Thesis

Modelling the Fragmentation of Cordilleran
Orogens: Applications to the Lord Howe Rise
Bailey Payten
Supervisors: A-Prof. Patrice Rey, Dr. Simon Williams, Prof.
Dietmar Mueller
Thesis submitted in partial fulﬁlment of the requirements for the degree of Bachelor of
Science (Honours)
The University of Sydney
2016
Word Count: 18665
Submitted: 21/10/2016

Acknowledgements
Firstly I would like to thank my supervisors for helping me so much this year. Thanks to Patrice for
aiding and shaping my entire approach to modelling; the results of this project would be far below
what they are now without your guidance. Thanks also to Simon for helping me to understand
much of the conceptual information that underpins this topic. This was particularly important early
in the study and on the research cruise. Thanks also to everyone in Eathbyte who offered advice
all through this year.
A special thanks goes out to James for keeping me sane this year. Lastly, thanks to Georgia for
keeping me on the right track and reminding this was the right thing to do!

Abstract
The fragmentation of Eastern Gondwana led to the formation of the worlds largest continental
ribbon, the Lord Howe Rise. This fragmentation occurred through west-dipping subduction of the
Paciﬁc plate and subsequent back-arc extension in the Eastern Gondwana Cordillera; the
mechanisms of which are not well constrained. Through thermo-mechanical numerical modelling,
this study explored the impacts of the Cordillera itself, a buoyant mantle wedge and kinematic plate
movement on back-arc extension and subduction. A palinspatic reconstruction of the Lord Howe
Rise and Eastern Australian margin was performed, with reconstructed extensional velocities,
crustal thicknesses and crustal geometries derived for comparison to the numerical model output.
Presence of a Cordillera introduces a gravitational potential energy anomaly resulting in orogenic
collapse which then promotes trench retreat. Removal of this orogen caused a ~60 km decrease in
continental extension. Introducing a buoyant mantle increases continental extension by ~70 km
through development of higher stresses promoting lateral extension in the overriding continental
crust. Most importantly, the study found that removing horizontal oceanic plate velocity to develop
increased extension by ~90 km, with continental crust hyperextension and asymmetry matching
the narrow Eastern Australian margin and hyperextended Lord Howe Rise. Rift migration is
proposed to be the primary factor that controlled the formation of this crustal geometry.

Table of Contents
1. Introduction……………..……………………………………………………………………………1
2. Background…………..………………………………………………………………………………2
2.1. Tectonics of the Lord Howe Rise………………………………………………………………2
2.1.1. Current Geologic Setting………………………………………………………………….2
2.1.2. Early to Mid Cretaceous…………………………………………………………………..5
2.1.3. Late Cretaceous to Eocene………………………………………………………………6
2.1.4. Eocene to Present Time…………………………………………………………………10
2.1.5. Comparable Global Environments: The North West Shelf and the Basin and Range
Province………………………………………………………………………………………11
2.2. Continental Extension at Active Margins………………………………………………………14
2.2.1. Fragmented Crustal Block Types and Their Composition……………………………14
2.2.2. Active and Passive Rifting……………………………………………………………….14
2.2.3. Causes of Rifting at Active Margins………………………………………………….…15
2.2.4. Rift Architecture and Evolution………………………………………………………..…19
2.3. Knowledge Gaps…………………………………………………………………………………22
3. Methodology………………..………………………………………………………………………23
3.1. Developing a Palinspatic Reconstruction of Eastern Gondwana…………………………..23
3.1.1. Crustal Thickness Map…………………………………………………………………..23
3.1.2. Bouguer Gravity and Total Horizontal Derivative Maps………………………………24
3.1.3. Determining the Unstretched Continental Crust Limit and Continent - Ocean
Boundaries...…………………………………………………………………………………25
3.1.3.1. Eastern Australian Margin…………………………………………………………28
3.1.3.2. Lord Howe Rise………………….…………………………………………………28
3.1.4. Stage Pole Rotation of the Lord Howe Rise Relative to Australia…….……………29
3.1.5. Reconstruction of the Paleo-Crust……………………………………………………..29
3.1.6. Determining the Extensional Velocity…………………………………….……………30
3.2. Thermo-Mechanical Modelling with Underworld…………………………………………….31
3.2.1. Overview……………………………………………………………………………………31
3.2.2. General Equations…………………………………………………………………………32
3.2.3. Lithospheric Modelling Recipe…………………………………..………………………34
3.3. The Reference Model…………………………………………………………………………35
3.3.1. Outline and Development……………………………………..…………………………35
3.3.2. Model Architecture and Rheology………………………………………………………35
3.3.3. Initial Temperature Conditions………………………………..…………………………37
3.3.4. Imposed Velocity Conditions………………………………….…………………………38

3.3.5. Surface Processes and Partial Melting………………………………………………38
3.4. Model Roster………………………………………………………………………………..…41
3.4.1. Influence of a Buoyant Mantle Wedge……………………………………………….…41
3.4.2. Anorogenic Continental Crust……………………………………………………………41
3.4.3. Subduction of an Oceanic Lithosphere with Zero Horizontal Velocity………………41
4. Results……………………..…………………………………………………………………………42
4.1. Palinspatic Reconstruction and Extensional Velocities in Eastern Gondwana……………42
4.2. Numerical Models………………………………………………………………………………44
4.2.1. Evolution of the Reference Model………………………………………………………44
4.2.2. Influence of a Buoyant Mantle Wedge…………………………………….……………46
4.2.3. Influence of an Oceanic Plate with Zero Horizontal Velocity…………………………47
4.2.4. Influence of an Anorogenic Continental Crust…………………………………………51
5. Discussion……………..……………………………………………………………………………54
5.1. Comparisons of the Models to the Lord Howe Rise…………………………………………54
5.2. Slab Dynamics….…….……………………………………….…………………………………57
5.3. Limitations…….……………………………………………….…………………………………58
5.4. Future Work…..……………………………………………….…………………………………59
6. Conclusion………………………………………………..…………………………………………61
7. References……………………………………………………..……………………………………62
8. Appendices……………………..……………………………..……………………………………70
8.1. Palinspatic Reconstructions and Extensional Velocities along the Eastern Australian
Margin, Tasman Basin and Lord Howe Rise for a Pre-Deformational 60 km Thick
Continental Crust……………………………………………………..………………………….70
8.2. XML Input Files for the Models…………………………………………………………………73
8.3. Movies of Model Evolution………………………………………………………………………73

Table of Figures
Figure 2.1: Detailed geological map of the Lord Howe Rise region From: Higgins et al. (2015)…..3
Figure 2.2: Bathymetry of the Southwest Pacific region………………………………………………..4
Figure 2.3: Schematic representation of subduction and back-arc extension at the Eastern
Gondwana margin from the Late Cretaceous to Early Eocene. From: Cluzel et al. (2001)……….…8
Figure 2.4: Diagrammatic representation of the various models for the subduction history East of
the Lord Howe Rise. From: Matthews et al. (2015)………………………………………………………9
Figure 2.5: Modern day location of the Lhasa terrane with locations of arc basalts and eclogites
within it visualised. From: Metcalfe (2013)……………………………………………………………….12
Figure 2.6: Graphic representation from 250 - 145 Ma of proposed back-arc extension and
fragmentation of the Lhasa block. From: Zhu et al. (2011)…………………………………………….12
Figure 2.7: Schematic representation showing delamination of an oceanic slab resulting in back-
arc extension in the overriding continental crust. From: Sutherland et al. (2010)………………..…16
Figure 2.8: Influence of a buoyant mantle on back-arc extension. From: Rey and Mueller (2010)
………………………………………………………………………………………………………………..18
Figure 2.9: Schematic representation of the influence of a mantle plume on rift migration. From:
Mueller et al. (2001)….……………………………………………………..……………………………..19
Figure 2.10: The nature of rift migration in varying extensional velocity environments. From: Brune
et al. (2014)….……………………………………………………..……………………………………….20
Figure 3.1: Crustal thickness map of the Eastern Australian margin.………………………………..25
Figure 3.2: Crustal thickness map of the Lord Howe Rise region.……………………………………25
Figure 3.3: Total horizontal derivative of the Bouguer gravity map for the Eastern Australia - Lord
Howe Rise region.…………………………………..…………………………………..………………….26
Figure 3.4: Bouguer gravity map for the entire Eastern Australia - Lord Howe Rise region……….27
Figure 3.5: Initial material geometry and temperature field for the reference model……………….36
Figure 3.6: Development of the subducting slab used as the initial material geometry for the
reference model.…………………………………..…………………………………..……………………37
Figure 4.1: Locations of the small circle traces used for extensional velocity reconstructions laid
over crustal thickness maps in the Eastern Australian and Lord Howe Rise regions……………….43
Figure 4.2: Evolution of the reference model in it’s entire domain until 6 Myr………………………44
Figure 4.3: Evolution of the reference model zoomed in on a 800 km by 200 km segment
concentrating on the continental crust..…………………………………..………………………..……45
Figure 4.4: Evolution of the buoyant mantle wedge model zoomed in on a 800 km by 200 km
segment concentrating on the continental crust.…………………………………..……………………46

Figure 4.5: Evolution of the zero horizontal oceanic plate velocity model in its entire domain until
7.50 Myr.…………………………………..…………………………………..…………………………….48
Figure 4.6: Evolution of the zero horizontal oceanic plate velocity model zoomed in to a 850 km by
200 km segment concentrating on the continental crust.………………………………………………49
Figure 4.7: Viscosity profile of the zero horizontal oceanic plate velocity model at 4.00 Myr……..49
Figure 4.8: Velocity profiles across the continental crust for the zero horizontal oceanic plate
velocity model.…………………………………..…………………………………..……………………..50
Figure 4.9: Strain rate profile across the continental crust for the zero horizontal oceanic plate
velocity model.…………………………………..…………………………………..……………………..50
Figure 4.10: Evolution of the anorogenic continental crust model zoomed into an 800 km by 200
km segment.…………………………………..…………………………………..………………………..51
Figure 4.11: Velocity profiles along the continental crust up to the trench for each model at 1.5
Myr..…………………………………..…………………………………..…………………………………52
Figure 4.12: Displacement of the trench from the initial timestep through time for all of the
models……………………………………..…………………………………..……………………………52
Figure 5.1: Comparison of the zero horizontal plate velocity model to rifting models by Brune et al.
(2014)…………………………………..…………………………………..……………………………….56
Figure 5.2: Comparison of models with zero horizontal plate velocity and models with free oceanic
plate movement by Quinquis et al. (2011) with the models from this study.…………………………57
List of Tables
Table 3.1: Stage pole rotation values for the central and northern Lord Howe Rise.……………….29
Table 3.2: Rheological properties of the reference model.…………………………………………….39
Table 4.1: Reconstructed properties and ensuing extensional velocities of a number of chosen
traces over the Eastern Australian margin and Lord Howe Rise.……………………………………..42
Table 4.2: Major properties of each model.…………………………………..…………………………53

