In an effort to improve our understanding of the seismic
character of the crust beneath southeast Australia and how it relates to
the tectonic evolution of the region, we analyse teleseismic earthquakes
recorded by 24 temporary and 8 permanent broadband stations using the
receiver function method. Due to the proximity of the temporary stations to
Bass Strait, only 13 of these stations yielded usable receiver functions,
whereas seven permanent stations produced receiver functions for subsequent
analysis. Crustal thickness, bulk seismic velocity properties, and internal
crustal structure of the southern Tasmanides – an assemblage of Palaeozoic
accretionary orogens that occupy eastern Australia – are constrained by
H–κ stacking and receiver function inversion, which point to the following:
a ∼ 39.0 km thick crust; an intermediate–high Vp/Vs ratio
(∼ 1.70–1.76), relative to ak135; and a broad (> 10 km) crust–mantle transition beneath the Lachlan Fold Belt. These results are
interpreted to represent magmatic underplating of mafic materials at the
base of the crust.
a complex crustal structure beneath VanDieland, a
putative Precambrian continental fragment embedded in the southernmost
Tasmanides, that features strong variability in the crustal thickness (23–37 km) and Vp/Vs ratio (1.65–193), the latter of which likely represents
compositional variability and the presence of melt. The complex origins of
VanDieland, which comprises multiple continental ribbons, coupled with
recent failed rifting and intraplate volcanism, likely contributes to these
observations.
stations located in the East Tasmania Terrane and
eastern Bass Strait (ETT + EB) collectively indicate a crust of uniform
thickness (31–32 km), which clearly distinguishes it from VanDieland to the
west.
Moho depths are also compared with the continent-wide AusMoho model in
southeast Australia and are shown to be largely consistent, except in
regions where AusMoho has few constraints (e.g. Flinders Island). A joint
interpretation of the new results with ambient noise, teleseismic tomography,
and teleseismic shear wave splitting anisotropy helps provide new insight
into the way that the crust has been shaped by recent events, including
failed rifting during the break-up of Australia and Antarctica and recent
intraplate volcanism.
Introduction
The Phanerozoic Tasmanides (Collins and Vernon, 1994; Coney, 1995; Coney et
al., 1990) comprise the eastern third of the Australian continent and,
through a process of subduction accretion, were juxtaposed against the
eastern flank of the Precambrian shield region of Australia beginning in the
late Neoproterozoic and early Palaeozoic (Foster and Gray, 2000; Glen, 2005;
Glen et al., 2009; Moresi et al., 2014) (Fig. 1). Persistent sources of
debate that impede a more complete understanding of the geology of the
Tasmanides include (1) the geological link between Tasmania – an island
state in southeast Australia – and mainland Australia, which are separated
by the waters of Bass Strait; and (2) the presence and locations of
continental fragments from Rodinian remnants that are entrained within the
accretionary orogens. Furthermore, the lateral boundaries between individual
tectonic blocks and their crustal structure are often not well defined. To
date, few constraints on crustal thickness and seismic velocity structure
have been available for regions such as Bass Strait. Therefore, constraints on the Moho
transition, crustal thickness, and velocity structure beneath Bass Strait
derived from receiver functions (RFs) can provide fresh insight
into the nature and evolution of the Tasmanides.
Regional map of southeastern Australia that shows key
geological boundaries and the location of observed or inferred tectonic
units (modified from Bello et al., 2019a). Thick black lines delineate
structural boundaries, and the thick brown dashed line traces out the
boundary of VanDieland. The following abbreviations are used in the figure: HF – Heathcote Fault; GF – Governor Fault; BF – Bootheragandra Fault; KF – Koonenberry Fault; THZ – Torrens Hinge Zone; MA – Macquarie Arc; NVP – Newer Volcanics Province; KI – King Island in Bass Strait; FI – Flinders Island in Bass Strait; WTT – West Tasmania Terrane; ETT – East Tasmania Terrane;
AL – Arthur Lineament; TFS – Tamar Fracture System; and
RCB – Rocky Cape Block. Outcrop boundaries are sourced
from Rawlinson et al. (2016).
Previous estimates of crustal thickness and structure beneath southeastern
Australia have been obtained from deep seismic reflection transects,
wide-angle seismic data, topography, and gravity anomalies (e.g. Collins,
1991; Collins et al., 2003; Drummond et al., 2006; Kennett et al., 2011).
