S' 52 354 63 'E' 6 'A' 'S' 54 155 89 'W' 335 'A' 'S' 55 172 89 'W' 352 'A' 'S' 58

0 80 'E' 5 'A' 'S' 59 20 80 'E' 25 'A' 'S' 64 22 74 'E' 29 'A' 'S' 65

Lake Pukutahi (LP) dextral population: LP' n = 7 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

46 84 'E' 47 'A' 'D' 1 33 64 'E' 47 'A' 'D' 2

218 80 'W' 35 'A' 'D' 3 36 60 'E' 51 'A' 'D' 4

214 63 'W' 25 'A' 'D' 5 92 74 'S' 126 'A' 'S' 6

194 88 'W' 10 'A' 'D' 7

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Mount Ongaruanuku (MO) oblique thrust population: 'MO' N = 6 'az' 'dip' 'dd' 'trend' 'quad' 'type' 'id'

342 55 'N' 54 'A' 'D' 3 344 16 'N' 33 'A' 'D' 4 330 36 'E' 36 'A' 'D' 6

2 80 'E' 13 'A' 'D' 10 37 74 'E' 48 'A' 'D' 15 26 56 'E' 59 'A' 'D' 16