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1. Introduction
The city of Sydney lies on the Eastern Australian margin, a region of relative tectonic quiescence
since the fragmentation of the Eastern Gondwana Cordillera during the Late Cretaceous-Eocene
(Tulloch and Kimbrough, 1989; Gaina et al., 1998; Matthews et al., 2015). This fragmentation
caused the formation of the Lord Howe Rise, the world’s largest continental ribbon and part of the
submerged continent Zealandia. In considering how this occurred, a model of particular
prominence is the notion that fragmentation occurred due to west-dipping Paciﬁc plate subduction
and subsequent back-arc extension in the Eastern Gondwana Cordillera (Schellart et al., 2006;
Whattam et al., 2008; Crawford et al., 2003).
The nature of back-arc extension at active margins is dependent on the interplay between slab
buoyancy, lateral slab forces, trench rollback, kinematic plate movement, gravitational potential
energy and mantle ﬂow (Karig, 1971; Scholz and Campos, 1995; Schellart et al., 2006; Sdrolias
and Mueller, 2006; Rey and Mueller, 2010). Although active margins have the capacity to become
sites of extension, they often display signs of orogeny due to prior contractional tectonics (Russo &
Silver, 1996). This study seeks to explore the evolution of cordilleras during continental
fragmentation in the context of an active margin. This aids our knowledge in continental margins in
general; regions known to be preferential sites for hydrocarbon formation.
Numerical and analogue subduction models have been produced that emphasise alternative
controlling factors in back-arc extension (Schellart and Moresi, 2013; Quinquis et al., 2011). Many
of these models however lack realistic Earth-like rheologies and material geometries or lack
resolution. This study seeks to establish a high-resolution subduction model with realistic
rheological parameters able to explore the nature of back-arc extension in the overriding
continental Cordillera under a variety of conditions. In particular, the study seeks to understand the
impact of kinematic oceanic plate movement, introducing a buoyant mantle wedge and the
Cordillera itself on back-arc extension. Although this model is placed in the context of the Lord
Howe Rise’s fragmentation from Gondwana, it is not to be taken as a simulation but rather a
generic model of the processes involved.

2
2. Background
2.1 Tectonics of the Lord Howe Rise
2.1.1 Current Geologic Setting
The Lord Howe Rise is the worlds largest continental ribbon and part of the submerged continent
Zealandia (Grobys et al., 2008). It extends from the Chesterfield Plateau west of New Caledonia
down to the Challenger Plateau off New Zealand. It is contained within the Southwest Pacific, a
region formed through the fragmentation of Eastern Gondwana and back-arc extensional
processes (Schellart et al., 2006; Matthews et al., 2015).
The Lord Howe Rise has been divided into 3 distinct structural zones from east to west known as
the Lord Howe platform, the Central rift province and the Western rift province (Stagg et al., 1999;
Jongsma and Mutter, 1978). The Lord Howe platform is 1000-1300 m deep, displaying limited
Cretaceous extension or rift depocentre development (Higgins et al., 2015). To its west lies the
central rift province at depths of 1300-1700 m encompassing the the Faust, Eastern Gower and
Moore basins which are a series of rift depocentres. The western rift province displays even larger
rift depocentres (Van de Beuque et al., 2003) and encompasses the Capel, Western Gower and
Monawai basins. This increasing rift development indicates increasing crustal thinning westward
from the Lord Howe platform (Higgins et al., 2015). Immediately around the Lord Howe Rise lay a
number of sag basins including the Middleton basin to its west along with the Fairway and New
Caledonia basins to its east (Lafoy et al., 2005; Exon et al., 2007; Norvick et al., 2008).
The Middleton and Lord Howe basins formed directly to the west of the Lord Howe Rise following
rifting with the Dampier Ridge and are underlain by continental crust (Gaina et al., 1998a). Rock
sample dredging by McDougall et al. (1994) has also revealed the continental crust composition of
the Dampier Ridge. North of the Dampier Ridge lies the Capel and Faust basins along with the
northern section of the Lord Howe Rise comprising the Chesterfield, Kenn and Hellish plateaus
(Van de Beuque et al., 2003) composed of continental crust rifted from Australia (Exon et al.,
2007). South of the Lord Howe Rise lies the non-volcanic, continental Challenger Plateau which
underwent transtension with the Lord Howe Rise causing the formation of the Bellona Trough prior
to formation of the Tasman basin (Gaina et al., 1998a; Wood, 1991).

3
Figure 2.1: Detailed geological map of the Lord Howe Rise region
From: Higgins et al. (2015)

4
Immediately east of the Lord Howe Rise lies the Fairway, Aotea and New Caledonia basins. The
furthest west of these 3 basins is the Fairway basin which which is a north-south trending, 120-200
km wide structural and sedimentary (not bathymetric) basin (Exon et al., 2007). Lafoy et al. (2005)
used deep reflection-seismic profiling and refraction seismic profiling to assert that underlying crust
is continental. Conversely, gravity, magnetic and seismic data used by Van de Beuque et al.
(2003) indicates the basement is either partly oceanic in nature or at the least contains highly
extended continental crust. The Fairway basin has been incorporated into the New Caledonia
basin which lies immediately to its East by a number of authors due to its bathymetry, however
seismic profiles highlights the presence of a basement ridge known as the Fairway ridge between
Figure 2.2: Bathymetry of the Southwest Pacific region.
AUS - Australia, CP - Challenger Plateau, CSB - Coral Sea Basin, DR - Dampier Ridge, LB -
Lau Basin, NC - New Caledonia, NCB - New Caledonia Basin, NFB - North Fiji Basin, NHT -
New Hebrides Trench, NLoB - North Loyalty Basin, NR - Norfolk Ridge, NZ - New Zealand,
SFB - South Fiji Basin, TKR - Three Kings Ridge. Made using GMT software

5
them (Collot et al., 2009; Exon et al., 2007). The New Caledonia bathymetric basin is underlain by
partly oceanic and continental crust (Matthews et al., 2015, Sutherland et al., 2010; Cluzel et al.,
2002; Lafoy et al., 2005) although the sedimentary and structural basin below it is known as the
Aotea basin (Exon et al., 2007) which is continuous with the Fairway basin (Collot et al., 2009) and
itself continues through to the Taranaki basin northwest of New Zealand. East of the Fairway basin
lies the Norfolk basin, underlain by partly oceanic and continental crust (Matthews et al., 2015,
Sutherland et al., 2010; Cluzel et al., 2001) separated by the Norfolk ridge. Collot et al. (2009) and
Sdrolias et al. (2003) contend this ridge may be have been formed by fragmentation of a pre-
existing Early Cretaceous continental arc.
The wider Southwest Pacific region is defined by the presence of volcanic arcs and back-arc
extensional processes. Underlain by oceanic crust these basins largely vary in trend from north-
south to northwest-southeast and include the Loyalty, South Fiji , North Fiji, Norfolk and Lau Basins
as seen in Figure 2.2 (Whattam et al., 2008; Sdrolias et al., 2003; Cluzel et al., 2012). Many of the
surrounding arcs including the New Hebrides, Loyalty-Three Kings, d-Entrecasteaux, Fiji-Lau-
Colville and Tonga-Kermadec display volcanic arc remnants, alluding to the complex subduction-
influenced tectonic history of the region (Gaina et al., 1998a, Norvick et al., 2001; Schellart et al.,
2006; Whattam et al., 2008). The Australian plate is currently subducting northward underneath the
Pacific plate along the New Hebrides Trench whilst the Pacific plate is subducting westward
underneath the Australian plate at the Tonga-Kermadec Trench (Schellart et al., 2006).
2.1.2 Early to Mid Cretaceous
The Early Cretaceous history of Gondwanaland involved substantial west-dipping subduction of the
Pacific plate at its eastern margin. This long-lived subduction was present from approximately 250
Ma (Veevers, 1984 & Bradshaw, 1989) until about 105-99 Ma (Schellart et al., 2006; Laird and
Bradshaw, 2004) creating both an enlarged fore-arc region taking in the Lord Howe Rise and
adjacent Cordilleran orogenic plateau (Veevers, 2006). A result of this widespread subduction was
the initiation of partial melting all along the Gondwana-Pacific margin resulting in the emplacement
of the Silicic Whitsunday Large Igneous Province (Bryan et al., 2000). This emplacement was the
major preceding magmatic event before rifting along the Eastern Australian margin began with
remainders of this Silicic Large Igneous Province recorded in onshore and offshore sedimentary
basins (Bryan et al., 2000).
Subduction cessation followed a major plate reorganisation 105-100 Ma due to one of two potential
proposed factors (Matthews et al. 2012). Firstly, collision of the Hikurangi Plateau with the Eastern
Gondwana trench near the Chatham rise may have caused a choking of the subduction zone

6
resulting in trench retreat and onset of back-arc extension parallel to strike at the point of collision
(Matthews et al., 2015). Timing of the Hikurangi collision is debated with Sutherland and Hollis
(2001) and Davy et al. (2008) proposing collision at 105-100 Ma due to the timing of onset of
extension in New Zealand whilst Seton et al. (2012) and Worthington et al. (2006) propose a later
collision from 86-80 Ma. Alternatively, Bradshaw (1989) and Luyendyk (1995) contend that the
interaction of a mid-ocean ridge system and subduction zone caused the plate reorganisation.
Luyendyk (1995) propose that as a mid-ocean ridge subparallel to a trench moves towards it, the
down-going slab material becomes too young and small to subduct causing cessation of the
spreading process. However, the timeframe for the subduction cessation and initiation of extension
is clouded due to indications of deformation at Marie Byrd Land and southern New Zealand
suggesting that subduction continued beneath the Antarctic Peninsula until 85 Ma (McCarron and
Larter, 1998; Matthews et al., 2015).
Preservation of the Paparoa metamorphic core complexes on the South Island of New Zealand
indicate the presence of extensional tectonics at Gondwana in the Early-Mid Cretaceous period
(Spell et al., 2000; Schulte at al.,2014). Syntectonic granitic pluton presence infers the presence of
magmatism associated with the core complex formation. Dating and analysis of the cooling rates of
lower plate rocks by Spell et al. (2000) and Schulte et al. (2015) indicates continental extensional
tectonics at Gondwana ~116-90 Ma preceding Tasman spreading and conditioning the crust for this
break-up. Further indicating the presence of Early-Mid Cretaceous extension in southern New
Zealand is the formation of the Doubtful Sound shear zone from 106-97 Ma displaying
decompression, cooling and lateral flow diagonal to the trend of the Cretaceous arc. This formation
coincided with creation of the Paparoa metamorphic core complex and was followed by formation
of the Resolution Island shear zone from 95-89 Ma (Klepeis et al., 2016). This extension
propagated wide-ranging heat flow synchronous with Tasman Sea spreading. Metamorphic core
complexes require the exhumation of rocks from below the brittle-ductile transition and significant
amounts of extension to occur (Platt et al., 2014). This extensional episode is supported by the
upper Jurassic-Cretaceous stratigraphy of the Taranaki Basin (Uruski, 2003).
2.1.3 Late Cretaceous to Eocene
Significant plate reorganisation 105-99 Ma (Bradshaw, 1989; Veevers, 2000) caused a change
from direct subduction at the Eastern Gondwana margin to oblique convergence resulting in
cessation of contractional tectonics at the margin. Strike-slip motion along the margin began to
initiate (Sutherland and Hollis, 2001) followed by continental extension and fragmentation of the
Cordillera due to transtensional forces along the margin (Gaina et al., 1998a; Matthews et al.,
2015; Van de Beuque et al., 2003). Eastward movement of the fore-arc resulted in wide-scale