Earlier RF studies in southeast Australia (Shibutani et al., 1996; Clitheroe
et al., 2000; Tkalčić et al., 2011; Fontaine et al., 2013a, b)
suggested the presence of complex lateral velocity variations in the
mid-lower crust that probably reflect the interaction of igneous
underplating, associated thinning of the lithosphere, recent hotspot
volcanism, and uplift. Furthermore, the intermediate to high crustal Vp/Vs ratio
of 1.70–1.78 in this region (Fontaine et al., 2013a), relative to ak135
continental crust where Vp/Vs is ∼1.68, may indicate a mafic composition
that includes mafic granulite rocks, granite gneiss, and biotite gneiss.
Body- and surface-wave tomography (Fishwick and Rawlinson, 2012; Rawlinson
et al., 2015) revealed P- and S-wave velocity anomalies in the uppermost mantle
beneath Bass Strait and the Lachlan Fold Belt. Ambient noise surface wave
tomography (Bodin et al., 2012b; Young et al., 2012; Pilia et al., 2015b,
2016; Crowder et al., 2019) of the southern Tasmanides revealed significant
crustal complexity, but it is unable to constrain crustal thickness or the
nature of the Moho transition.
The goal of this paper is to provide fresh insight into the crust and Moho
structure beneath the southern Tasmanides using P-wave RFs and to explain the
origin of the lateral heterogeneities that are observed. This will allow us
to explore the geological relationship between the different tectonic units
that constitute the southern Tasmanides and to develop an improved
understanding of the region's tectonic history.
Geological setting
The Palaeozoic–Mesozoic Tasmanides of eastern Australia form part of one of
the most extensive accretionary orogens in existence and evolved from
interaction between the East Gondwana margin and the proto-Pacific Ocean.
The tectonic evolution of the Tasmanides is complex, and large-scale
reconstructions have proven difficult. This is evident from the variety of
models that have been suggested to explain how the region formed (Foster and
Gray, 2000; Spaggiari et al., 2003; Teasdale et al., 2003; Spaggiari et al.,
2004; Boger and Miller, 2004; Glen, 2005; Cawood, 2005; Glen et al., 2009;
Cayley, 2011a, b; Gibson et al., 2011; Moresi et al., 2014; Pilia et al.,
2015a, b). Particular challenges arise from multiple subduction events,
multiple phases of metamorphism, entrainment of exotic continental blocks,
the formation of large oroclines, recent intraplate volcanism, and subsequent
events, including the separation of Antarctica and Australia and the
formation of the Tasman Sea. These challenges are compounded by the presence
of widespread sedimentary sequences that hinder direct access to basement
rocks (Fig. 1).
The Tasmanides consist of four orogenic belts, namely the Delamerian,
Lachlan, Thomson, and New England orogens. The Delamerian Orogen – located in
the south – is the oldest part of the Tasmanides and has a southward
extension across Bass Strait from Victoria into western Tasmania, where it
is commonly referred to as the Tyennan Orogen (Berry et al., 2008). Between
about 514 and 490 Ma, the Precambrian and early Cambrian rocks that
constitute the Delamerian Orogen were subjected to a contractional orogenic
event along the margin of East Gondwana (Foden et al., 2006). Subsequently,
the Lachlan Orogen formed in the east, which contains rocks that vary in age
from Ordovician to Carboniferous (Glen, 2005). Gray and Foster (2004) argued
for a tectonic model of the Lachlan Orogen that involved the interaction of a
volcanic arc, oceanic microplates, and three distinct subduction zones. Each
subduction zone is linked to the formation of a distinct tectonic terrain:
the Stawell–Bendigo zone, the Tabbarebbera zone, and the Narooma accretionary
complex. The limited rock exposure in the Tasmanides as a whole has made
direct observation of the Lachlan Orogen difficult; this is attributed to a
large swath of Mesozoic–Cenozoic sedimentary cover and more recent
Quaternary volcanics, which obscure a large portion of the underlying
Palaeozoic terrane. However, the Lachlan Orogen contains belts of Cambrian
rocks in Victoria and New South Wales that are similar in age to the
Delamerian Orogen (Gray and Foster, 2004).
The presence of Precambrian outcrops in Tasmania and the relative lack of rocks that are similar in age in adjacent mainland Australia has led to different models
which attempted to explain the existence of Proterozoic Tasmania. For
instance, Li et al. (1997) suggested that western Tasmania may be the
remnant of a continental fragment set adrift by Rodianian break-up, whereas
Calvert and Walter (2000) proposed that King Island, along with western
Tasmania, rifted away from the Australian craton around ∼ 600 Ma (Fig. 1). Other researchers have developed scenarios in which the island
of Tasmania was present as a separate microcontinental block that was
positioned outboard of the eastern margin of Gondwana before reattaching at
the commencement of the Palaeozoic (Berry et al., 2008).