7
continental lithospheric boudinage in some areas with others displaying successful rifting of
continental crust separated by oceanic basin.
Fairway basin opening occurred in the 105-100 Ma behind the Norfolk ridge volcanic arc and away
from the Lord Howe Rise (Uruski and Wood, 1991; Lafoy et al., 2005; Collot et al., 2009). Following
this, rifting began between the Challenger Plateau and Lord Howe Rise forming the Bellona
Trough. The Challenger Plateau rifted from Australia 83 Ma initiating Tasman Sea spreading
(Gaina et al., 1998a). Spreading migrated northwards, with strike-slip motion between the Gilbert
Seamount complex and South Tasman Rise and between the middle and northern Lord Howe Rise
reflected in gravity and seismic reflection data. This northwards migration continued until Chron 28
(61 Ma) at which point the Chesterfield Plateau rifted from the Australian plate and a localised,
continuous spreading centre formed in the Tasman basin (Gaina et al., 1998a; Van de Beuque.,
2003).Through the use of satellite altimetry, magnetic, seismic and geological data, tectonic activity
of the Tasman Sea can be modelled using the motion of 13 micro-continental blocks (Gaina et al.
1998a) rather than through the 2 plate system proposed by Hayes and Ringis (1973) due to the
propagation of rifting from south to north in a number of stages.
The mechanics explaining the cause behind Southwest Pacific extension following the cessation of
Pacific plate subduction under the Gondwana plate are not well constrained with a variety of
models proposed to explain evolution up til the mid-Eocene. These unique models of tectonic
history all seek to explain the extensional forces that caused the deformed continental crust and
pattern of back-arc basins and ridges of the region.
Schellart et. al (2006) emphasises a model of continuous west-dipping subduction from 82 Ma
through til present-day with east-north east directed slab rollback of the Pacific Plate. This slab
rollback is proposed to be 1200 km long accomodating the opening of the New Caledonia, South
Loyalty, Coral Sea and Pocklington basins, along with the Tasman Sea Basin from 80-52 Ma
(Schellart, 2006; Cluzel, 2010). Approximately 50-45 Ma, the eastward-subducting New Caledonia
and northward-subducting Pocklington subduction zones formed, subducting the South Loyalty and
Pocklington basins with their remnants displaying back-arc geochemical signatures (Aitchison et
al., 1995; Cluzel et al., 2001). Models by Whattam et al. (2008), Cluzel (2001) and Crawford et al.
(2003) offer an alternative model without continuous west-dipping subduction. The model of
Crawford et. al (2003) proposes continuous subduction until 55 Ma at which point there is
subduction polarity reversal and formation of an east-directed subduction zone at the extinct South
Loyalty Basin forming the Loyalty-D’Entrecesteaux arc.

8
Conversely, models by Hall (2002), Sdrolias et. al (2003), Matthews et al. (2012) and Seton et al.
(2012) dispute a west-dipping subduction regime in favour of east-dipping subduction. Sdrolias et
al. (2003) proposes the formation of a continuous eastward-dipping subduction zone active from 90
Ma to 45 Ma initiating along the Loyalty-Three Kings Ridge. The extensional forces required to rift
the opening of the Tasman Sea are proposed to be due to the influence of slab pull. East-dipping
subduction in the model of Seton et al. (2012) is supported through the distribution of slab material
imaged by seismic tomography showing continuous slab material extending from the surface to
1000 km depth whereabouts 85 Ma slabs are predicted to be.
The presence of any subduction happening to the east of Australia at all is debated in the models
of Steinberger et al. (2004) and Sutherland et al. (2010). Steinberger et al. (2004) propose a
configuration of the Southwest Pacific which places no plate boundary between the Pacific plate
and the Lord Howe Rise before Chron 20 (43 Ma). This conclusion is drawn through a revision of
Figure 2.3: (A) - (B) show the effect of slab rollback in promoting trench retreat and accomodating back-
arc opening. Note the lithospheric boudinage and multiple rifts at Eastern Gondwana occurring up to 50
Ma forming the Dampier Ridge, Lord Howe Rise and Norfolk Ridge. (C) shows subduction polarity
reversal from west-dipping to east-dipping at 50 Ma in agreement with models of Whattam et al. (2008)
and Crawford et al. (2003).
Edited From: Cluzel et al. (2001)
A
B
C

9
Figure 2.4: Diagrammatic representations for subduction history east of the Lord Howe Rise. Model
1a, based off Schellart et al. (2006) has continuous west-dipping subduction accomodating back-arc
extension of the South Loyalty Basin before the initiation of east dipping subduction consuming the
South Loyalty Basin and causing obduction of the Loyalty Arc. Model 1b is based off Whattam et al.
(2008) and Crawford et al. (2003). The major difference between 1b and 1a is that west-dipping
subduction ceases 55-50 Ma when east-dipping subduction initiates before re-initiating 50-45 Ma.
Pacific plate movement in 1b is also not defined. Model 1c illustrates the model of Sdrolias et al.
(2003) and Hall (2002) with continuous east-dipping subduction from 85 Ma to 45 Ma. Model 2
represents the models of Steinberger et. al (2004) and Sutherland et al. (2010). There is no proposed
subduction east of the Lord Howe Rise and it is proposed to be part of the Pacific plate from 83-45
Ma.
C - Colville Arc, FB - Fairway Basin, L - Loyalty Arc, LH - Lord Howe Rise, N - Norfolk Ridge, NB -
Norfolk Basin, NLB - North Loyalty Basin, PAC - Pacific Plate, PT - Proto-Tonga-Kermadec Arc, SLB
- South Loyalty Basin, V - Vitiaz Arc
From: Matthews et al. (2015)

10
existing global plate motion to account for the bend in the Emperor-Hawaii seamount chain.
However, this bend can also be explained through the interaction of plume tilt and the lateral
advection of the plume source (Hassan et al., 2016). Furthermore, continued subduction during
this period would expect to form wide-scale volcanic arcs in the region. However, there is a lack of
Late Cretaceous to mid-Eocene volcanic arcs present in the Southwest Pacific. Conversely, arc
remnants from the west-Dipping late Palaeozoic subduction beneath Eastern Gondwanaland are
widespread throughout Australia and Zealandia (Matthews et al., 2015). However, rapid eastward
flow of the upper asthenosphere during stretching however may have prevented uplift of
metasomatized asthenosphere (controlled by the mantle wedge) to occur, causing no volcanic arc
to form (Cluzel et al., 2010). The extensional forces that occurred in the region at the time are
proposed by Steinberger et al. (2004) and Sutherland et al. (2010) to be independent of subduction
zone and back-arc basin processes.
An alternative explanation for extensional forces is that orogenic collapse could have precipitated
rifting of the Lord Howe Rise (Rey and Mueller, 2010). Subduction of a west-dipping early
Cretaceous magmatic arc promoted frictional and viscous stresses along the Eastern Gondwana-
Pacific margin with the capability to promote surface uplift and the formation of Cordilleran-type
orogens. Following the plate reorganisation 105 Ma and the cessation of this west-dipping
subduction, dissipation of the forces promoting orogenic uplift occured. Orogenic collapse and
trench rollback with the aid of a buoyant mantle wedge are proposed by Rey and Mueller (2010) to
have caused the rifting of the Lord Howe Rise. This low density, buoyant upwelling mantle
promotes tensional forces on the overriding plate aiding the collapse. Weakness and rifting can
occur in orogenic belts compared to cratons due to higher heat due to shallow asthenosphere
convection aided by water from the underlying subducted slab (Hyndman et al., 2005). Sutherland
et al. (2010) also proposed that preservation of metamorphosed oceanic crust in the Gondwana
slab preconditioned chemical weakness in the thickened Gondwana crust. In contrast, Crawford et
al., 2003 emphasises the possibility of mantle plume influence on the region in which it thermally
weakens the rifted margin allowing spreading to propagate and micro continents to more easily
slice off (Gaina et al., 2003).
2.1.4 Eocene to Present Time
A secondary change in major plate motion occurred 45 Ma due to fusion of the Indian and
Australian plates following the collision of India with Asia (Veevers, 2000). This resulted in collision
of the Loyalty Arc with New Caledonia causing blockage of the subduction zone and obduction of
ophiolites onto New Caledonia (Aitchison et al. 1995). Due to this ‘choking’ event 45 Ma east-
dipping subduction was forced to relocate to the east of the Loyalty-Three Kings Ridge as well as

11
change subduction polarity to a west-dipping subduction zone according to Sdrolias et al. (2003).
However, dating of bonninitic lavas in the Mariana-West Philippine region show emplacement
52-48 Ma indicating a potential earlier subduction date for the region (Crawford et al., 2003).
Further bonninitic presence beneath the New Caledonia ophiolite and the forearm region of the
Tongan arc indicate a potential change in Plate motions in the Southwest Pacific at 55 Ma rather
than 45 Ma.
In contrast, Schellart et al. (2006) proposes an eastward dipping subduction zone occurring at a
newly formed New Caledonia subduction zone initiating 50 Ma in tandem with the modelled
continuing west-dipping subduction zone to its east. East-dipping subduction caused westward
rollback consuming the South Loyalty Basin and resulting in thrusting of the Poya and Pouebo
terranes (Aitchison et al., 1995), southwestward obduction of fore-arc ophiolite onto New
Caledonia (Cluzel et al., 2001) and opening of the North Loyalty Basin. This subduction is
proposed to continue until 27 Ma at which point the subducting slab detaches, becoming a remnant
subduction zone (Schellart et al., 2006). Subsidence of the New Caledonia Basin 36 Ma is claimed
to have formed due to lithospheric delamination and proposed initiation of the Australia-Pacific
boundary (Sutherland et al., 2010).
650 km of eastward slab rollback of the Pacific plate along the outer Tonga-Kermadec trench from
the Oligocene to the early Miocene begun in the north at the proto-Tonga section before migrating
southwards to the proto-Kermadec section (Schellart et al., 2006). This caused asymmetric
opening of the South Fiji and Norfolk Basin’s and propagation of the Three Kings and South
Loyalty ridges (Schellart et al., 2006; Mortimer et al., 2007). Continuing roll-back of the Tonga-
Kermadec Trench 400 km opened the North Fiji basin due to clockwise rollback of the New
Hebrides trench and the Lau-Havre basins from the Miocene to present (Schellart et al., 2006;
Hathway, 1993).
2.1.5 Comparable Global Environments: The North West Shelf and the Basin and
Range Province
The continental fragmentation of the North West shelf of Australia and subsequent drift,
amalgamation and accretion of continental terranes are proposed to have formed modern day Asia
(Metcalfe, 2013). For the Lhasa terrane (found in modern day Tibet), it’s Triassic separation from
Australian Gondwana is said to have occurred due to south dipping subduction of Meso-Tethys
oceanic crust underneath it, resulting in back-arc extension and continental fragmentation as seen
in Figure 2.6 (Metcalfe, 2013). This is supported by the presence of a Permian eclogite belt and
island arc basalts identified by Yang et al. (2009) in the Lhasa block seen in Figure 2.5.

12
Above - Figure 2.5: The modern
day location of the Lhasa terrane
following fragmentation, movement
and accretion. Locations of arc
basalts and eclogites are
presented, emphasising its
potential back-arc source.
From: Metcalfe (2013)
Left - Figure 2.6: A representation
from 250-145 Ma of the proposed
back-arc extension and
fragmentation of the Lhasa block.
From: Zhu et al. (2011)

13
The topography of the Basin and Range province displays narrow, highly faulted Cordilleran
orogens and flat valley and basin regions (Stewart, 1978 and Eaton, 1982). This topographic
variance reflects the historic normal faulting present in the region causing the formation of horst
and graben structures indicative of extensional tectonics. Initially, subduction of the Farallon plate
underneath the North America plate caused the formation of a Cordilleran orogen along western
North America (Coney, 1987., Sigloch and Mihalynux, 2013). Interaction of the Pacific-Farallon
spreading ridge with the subduction zone (Brothers et al., 2012) caused a Cenozoic slowdown of
Farallon plate subduction and slab removal by 20 Ma. This then caused slab-driven trench retreat
at the North American-Farallon margin caused extensional forces and formation of the Basin and
Range province inland in the North American continental crust (Schellart et al., 2010 and
Humphreys, 1995). The presence of metamorphic core complexes in the North America Cordillera
(Methner et al., 2015) indicates the presence of significant extensional tectonics in the region.

14
2.2 Continental Extension at Active Margins
Continental extension and rifting is a dynamic process of integral importance to plate tectonic
theory. Rifting and break-up of existing continental crust allows the formation of ocean basins and
the onset of seafloor spreading to occur. The nature of how this occurs in the context of an active
margin is critically important in understanding the Late Cretaceous-Eocene geological evolution of
the Lord Howe Rise.
2.2.1 Fragmented Crustal Block Types and Composition
Fragmentation at the continental margin due to extensional forces causes the formation of a
number of differing crustal block types depending on the extent to which this rifting occurs and the
relationship of the rifted block with the continent. Rifting preferentially occurs in continental crust at
margins rather than the nearby oceanic crust due to its greater rheological weakness (Vink et al.,
1984).
Peron-Pinvidic and Manatschal (2010) utilises the term ‘micro-continent’ as being an isolated piece
of continental crust entirely rifted from the main continent. This differs to a ‘continental ribbon’
defined by Lister et al. (1986) in which thinned continental fragments remain attached to the intact
plate. These crustal fragments retain the general composition of the continental plate from which
they rifted (Tetreault and Buiter, 2014). However, extensional episodes are able to highly modify
the stratigraphy of the crust and can also promote the presence of magmatism, dyking and
hydraulic fracturing. Christensen and Mooney (1995) determined the average crustal density of
continental fragments to be 2.81g cm-3, which, despite their thinned nature, is similar to that of
typical continental crust.
2.2.2 Active and Passive Rifting
Continental rifting is able to occur through two processes known as active rifting and passive
rifting. Active rifting involves the insertion of a mantle plume up into the base of the continental
lithosphere. This upwards flow of hot mantle material causes thinning and failure of the lithosphere.
Failure occurs due to tensional force build-up from crustal doming and tractional forces occurring at
the base of the lithosphere (Neugebauer, 1978). Turcotte and Emerman (1983) suggest that hot
mantle rock from the asthenosphere reaches the base of the continental crust through diapiric
penetration. Active rifting is illustrated in the evolution of the East African rift, which has developed
due to mantle convection and hotspot magmatism (Regenauer-Lieb et al., 2008 and Corti, 2009).
Whilst plume impingement is the major factor in active rift formation, it also has an ability to

15
influence other rifted margins through heating the lithosphere and allowing migration of spreading
centre’s to weaker zones (Mueller et al., 2001).
Passive rifting differs from active rifting in that it is not dependent on dynamic processes in the
mantle. Passive rifting occurs due to extensional forces caused by plate boundary forces as well as
lateral variations in gravitational potential energy (Regenauer-Lieb et al., 2008). These extensional
forces cause thinning and rifting of the lithosphere ultimately causing a rise in the asthenosphere
upwards (Huismans et al., 2001). This ascension of the asthenosphere occurs as a consequence
rather than cause of rifting.
2.2.3 Causes of Rifting at Active Margins
Continental crust at active margins have the capacity to undergo extension and form back-arc
basins. These basins often form behind island arcs and can also be know as ‘marginal basins’ due
to their proximity to continental margins and because they are produced through interactions at
convergent plate margins (Karig, 1971). However, not all subduction zones and convergent
margins cause back-arc extension with the right conditions necessary for it to occur.
Interaction between the buoyancy force of the down-going slab, forces on the slab resisting lateral
motion, trench rollback, kinematic plate motion, gravitational potential energy and mantle effects
are all proposed to play varying influences on the nature of back-arc extension. Karig (1971), first
emphasised the link between mantle processes and back-arc extension by proposing an upwelling
diapir of shear-heated mantle material rising from the asthenosphere to the overlying lithosphere.
Sleep and Toksov (1971) furthered the idea of mantle influence through suggestion of secondary
convection cells behind arcs causing rifting to occur. These models fail to explain however why it is
that only a select number of arcs display active back-arc spreading.
Forsyth and Uyeda (1975) proposed the term ‘trench suction’ to explain a vertical downwards force
caused by subducting slabs that drags the continental crust with it (Chase, 1978). This vertical
suction force occurs due to slab pull forces becoming unbalanced due to high subduction
resistance of the subducting slab (Scholz and Campos, 1995; Shemenda ,1993). Scholz and
Campos (1995) also highlight the influence of ‘sea-anchor force’ in which hydrodynamic resistance
to lateral movement of the subducting slab through a viscous fluid then results in lateral movement
of the overriding continental plate from the trench.
Influence of toroidal mantle flow in the overriding plate has also been proposed to produce back-
arc extension in narrow subduction zones (Schellart and Moresi, 2013; Sternai et al., 2014; Chen

16
et al., 2016). Chen et al. (2016)
used dynamic laboratory models
in combination with a
stereoscopic Particle Image
Velocimetry technique to map the
deformation of the overriding
plate as well as subduction-
induced mantle flow under and
around the overriding plate. Their
results indicate maximum
extension 300-500 km from the
trench due to toroidal mantle flow
migrating from the sub-slab
region around the lateral slab
edge to the mantle wedge
causing basal drag force at the
base of the overriding plate.
Subducting slabs are colder and
denser than the surrounding
mantle, resulting in a negative
buoyancy within its environment.
Schellart et al. (2006) proposed
that this negative buoyancy of the
down-going slab causes a
backwards sinking at an oblique
angle to the initial dip of the slab. This then results in seaward retreat of the subducting slab and
extension of the overriding plate as it collapses towards the trench. As the overriding plate moves
towards the trench the relative slip rates between the two plates increases resulting in further
stress in the back-arc region. This causes a feedback mechanism that propagates steady back-arc
spreading (Hashima et al., 2008).
Sutherland (2010) argues that in the context of the New Caledonia Basin the slab undergoes
delamination before rolling back. Gravitational instability of the lithospheric root causes
delamination of the lower crust, resulting in sinking of the slab before slab rollback and trench
retreat began to occur as seen in Figure 2.7. Sutherland (2010) further emphasises the impact of
resulting upwards mantle flow driving further back-arc extension.
A
B
C
Figure 2.7: Diagram showing how delamination of an oceanic slab
results in back-arc extension in the overriding continental crust.
Edited from: Sutherland et al. (2010)

17
Although back-arc basins are seen in environments worldwide they do not occur at every
subduction zone margin. Sdrolias and Mueller (2006) present a number of conditions necessary for
back-arc extension to occur at a convergent margin.
1. Back-arc basins may only develop when the age of the subducting oceanic lithosphere is
over 55 million years.
2. The average dip angle of the subducting slab must be greater than 30 degrees.
3. Back-arc basin formation must be preceded by absolute motion of the plate away from the
trench so that ‘accomodation space’ is created between the subducting and overriding
plates. Following this, accomodation space continues to form through rollback of the
subduction hinge regardless of overriding plate motion.
Back-arc extension is quite episodic in nature with Lister et al. (2001) proposing this to be due to
collision of remnant volcanic arcs and ocean plateaus with the subduction zone. Conversely,
Schellart et al. (2006) emphasises the interaction of a down-going slab with the upper and lower
mantle as being responsible for this episodic nature.
Back-arc basins also have the potential to form due to orogenic gravitational collapse and rifting,
seen in Figure 2.8 (Rey and Mueller, 2010). Differences in crustal thicknesses can cause lateral
variations in gravitational potential energy across a plate due to locations of upwelling mantle.
These variations cause the lateral movement of crustal material in compensation (Artyushkov,
1973). Whilst these variations also have the potential to drive contraction, in the context of back-
arc basin formation they promote extension in regions of thickened crust (Rey et al., 2001).
Subduction at a convergent margin can cause lithospheric thickening and surface uplift in the
overriding plate which is sustained by frictional and viscous forces from compressional stress due
to plate motions and basal shear stresses (Sutherland, 2010; Rey and Mueller, 2010; Rey et al.,
2001). Therefore, once the forces promoting convergence and thus lithospheric thickening
decrease enough that the forces caused by gravitational potential energy overcome it, extension in
the region may occur. Conductive heating of the overriding plate structurally weakens it making it
more susceptible to these rifting processes (England et al., 1988). Rey and Mueller (2010) also
propose the existence of a buoyant mantle wedge in subduction zones as helping to drive orogenic
gravitational collapse and the formation of back-arc basins. This wedge produces horizontal
extensional stresses in the overriding plate which couple with forces due to gravitational potential
energy to promote extension. Rey and Mueller’s (2010) numerical modelling using Ellipsis shows
that the extent of trench retreat, microcontinent fragmentation and crustal boudinage is proportional
to the buoyancy of the mantle wedge.

18
Hyndman et al. (2005) proposes the idea that deformation during collisional orogeny concentrates
among weak former back-arcs. This deformation causes the formation of ‘mobile mountain belts’ at
the site of these former back-arcs because of their susceptibility to plate boundary forces due to
their structural weakness and hot temperature. This is an interesting concept to pair with the notion
of divergence due to gravitational collapse as it provides a cyclical nature of back-arc basin and
orogenic mountain building processes.
The phenomenon of ridge jump is a major controlling factor on the initiation of rifting. Ridge jump
occurs when a nearby spreading centre migrates landwards onto the thinned continental margin
(Mittelstaedt et al., 2011; Misra et al., 2015). Strength of the continental lithosphere at continental
margins is weaker than the neighbouring oceanic crust allowing the continental crust to become a
potential zone for preferential rifting (Mittelstaedt et al., 2011). Re-rifting can occur in continental
margins younger than 25 My old if they interact with a mantle plume stem and cause the formation
of microcontinents (Mueller et al., 2001). Thermal heating by the plume causes spreading centres
Figure 2.8: The inﬂuence of a buoyant mantle on lithospheric extension. (A) shows the initial boundary
conditions and composition of the model. (B) - (D) show an increasing density and thus decreasing
buoyancy mantle wedge. As this density increases, the mantle wedge surface extension decreases at the
overriding plate. (E) has no buoyant mantle wedge and is underlain by typical asthenosphere. Here the
slight surface extension can be attributed to gravitational collapse. The densities in this model are
calculated at 1000°C and in the mantle wedge a melt fraction of 8% is assumed.
Modiﬁed from: Rey and Mueller (2010)
A
B C
D E

19
to jump to zones of weakness within
the continental crust on the landward
edge of the rifted margin. These ridge
jumps occur both more often and
rapidly towards the plume under high
magmatic heating (Misra et al., 2015).
Mueller et al. (2001) illustrates the
asymmetric nature of spreading that
occurs due to the plume interaction
as seen in the East Tasman Plateau
and Gilbert Seamount Complex
(Figure 2.9). Numerical modelling by
Brune et al. (2014) indicates that the
rift migration is controlled by lower
crustal flow, which is itself a function
of the thermal properties, strain rate
and crustal composition. Yamasaki
and Gernigon (2010) also proposed
the notion of magmatic underplating
as a cause for re-rifting. This can
cause thermal weakening of the
lithosphere without actually reaching
the surface; changing rock chemistry
and increasing levels of heterogeneity
(Regenauer-Lieb et al., 2008).
2.2.4 Rift Architecture and
Evolution
Rift evolution and formation is heavily
dependent on crust-mantle interaction
and external forces. These
parameters influence the mechanism
of upper crustal brittle deformation and lower crustal ductile flow ultimately causing causing
variance in rift architecture.
Figure 2.9: The influence of a mantle plume on rift
migration. The mantle plume stem on Plate 1 causes the
formation of a graben due to tensional forces on the
overriding plate as seen in C. The centre of rifting then
jumps over to the hotspot in D as it has become a weaker
zone in comparison to the existing rift. This new rifting
causes the formation of another oceanic spreading centre
on plate 1 in E and the formation of a microcontinent.
Subsequent rift migration is visible in E as the ridge
migrates to the hotspot location.
From: Mueller et al. (2001)

20
Peron-Pinvidic and Manatschal (2010) propose a system for rift evolution over time. Initial
stretching produces pure shear and deformation of the continental crust which is largely
delocalised aside from areas with pre-existing weaknesses (Lavier and Manatschal, 2006). This is
followed by the thinning phase in which deformation localises along a number of major faults at the
edge of strong crustal bodies resulting in the formation of continental ribbons at different places
within the lithosphere. After thinning to a point where continuous ductile layers are no longer
present in the crust, coupling of brittle layers and large scale detachment faults form in what is
known as the exhumation phase. Following this exhumation, strength softening by input of magma
contributes to break-up of continental crust and the onset of seafloor spreading.
Whilst the many origins for rifting have been discussed above, other internal factors influence the
style and pattern of how it occurs. Boundary conditions, rheology of the crust and mantle, coupling
effects as well as heterogeneities in the crust and mantle all influence lithospheric strength and rift
architecture. These factors do not operate individually, but are interlinked and can contribute to one
another during the rifting process.
Strain rate and extensional velocity directly impacts on the symmetry of rifting and the margin
width. Brune et al. (2014) used thermo-mechanical modelling to indicate that faster extension
results in wider margins displaying strong asymmetry. Wide margins rely on the ability for ductile
Figure 2.10: Diagram highlighting rift migration occurs in varying extensional velocity environments. Higher rift
velocity leads to shallower advection of the isotherm heating the lithosphere. The lower crustal flow seen in g, h, i
counteracts the crustal thinning due to faulting allows the phase of rift migration to increase, ultimately leading to
wider margins.
From: Brune et al. (2014)

21
lower crust to flow into the area of fault termination. This is supported by high extension rates
causing heat advection from the mantle to shallower depths. This leads to heating of the lower
crust, forming a weaker and larger point of deformation and causing higher rates of rift migration.
Rift migration occurs due to softening at the tip of the active fault and strengthening at its footwall
which generates a lateral strength gradient (Figure 2.10). These fast-extension scenarios produce
less coupling between the upper continental crust and continental mantle and the detachment
depth rift-related faults in the upper crust is less than those in slow-extension settings (Nemcok et
al., 2012). In contrast to Brune et al. (2014), Huismans and Beaumont (2007) used lithospheric-
scale models with a strong lower crust and a zone of inherited weakness to determine that rifting
asymmetry was greatest for low extensional velocities. Strain softening resulted in localisation of
shear zones causing preferential weakening with lithospheric detachment and asymmetric
spreading. These models resulted in initially wide rift modes before transitioning to late stage
narrow rifts with significant mantle upwelling.
Temperature of the crust and mantle greatly affect the nature of rifting and the impact of
extensional velocity on this has been previously stated. Thermal properties affects viscosity of the
upper and lower crust and the depth of the brittle-ductile transition. Old, cold and brittle crust tends
to display narrow rifting, whilst younger, hotter and more ductile crust displays patterns of wide
rifting (Brune et al., 2014). However, regions with thinned lithosphere and thus hotter crust resulted
in symmetrical rifting with no rift migration, whilst a model with a thicker lithosphere and cooler
thermal profile resulted in rift migration and asymmetric rifting due to greater coupling effects
(Brune et al., 2014). This echoes Précigout et al. (2007) who found that a low temperature, high
strength mantle lead to greater crust-mantle coupling and the localisation of strain at certain points
with narrow rifting and lithospheric necking. Conversely, lower strength mantle models with low
crust-mantle coupling show patterns of wide rifting with more evenly spread strain (Nemcok et al.
2012). Gueydan et al. (2007) found that mantle temperatures greater than 700°C produced a sharp
reduction in upper mantle strength. In the context of a back-arc basin, the initial location for rifting
does not necessarily occur in the region of highest stress accumulation but in the region of greatest
lithospheric weakness which is mainly controlled by temperature (Hashima et al., 2008). However,
high strain locations can become affected through shear heating (Regenauer-Lieb et al., 2008;
Kaus and Podladchikov, 2006).
Aside from major variations in the crust-mantle composition, viscosity of the lithosphere can be
affected by both thermal properties as stated above, and grain size differences. High strain can
result in reduction of grain size within ductile shear zones due to dislocation creep and
recrystallisation (Brune et al. 2014). This lowering of grain size creates viscosity reduction within
the region and promotes diffusion creep (Austin & Evans, 2009).

22
Heterogeneities within the lithosphere can occur due to pre-existing zones of weakness along with
the presence of melt and serpentinisation (Buck, 1991; Ziegler and Cloetingh, 2004; Huismans and
Beaumont, 2007; Regenauer-Lieb et al., 2008). Heterogeneities can cause strain localisation in
these zones of greater weakness (Peron-Pinvidic and Manatschal, 2010) and may cause the
presence of rift asymmetry (Ziegler and Cloetingh, 2004). Pre-existing zones of weakness can
appear in the lithosphere due to pre-existing faults and regions of differing crustal composition.
Numerical modelling by Brune et al. (2014) shows how the presence of a pre-existing fault can
allow asymmetric mantle upwelling and asymmetric lower crustal weakening. Mantle
decompression and heating can cause the production of melt, which influences rheology by
weakening parts of the crust (Buck, 1991). This melt can manifest itself through underplating and
intraplating of basaltic magmas as found by Yamasaki and Gernigon (2010) greatly increasing
heterogeneity of the lithosphere. Buck (1991) also emphasises the impact of serpentinisation by
which water influx into mantle rock produces localised weakening.
2.3 Knowledge Gaps
When considering the fragmentation of the Lord Howe Rise through the prism of back-arc
extension, a number of knowledge gaps appear that this study will seek to address. Although
geological evidence allows an understanding of the region as it lies today, it is through modelling
that an understanding of how it formed through time can be determined. Plate reconstructions have
constrained kinematic movement of the plates, however current literature lacks a cross-sectional
modelling approach on the crust-mantle scale which can help form a more refined view of the
region. Therefore through creating a high resolution model of subduction and back-arc extension
with realistic rheological parameters, this study seeks to explore a few key questions:
1. Why is the Eastern Australian margin so narrow when compared to the hyperextended Lord
Howe Rise?
2. What role does the Eastern Gondwana Cordillera play in extension in the margin?
3. How does movement of the subducting oceanic plate influence back-arc extension?
4. Finally, what role does a buoyant mantle wedge play in promoting back-arc extension with
the presence of a subducting slab? This slightly expands on the work by Rey and Mueller
(2010) by including the oceanic slab material within the model.

23
3. Methodology
In order to address the gaps in knowledge previously addressed, this study sought to analyse
back-arc extension and subduction under a range of conditions. This was achieved through
thermo-mechanical modelling of the processes in question. In order to compare and contrast
model outputs to the Lord Howe Rise, a palinspatic reconstruction of the Eastern Gondwana region
was developed for comparison to the models.
3.1 Developing a Palinspatic Reconstruction of Eastern Gondwana
Reconstruction of deformed crust to a pre-extensional state was performed in order to recreate the
environment of pre-fragmentation Eastern Gondwana. This crustal reconstruction to pre-
extensional thickness was achieved through partially following the methodology of Williams et al.
(2011). This allowed for the calculation of extensional velocities across the Tasman Basin and
within the deformed crust through geological time. As an overview, the following are the major
steps that I undertook in order to reconstruct the crustal thicknesses, margin widths and
extensional velocities:
1. Formation of crustal thickness, Bouguer gravity and total horizontal derivative maps for the
region.
2. Ascertaining the location of the Unstretched Continental Crust Limit (UCCL) and Continent
Ocean Boundary (COB) for the Eastern Australian margin and the locations of both COB’s
for Zealandia.
3. Emplacement of small circles in GPlates representing the direction of internal deformation
and extension at the Lord Howe Rise and Eastern Australia pre and post breakup.
4. Using Generic Mapping Tools software, determine the width, thickness and cross-sectional
area of the deformed crust along the small circle path representing direction of extension.
5. Assuming a rectangular initial crustal proﬁle, determine the initial width for a 60 km thick
crust.
6. Determine the times of initial extension, break-up and cessation of movement to calculate
the extensional velocities during the periods of pre-breakup internal deformation and post-
breakup movement of the Lord Howe Rise away from the Eastern Australian margin.
3.1.1 Crustal Thickness Map

24
The crustal thickness for the Eastern Australian margin was found through adding bathymetry and
topography data to Moho depth estimates (Figure 3.1). Aitken (2010) estimated the depth to the
Moho through his Moho geometry gravity inversion experiment (MoGGIE). This model is based on
constraining inversion of free-air gravity data through seismic Moho depth results along with
assumption of a laterally non-uniform lithosphere. Bathymetry and topography data with a
resolution of 9 arc seconds was taken from Whiteway (2009) and formed through a joint project
between Geoscience Australia and the National Oceans ofﬁce.
A crustal thickness grid produced by Grobys et al. (2008) was used for the majority of the studied
region in Zealandia. This grid was produced through combining new and published 2D and 3D
crustal thickness models constrained by satellite altimetry, free-air gravity, seismic and sedimentary
thickness data (Figure 3.2). This grid does not cover the entire Lord Howe Rise and Zealandia
region due to sparse data coverage in a number of regions including the Macquarie, Colville and
Kermadec ridge systems. The data was gridded with a spacing of 3 x 3 arc minutes and extends
from 175°W to 157°E and from 57°S to 17°S. For areas of interest outside of this region the global
CRUST2.0 grid made by Bassin et al. (2000) is utilised as per the methodology of Williams et al.
(2011). The data is plotted on a 2° x 2° grid which makes it far less detailed than the grid provided
by Grobys et al. (2008) however it only minimally overlaps with the studied region of interest and
thus is useful for the reconstruction.
3.1.2 Bouguer Gravity and Total Horizontal Derivative Maps
Mapping the Bouguer gravity results for the region of interest seen in Figure 3.4 was done through
using the global high resolution grid of Bouguer gravity anomalies created by Bonvalot et al. (2012)
as part of the Bureau Gravimetrique International. This dataset was the ﬁrst to account for a
realistic Earth model that considered contributions of heterogeneous surface masses and is at a 1
x 1 arc minute resolution.
The total horizontal derivative (THD) of a grid emphasises regions with the highest rate of change
of data values. Through calculating the total horizontal derivative of the Bouguer gravity data this
allows us to more easily discern the point of transition from deformed to undeformed continental
crust along Eastern Australia as seen in Figure 3.3. In order to do this, a fast-fourier transformation
was done to the Bouguer gravity data so that only the long wavelength changes in Bouguer gravity
data remained. The gradient of the grid in both the vertical and horizontal directions was then
found. If vertical gradient = v and the horizontal gradient = h, then the total horizontal derivative of
the grid is equal to (v2 +h2)1/2.

25
3.1.3 Determining the Unstretched Continental Crust Limit and Continent Ocean
Boundaries
In order to reconstruct the paleothicknesses of the Eastern Australian margin and Lord Howe Rise,
parts of the methodology of Williams et al. (2011) was used. In their reconstruction of the
Australian-Antarctic margin, Williams et al. (2011) employed a method by which the boundary
Figure 3.2: Crustal thickness map of the Lord
Howe Rise region. The COB’s on either side are
in white with the small tracers representing plate
movement in brown.Figure 3.1: Crustal thickness map of the Eastern
Australian margin. The COB and UCCB are in
white with the small tracers representing plate
movement in purple.

26
Figure 3.3: Total horizontal derivative of the Bouguer gravity map for the Eastern Australia-
Lord Howe Rise region. The COB’s and UCCB for the region are in white

27
Figure 3.4: Bouguer gravity map for the entire Eastern Australia-Lord Howe Rise region. In
white are the COB’s and UCCB for the region. The small traces for plate movement are
represented in purple for the Eastern Australia margin, grey for Tasman Sea and brown for
the Lord Howe Rise.

28
between the unstretched continental crust and stretched continental crust was found along with the
continental crust and oceanic crust boundary. The locations of these boundaries were constrained
through geophysical and geological data of the region. In this study, the crustal thickness, bouguer
gravity and total horizontal derivative maps was deemed enough to delineate the boundaries.
3.1.3.1 Eastern Australian Margin
The Bouguer gravity map shows a transition from low to high values as continental crust transitions
to oceanic crust. This occurs due to shallowing of the Moho from thick continental crust through to
thinner oceanic crust. As stated above, the long-wavelength total horizontal derivative map best
emphasises this change in bouguer gravity. The boundary for the Eastern Australian UCCL lies at
the western edge of the high THD value points. This edge is where the Bouguer gravity no longer
experiences a high of rate of change and thus the Moho no longer experiences signiﬁcant depth
variance. This indicates a region of homogenous crustal thickness. For the Eastern Australian
margin, analysis of the Bouguer gravity and crustal thickness along with high-resolution maps by
Sandwell and Smith (1997) provides a fairly uniform outline for the boundary of continental and
oceanic crust. This was constrained by comparison to previous continental-oceanic crust
boundaries provided by Van de Beuque et al. (2003) and Brown et al. (2003). These comparisons
also helped to assure that the eastward margin lay landward of the ﬁrst magnetic isochron values
which delineate oceanic crust. For the southern region of the margin the COB lies landward of the
East Tasman region following the trace of Brown et al. (2003) owing to poor ability to constrain the
boundary with the Bouguer and crustal thickness maps of this study.
3.1.3.2 Lord Howe Rise
For the Lord Howe Rise region the concept of UCCL does not apply as the entire continent region
has been assumed to have undergone a degree of deformation from its initial thickness and
therefore the COB for both sides of the Lord Howe rise are taken. For the western COB tracing the
boundary between the Lord Howe Rise and the Tasman Basin, a similar method was employed as
in the eastern COB for the Eastern Australian margin. Along with the Eastern Australian margin, the
COB extended northwards up to the Cato trough. The southern margin of the trace extends to the
Challenger Plateau however does not continue through to the Bellona Trough. The eastern COB of
the Lord Howe Rise is not as easily constrained due to varying interpretations of the crust type and
evolution. For this study the New Caledonia basin is assumed to be completely underlain by
continental crust as proposed by Lafoy et al. (2005). By assuming the basin is entirely underlain by
continental crust the entire transect (for which cross-sectional area will be taken for) across the
Lord Howe Rise can be reconstructed to pre-deformation levels. The eastern margin of the Norfolk

29
ridge was taken to be the COB for the southern portion of the region. Proposed opening of the
Norfolk basin occurred from the Oligocene to early Miocene (Schellart et al., 2006; Mortimer et al.,
2007) due to rollback of the Tonga-Kermadec Trench. This occurred after the major Cretaceous
extensional episode which caused Gondwana fragmentation. As it occurred following the major
extensional episode this study concentrates on and the fact that it is underlain by both highly
thinned continental crust and oceanic crust (Matthews et al., 2015, Sutherland et al., 2010) the
eastern margin of the Norfolk ridge can be taken as a good approximation for the COB. The
boundary with exclusively oceanic crust is impossible to ascertain with absolute certainty however
due to sparse data coverage in the region. For the northern portion of the region the COB is taken
as the point of contact of the Norfolk ridge/New Caledonia with the South Loyalty basin.
3.1.4 Stage Pole Rotation of the Lord Howe Rise Relative to Australia
GPlates is an interactive tool that allows users to visualise and edit global plate movements. Within
GPlates, the small circle function was utilised in order to visualise the stage pole rotation of plate
movement. The Central and Northern Lord Howe rise are taken as two separate plates and the
stage poles for these were taken from the dataset of Mueller et al. (2016) shown in Table 3.1.
Stage pole rotation for the central Lord Howe Rise is the same as that provided by Gaina et al.
(1998) whilst for the northern Lord Howe Rise the longitude and latitude values differ by less than
3°. These rotations incorporate movement during continental extension followed by break-up.
One rotation for each plate was used for ease of use rather than splitting it into two rotations for
internal deformation and post-breakup movement. The direction of internal deformation and rifting
movement differ by a minute amount with the ultimate differences in results inconsequential.
3.1.5 Reconstruction of the Paleo-crust
Using the Generic Mapping Tools software, small circle profiles transecting continental crust for
both the Eastern Australian margin and the Lord Howe Rise (this excludes the profiles across the
Tasman basin) were traced separately along the produced crustal thickness grids (Figures 3.1 and
3.2). By doing this, the crustal thickness at discrete points along the profile was deduced and from
Plate Latitude Longitude Timeframe
Northern Lord Howe
Rise
0.75° -45.56° 84Ma-Present
Central Lord Howe Rise 3.27° -42.59° 90Ma-Present
Table 3.1: Stage pole rotation values for the central and northern Lord Howe Rise.

30
that the approximate cross-sectional area of the proﬁle across the grid was established. By
assuming a pre-extensional rectangular crust shape, the width of the paleocrust for a determined
pre-extensional thickness can be determined by calculating Crust Area ÷ Thickness. In this study,
the pre-extensional thickness of the crust is given to be 60 km, the same as that invoked by Rey
and Mueller (2010).
3.1.6 Determining the Extensional Velocity
Extensional velocity was determined separately for the internal continental deformation and for the
post-breakup plate movement at each small circle trace. For internal continental deformation the
velocity was found by dividing the difference between the width of the present day continental crust
and the reconstructed initial width by the difference between the time of initial extension and time
of break-up. This assumes that continental deformation all occurred prior to break-up. For small
circle traces along the oceanic crust in the Tasman basin the extensional velocity was determined
by dividing the distance along the trace by the difference between the time of break-up and time of
plate movement cessation.
Values for time of initial extension were discerned from a number of sources. Extension at the
Doubtful Sound shear zone in New Zealand was found to have begun 106 Ma by Klepeis et al.
(2016) whilst Schwartz et al. (2016) found that extension occurred through orogenic collapse in the
region from 108-106 Ma. Schulte et al. 2014 found that the Pike detachment became active before
116.2 ± 5.9 Ma forming the Paparoa metamorphic core complex. Using these results an
approximate time of initial extension of 110 Ma was utilised for the mid Lord Howe Rise and 105
Ma for the northern Lord Howe Rise. The time of break-up of the Lord Howe Rise and Eastern
Australia was given to be 83 Ma for the mid Lord Howe Rise (Gaina et al.,1998b). For the northern
Lord Howe Rise, break-up between the Dampier ridge and eastern Australia begun 73 Ma (Gaina
et al., 1998a) and break-up in the northerly region of the northern Lord Howe Rise occurred 61.2
Ma (Gaina et al., 1998b). For the locations in between these points a linear progression was taken
in accordance with the nature of the ‘zipper’ opening of the system propagating northwards. The
cessation of spreading for all points was taken as being 52 Ma (Gaina et al., 1998a).

31
3.2 Thermo-Mechanical Modelling with Underworld
3.2.1 Overview
Underworld is a 3D-parallel computational modelling framework for geodynamic processes. It is a
top-level program within a hierarchical set of programs, that allow higher-level programs to build on
the outputs of lower-level programs. Underworld creates time-dependent thermal, mechanical,
tectonic and geodynamic experiments throughout an array of rheologies allowing elastic, plastic
and viscous behaviours (Rey and Mondy, unpublished). It utilises real-world processes such as
radiogenic heating and partial melting, which in turn influences the temperature, densities and
rheologies of the model through time. These features coupled with other processes such as
erosion and sedimentation allow for the simulation of real-world tectonic processes such as
lithospheric extension and mantle convection. (Rey and Mondy, unpublished). Within this study, the
Lithospheric Modelling Recipe developed by Luke Mondy for Underworld will be used to run
models more easily and efficiently.
To do this, Underworld utilises a Lagrangian particle-in-cell finite element scheme. Lagrangian
particles are able to track material deformation and are embedded within a mesh with variables
computed on the mesh nodes (Moresi et al., 2003). In order to generate an output, Underworld
solves the Stokes flow equations and energy advection/diffusion equation. Customisation of these
equations is possible in Underworld by adding things like force terms or constitutive behaviours in
order to create models of varying geophysical complexity (Rey and Mondy, unpublished).
The Stokes flow equations are solved in Underworld on a 2D or 3D cartesian grid. The nature of
Stokes flow is characterised by a small Reynolds number which is the ratio between inertial and
viscous forces interacting within a system. Earth’s geodynamic processes typically involve low
acceleration and inertia whilst dealing with high viscosity materials. This results in a low Reynolds
number and thus the Stokes flow equations become applicable. The flowing material itself is
assumed to be incompressible which helps to ensure the conservation of mass (Rey and Mondy,
unpublished). At a low Reynolds number as seen in Stokes flow, motion is smooth and laminar
whilst it becomes unstable and disordered for high values. As the Stokes flow equations are solved
in time; pressure, velocity, density, viscosity and strain-rate are computed, updated and re-stored
into the particles throughout the grid. This occurs as a continuing process throughout the run-time
of the model (Rey and Mondy, unpublished).

32
3.2.2 General Equations
Note: All equations were taken from Rey and Mondy (unpublished) other than the advection/
diffusion equation which was taken from (Moresi et al., 2003)
Incompressible Stokes Flow
The Stokes equations are solved on a cartesian grid and in a tensorial form the following equation
shows the conservation of momentum:
Where σij is the total stress tensor, τij is the deviatoric stress tensor, pδij the pressure tensor and
fi is the gravitational body force.
The relationship between the deviatoric stress tensor, the viscosity tensor and the strain rate
tensor is given by:
which results in the Stokes equation:
which can then be displayed in vector form to create the momentum equation:
where ∇2u is the velocity gradient, ∇p is the pressure gradient, η is the viscosity, ρg the driving
force and η∇2u is the stress gradient. This stokes equation is then solved on a 2D or 3D mesh with
values for pressure, velocity, density, viscosity, strain-rate and stress continuously updated and
assigned to particles which then advect through the grid (Rey and Mondy, Undated).
Advection/Diffusion
The advection-diffusion equation describes physical phenomena in which energy or other physical
loads are transferred within a physical system due to advection and convection.

33
xi is the spatial coordinates, ui is the velocity, T is the temperature, α is the thermal expansivity, ρ
is the fluid density, g is the gravitational acceleration, λ is the unit vector in the direction of gravity
and κ is the thermal diffusion.
Temperature
Temperature plays a crucial role within Underworld as it is a major influence on density and
viscosity of the system. This is factored in through the following equation:
Where, DF is partial melting, H is radiogenic heat production, κ is thermal diffusion and uz is heat
advection. This equation ensures a conservation of energy within the system and allows for density
and viscosity to become coupled with temperature as in the equations below.
Where, A is the pre-stress factor, n is the stress exponent, E is the activation energy and ἐ is the
strain rate.
Viscosity and Plastic Deformation
Above a particular yield stress, plastic materials become viscous causing plastic deformation. This
is considered in Underworld through the following formula:
Where τ is the yield stress, f(ε) is the strain weakening function (this causes existing fault zones to
become progressively weaker as strain accrue over time), CO is cohesion, μeff is the effective
coefficient of friction, pgz is the confining pressure
The post-yielding viscosity is then:

34
3.2.3 Lithospheric Modelling Recipe
The framework of Underworld allows a large variety of processes on a broad scale to be modelled.
In order to accurately model lithospheric deformation the Lithospheric Modelling Recipe (LMR) was
utilised. The LMR is designed to activate only the relevant segments of the Underworld framework
through the Python coding language which allows for coupled thermo-mechanical lithospheric
experiments to be performed. The input parameters are modiﬁed through a series of stacked XML
ﬁles that form the LMR.

35
3.3 The Reference Model
3.3.1 Outline and Development
The objective of the modelling process was to investigate how various conditions impact on the
nature of subduction of an oceanic slab and back-arc extension in the overriding continental crust.
The material geometry and temperature field for the reference model is the end-result of prior
numerical modelling. This prior modelling was utilised to develop a subducted slab upon which the
reference model could be based. Using modelling to define this slab geometry and temperature
field is necessary in order for isostatic compensation and to minimise model run-time. This
evolution of the reference model is seen in Figure 3.6.
3.3.2 Model Architecture and Rheology
The reference model seen in Figure 3.5 is 1536 km wide and 764 km tall with both a vertical and
horizontal resolution of 2 km. It is comprised of materials including air, sediment, continental crust,
oceanic crust, fault, lithospheric mantle and asthenosphere. Detailed information on the
compositional and rheological properties can be seen in Table 3.2. The top of the model is covered
by air which extends from the the top of the lithosphere to a height of 16 km. The continental
lithosphere overlays the oceanic lithosphere over a 80 km wide faulted region. Thickness of the
continental lithosphere is 141 km in the undeformed region with the crust 41 km thick and the
lithospheric mantle 100 km thick. In the 400 km wide Cordillera the thickness of the lithosphere is
89 km with the crust 60 km thick and the lithospheric mantle thinned to only 29 km.
The oceanic lithosphere is 100 km thick throughout the undeformed region; with the overlaying
crust 6 km thick. This changes at the far-right hand side of the model, with slight thinning due to
deformation caused by the introduction of a thermal anomaly seen in Figure 3.5 (b). Both the
oceanic and continental crustal domains are underlain by the same lithospheric mantle. The
inclusion of a 10 km wide fault between the oceanic and continental lithospheres was done to
encourage subduction. The low viscosity fault acts as a lubricant on the Benioff plane enabling
decoupling to occur during subduction. This minimises the frictional forces between the two plates
which can serve to weaken the subducting slab and produces contractional deformation in the
overriding plate. The oceanic crust also possesses a low viscosity and takes over as the lubricant
between the two plates following subduction of the fault material into the mantle. The presence of
weak crust in nature can occur due to water-sediment interaction and hydration of the crust
whether it be in the form of sediments, hydrothermally altered basalts or serpentinites (Crameri et

36
al.,
Figure 3.5 (a) - top, (b) - bottom: The initial material geometry (a) and temperature ﬁeld (b)
for the reference model. The white lines in (a) represent in descending order the
temperature contours of 400°C, 600°C, 800°C, 1000°C and 1200°C.
1536
764km
Lithospheric Mantle
Fault
Asthenosphere
Air
Oceanic Crust
Continental Crust
Sediments
LEGEND

37
2012). The asthenosphere resides
from the base of the lithospheric
mantle to the bottom of the model at
a depth of 752 km.
A strain weakening function utilised
to create a zone of weakness 12 km
wide and 60 km deep was added to
the far right hand side of the model
from the top of the oceanic crust at
a depth of 3 km down into the
lithospheric mantle. This region of
weakening was included in order to
allow the oceanic lithosphere to
decouple from the right-hand wall
and undergo acceleration due to
slab pull. This was aided by the
thermal anomaly emplaced in the
oceanic lithosphere at the right hand
side of the model. This decoupling
is important as it allows for greater
slab movement due to negative buoyancy.
3.3.3 Initial Temperature Conditions
Equilibration of the temperature field in the model is necessary prior to the coupled thermo-
mechanical modelling over time. Each material within the model contains a unique set of thermal
properties. The Underworld heat equation is solved through time for the initial model set-up with no
mechanical boundary conditions set or physical movement through time. The temperature field
then developed is used for coupled thermo-mechanical modelling over time. In this study, the air
remains constant at 20°C. The asthenosphere is initially 1300°C however this is not kept as a
constant and is able to evolve through time. The introduction of a 1300°C thermal anomaly in the
far right of the oceanic lithosphere was also imposed in order to weaken the crust and allow later
rifting. This temperature equilibration was undertaken for the initial material geometry seen in
Figure 3.6 prior to the formation of the reference model. Following the evolution of this model, the
temperature field was obtained at the point in which the reference model material geometry was
discerned. This temperature field for the reference model is seen in Figure 3.5 (b).
200 km
200 km
200 km
Figure 3.6: Development of the subducting slab to be used as
the initial material geometry for the reference model. A velocity
of 5 cm/yr was applied to the oceanic lithosphere at the right-
hand wall of the model in order to drive subduction. Red
shading indicates the presence of strain, and black arrows
show velocity.

38
3.3.4 Imposed Velocity Conditions
There is a velocity of 0 cm/yr set at each wall of the model. This means there is no inﬂow or outﬂow
of material and it is now a closed system. Therefore, the forces acting upon the system are all self-
sustaining.
3.3.5 Surface Processes and Partial Melting
An erosional threshold is set at an altitude of 5 km meaning that material cannot be displaced
above this altitude and is instead ‘eroded’ away. Furthermore, emplacement of sediments occurs
when the top of the crust becomes deeper than 5 km. This sedimentation is not reliant on the
erosional effects in other locations and no relationship exists between the two processes.
Partial melting is allowable within the crust with the maximum allowable melt shown in Table 3.2.
The resultant melt has a density 13% less than the surrounding material in both the crust and
mantle wedge. This melt is also allowable in the buoyant mantle wedge present in one of the later
models.

39
Air Sediment Continental
Crust
Lithospheric
Mantle
Asthenosp
here
Oceanic
Crust
Fault Mantle
Wedge
Density (kg
m-3)
1` 2200 2700 3395 3395 2900 3395 3300
Thermal
Expansivit
y (K-1)
0 3*10-5 3*10-5 3*10-5 3.5*10-5 3*10-5 3*10-5 3*10-5
Radiogenic
Heat
Production
(W m-3)
0 1.5*10-8 1.5*10-8 0.02*10-8 0.02*10-8 0 0 0
Viscous
Flow Law
Isoviscous Isoviscous Wet quartzite:
Paterson and
Luan, GeoSoc
London SP,
1990
Wet Olivine:
Hirth and
Kohlstedt,
Geophysical
Monograph
2003
Isoviscous Isoviscou
s
Isoviscous Wet Olivine:
Hirth and
Kohlstedt,
Geophysical
Monograph
2003
Isoviscosit
y Value
(Pa.s)
5*1018 5*1019 — — 5*1020 5*1019 5*1019 —
Stress
Exponent
— — 3.1 3.5 — — — 3.5
Activation
Energy (kJ
mol-1)
— — 135 520 — — — 520
Activation
Volume (m3
mol-1)
— — 0 23*10-6 — — — 23*10-6
Pre-
Exponentia
l Factor (Pa
s-1)
— — 1.66*10-26 1.6*10-18 — — — 1.6*10-18
Viscosity
Limiter
(Pa.s)
— — 5*1018 — 5*1023 5*1018 —
5*1023
5*1018 —
5*1023
— — 5*1018 —
5*1023
Initial
Reference
Temperatur
e (°C)
20 — — — 1300 — — —
Melt
Modiﬁer
— — Crustal Melt:
Hirth and
Kohlstedt,
Geophysical
Monograph,
2003
— — — — Mantle Melt:
Hirth and
Kohlstedt,
Geophysical
Monograph,
2003
Maximum
Melt
Fraction
(%)
— — 30 — — — — 8
Brittle Law — — Upper Crust:
Rey and Muller,
Nature 2010
Lithospheric
Mantle: Rey
and Muller,
Nature 2010
Lithospheri
c Mantle:
Rey and
Muller,
Nature
2010
— — Lithospheric
Mantle: Rey
and Muller,
Nature 2010
Cohesion
(MPa)
— — 10 10 10 — — 10

40
Cohesion
After
Softening
(MPa)
— — 2 2 2 — — 2
Friction
Coefficient
— — 0.577 0.577 0.577 — — 0.577
Friction
Coefficient
After
Softening
— — 0.1154 0.1154 0.1154 — — 0.1154
Maximum
Yield
Stress
(MPa)
— — 150 100 250 — — 250
Diffusivity
(m2 s-1)
10-6 10-6 10-6 10-6 10-6 10-6 10-6 10-6
Latent Heat
of Fusion
(kJ kg-1)
0 300 300 300 300 0 0 300
Specific
Heat (J.K
-1kg-1)
100 1000 1000 1000 1000 1000 1000 1000
Air Sediment Continental
Crust
Lithospheric
Mantle
Asthenosp
here
Oceanic
Crust
Fault Mantle
Wedge
Table 3.2: Rheological properties of the reference model

41
3.4 Model Roster
3.4.1 Inﬂuence of a Buoyant Mantle Wedge
A buoyant mantle wedge above the subducting slab was introduced to the reference model.
Dehydration of subducting slabs can lead to strong lithospheric mantle sitting above the Benioff
plane to become less viscous and more buoyant (Billeni & Gurnis, 2001). Modelling this allows for
an expansion on the work of Rey and Mueller (2010) who explored the role of a buoyant mantle
wedge on back-arc extension. However, whilst their models did not include the subducting slab
material in order to evaluate volume forces alone, this study will investigate the impact when the
subducting slab material is present. The wedge has allowable partial melt with its constraints and
rheological properties visible in Table 3.2. The remaining rheological, thermal and velocity
properties of the model are the same as the reference model.
3.4.2 Anorogenic Continental Crust
Introducing an anorogenic continental crust enables us to see what role the Cordillera itself plays in
back-arc extension. The thickness of the continental crust was changed so that it is kept constant
all the way to the continental margin with all other parameters the same as the reference model.
3.4.3 Subduction of an Oceanic Lithosphere with Zero Horizontal Velocity
The reference model decouples from the right hand wall of the model following break-up of the
oceanic lithosphere. This allows the oceanic lithosphere to undergo substantial horizontal
acceleration due to slab pull forces. In order to create a model whereby the slab is unable to break
off and undergo this acceleration, it was necessary to remove the sources of damage at the right
hand side of the model in the oceanic lithosphere. This includes both the imposed random damage
and the weakness produced by high temperature anomalies. The reference model was re-run with
this random damage removed and the addition of a constant thermal anomaly of normal oceanic
crustal temperature emplaced over the previous thermal anomaly.

42
4 Results
4.1 Palinspatic Reconstruction and Extensional Velocities in Eastern
Gondwana
Reconstructed values for the pre-deformation continental crustal width and the extensional velocity
during deformation were found for both the Lord Howe Rise and Eastern Australia margin along
each of the small circle traces in Figure 4.1. This reconstruction assumed a pre-deformation crustal
thickness of 60 km, equal to the thickness of the orogen in the Underworld models. Although the
width and velocity was calculated for each trace seen in Figure 3.4, a representative sample is
shown in Table 4.1 with their locations in Figure 4.1. These traces cover the entire region and all
the main features. Each trace in Eastern Australia is matched to an equivalent small circle trace in
the Lord Howe Rise.
Due to the fact that the Underworld models did not completely rift, the extensional velocity results
following rifting are included in the appendices section along with the results for every one of the
traces.
Trace Width (km) Cross-
Sectional
Area (km2)
Beginning of
Deformation
(Ma)
Time of
Break-up
(Ma)
Width of
Reconstructed
Crust (km)
Reconstructed
Extensional
Velocity (cm/
yr)
EA-1 237 5384 110 83 54.7 0.28
EA-2 120 2730 110 83 45.5 0.24
EA-3 172 172 105 70.4 103.2 0.20
EA-4 216 5792 105 66.4 96.5 0.31
EA-5 131 3354 105 61.2 55.9 0.17
LHR-1 851 15321 110 83 255.4 2.21
LHR-2 1123 18906 110 83 315.1 3
LHR-3 1578 23521 105 70.4 392 3.42
LHR-4 1232 19156 105 66.4 319.3 2.37
LHR-5 1024 13157 105 61.2 219.3 1.83
Table 4.1: Reconstructed properties and ensuing extensional velocities of a number of chosen
traces over the Eastern Australian margin and Lord Howe Rise.

43
EA-1
EA-2
EA-3
EA-4
EA-5
LHR-1
LHR-2
LHR-3
LHR-4
LHR-5
Figure 4.1: The locations of the small circle traces from Table 4.1 along with the COB’s and UCCB over the
crustal thickness maps from which the data was retrieved. Traces with same number in their name lie on the
same small circle.

44
4.2 Numerical Models
For each model a variety of factors were noted in order to properly evaluate the results. In
particular, the timing and pattern of slab movement and continental crust extension was
determined. At the continental crust, passive tracers were placed at the continent - ocean boundary
and also 50 km to the right of the landward side of the orogen. This allowed for measures of crustal
movement and average extensional velocities to be determined by calculating the distance moved
divided by time. Furthermore, the stretching factor for the continental crust for each of the models
was also calculated. This stretching factor is equal to the stretched width of the crust divided by the
initial width of the crust. The width of the orogen (400 km) was used as the initial crustal width.
4.2.1 Evolution of the Reference Model
During the first time steps of the reference model evolution, the model is searching for its isostatic
equilibrium. This is noticeable by the circular velocity arrows visible in Figure 4.2. At 0 Myr the
model seeks to lift the continental crust and lower the oceanic lithosphere near the subduction
300 km 300 km
300 km
Figure 4.2: Evolution of the reference model in it’s entire domain until 6 Myr. The initial timestep shows the model
searching for its isostatic equilibrium. The continental crust experiences extension up until 3 Myr, at which time the
oceanic slab is close to complete detachment from the right hand side of the wall. The slab continues to accelerate
downwards as it detaches from the wall before reaching a depth of 670 km by 6 Myr. The temperature profile at 6
Myr is also displayed with the scale in Kelvin. The temperature profile shows that the slab is not just different from
the surrounding asthenosphere due to material properties but that it is also thermally distinct.

45
zone. These isostatic adjustments slowly decrease through time until 1.2 Myr at which point the
model is in equilibrium.
Movement of the oceanic slab occurs very slowly during the first 2 Myr before subsequently
undergoing significant deformation in the oceanic lithosphere in the far right hand side of the
model. As this deformation increases so too does the downward velocity of the slab. The slab’s
downward velocity continues to increase until it eventually detaches from the wall at 3.4 Myr. The
slab continues to move deeper into the mantle no longer coupled to the wall until 6 Myr where it is
670 km deep as seen in Figure 4.2. As the slab begins to subduct, sediment begins to appear at
the continental - oceanic interface forming an accretionary wedge underlain by oceanic crust.
Extension of the overriding continental crust occurs instantaneously, concentrating initially in the
centre of the orogen as seen in Figure 4.3. Thinning of the lithosphere and subsequent mantle
upwelling begins to occur at the regions of high strain. Due to this upwelling, partial melting begins
to appear in the continental crust as the temperature exceeds the solidus. By 1.5 Myr extension
100 km
Figure 4.3: Evolution of the
reference model zoomed in
on a 800 km by 200 km
segment concentrating on
the continental crust. Active
strain in the continental
crust is shown in red and
partial melting is shown in
purple. The white dot acts
a passive tracer for a
particle at the continent-
ocean boundary and the
red dot is a passive tracer
initially 50 km from the
landward edge of the
orogen. Extension occurs
rapidly over the first 1.5
Myr of the model before
slowing down to a
complete stop by 2.5 Myr.
Over time the localisation
of strain and deformation
migrates throughout the
crust slowly moving
further landward. The
pattern of strain and the
extensional nature of the
system indicates
presence of normal
faulting with horsts and
grabens identifiable in the
topography.
100 km
100 km
100 km
26

46
occurs at multiple spots along the continental crust with this strain having migrated further
landward than it’s original location. As these regions of high strain experience crustal thinning,
partial melting begins to migrate to them. By 2.5 Myr, extension in the continental crust ceases as
the oceanic slab begins to accelerate and produce basal shear stresses on the continental crust.
The continental crust undergoes 141 km of extension giving it a stretching factor of 1.35. The
extensional velocity at the continent-ocean boundary is 5.64 cm/yr and 0.61 cm/yr at the landward
tracer.
4.2.2 Inﬂuence of a Buoyant Mantle Wedge
The buoyant mantle wedge was added to the mantle corner above the subducting slab as seen in
Figure 4.4. Strain initially localises in the centre of the orogen however this migrates into two sites
of deformation by 0.5 Myr. Migration continues to three sites of localised deformation by 1 Myr. By
1.3 Myr strain localised on the landward side of the orogen. This remained the site of preferential
extension until 2.5 Myr at which point extension ceases in the continental crust due to the basal
shear stresses from the oceanic slab as it begins to accelerate. This slab movement is similar to
100 km
Figure 4.4: Evolution of the
buoyant mantle wedge model
zoomed in on a 800 km by 200km
segment concentrating on the
continental crust. The mantle
wedge is in yellow. Like Figure 4.3
the model shows rapid extension
over the ﬁrst 1.5 Myr before
slowing to a complete stop by 2.5
Myr. However, this extension
continues 80 km further than the
reference model. Deformation
occurs in a similar pattern to that
seen in the reference model with
horst and grabens visible.
Increased partial melting also
occurs in the continental crust
above the mantle wedge.
Furthermore, partial melting
occurs in the mantle wedge as it
rises and undergoes
decompression melting. Oceanic
slab movement is extremely
similar to that of the reference
model.
100 km
100 km
23
100 km

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