A popular model that attempts to reconcile the geology observed in Tasmania
and adjacent mainland Australia is that of Cayley (2011a). This model
proposes that central Victoria and western Tasmania formed a
microcontinental block called “VanDieland” that fused with East Gondwana
at the end of the Cambrian, possibly terminating the Delamerian Orogeny.
VanDieland became entangled in the subduction–accretion system which built
the Palaeozoic orogens that now comprise eastern Australia (Fig. 1).
Delineating Precambrian continental fragments within southeast Australia has
proven difficult, partly due to more recent sedimentary cover that obscures
large tracts of the Tasmanides. However, if present, they likely have
distinctive structural and seismic velocity characteristics (Glen, 2013).
Previous geophysical studies
To date, a variety of geophysical methods have been employed to study the
crustal structure of the Tasmanides. Shibutani et al. (1996) applied a
non-linear inversion method to RF waveforms to constrain the shear wave
velocity beneath broadband seismic stations in eastern Australia. They found
that the Moho is relatively shallow (30–36 km depth) and sharp within the
cratonic region, and deeper (38–44 km) and transitional along the axis of
the Tasmanides. They suggested that crustal thickening of the fold belt by
underplating or intrusion of mantle materials may have contributed to this
observation. Clitheroe et al. (2000) built on this earlier work by inverting
RFs to map broad-scale crustal thickness and the Moho character across the
Australian continent. They found that there was generally good agreement
between xenolith-derived estimates of the Moho depth and those determined by RF
inversion, except beneath the Lachlan Fold Belt, where a broad Moho
transition may be present. Overall, however, the RF results were consistent
with those determined by Drummond and Collins (1986) and Collins (1991), who
used seismic reflection and refraction transects to determine that the
Lachlan Fold Belt includes the thickest crust (∼50 km) in eastern
Australia. A more recent study by Fontaine et al. (2013a) employed
H–κ stacking and non-linear RF inversion to investigate crustal
thickness, shear wave velocity structure, as well as dipping and anisotropy
of the crustal layers. Their results also indicated a thick crust (∼ 48 km) and an intermediate (2–9 km) crust–mantle transition beneath the Lachlan
Fold Belt, which could be attributed to underplating beneath the crust
and/or high concentrations of mafic rocks in the mid-lower crust. Their
results also showed a dipping Moho and crustal anisotropy in the
vicinity of three seismic stations (YNG, CNB, and CAN).
Over the last decade, ambient noise tomography has become a popular tool for
studying the structure of the Australian crust. Saygin and Kennett (2010)
produced the first group velocity maps of the Australian continent from
Rayleigh wave group velocity dispersion in the period range from 5.0 to 12.5 s. Limited spatial resolution (∼2∘×2∘) in our
study region means that this model is only able to represent the structure
beneath Bass Strait as a broad, low-velocity anomaly. However, the group
velocities exhibit a good correlation with known basins and cratons.
Subsequent studies using denser arrays covering southeast mainland Australia
(Arroucau et al., 2010), southeastern Australia (Young et al., 2013), and
northern Tasmania (Young et al., 2011) show good correlations between
both group and phase velocity maps and sedimentary and basement terrane boundaries.
In order to account for uneven data distribution, Bodin et al. (2012b) used
a Bayesian transdimensional inversion scheme to generate group velocity maps
that span the Australian continent from multi-scale ambient noise datasets.
However, in our study area, their model is of low resolution due to the
limited station coverage; hence, few details on crustal structure can be
inferred. Bodin et al. (2012a) subsequently applied Bayesian statistics to
reconstruct the Moho geometry of Australia using a variety of seismic
datasets, which gave an approximate Moho depth of ∼ 30 km beneath Bass
Strait. Pilia et al. (2015a, b) and Crowder et al. (2019) derived 3-D shear
wave velocity models of the Bass Strait region using ambient noise data from
the same array of temporary stations that we exploit in this study. They
were able to constrain the lateral and depth extent of the primary
sedimentary basins in the region as well as providing insight into the seismic
character of the Precambrian microcontinental block that appears to
underpin southern Victoria, northwestern Tasmania, and Bass Strait.
Teleseismic tomography has also been used to image the lithosphere beneath
southeast Australia, thanks in part to the prolific deployment of
short-period seismometers as part of the WOMBAT transportable array project
(Rawlinson and Kennett, 2008; Rawlinson et al., 2015, 2016). While the main
focus has been on the upper mantle, in Tasmania, where station spacing was
denser, some constraints on crustal velocity structure were possible.
Rawlinson and Urvoy (2006) found that the crust beneath the East Tasmania Terrane (ETT) was
significantly faster than the crust beneath central Tasmania, which may
represent a contrast between crust with oceanic provenance in the east and
Precambrian continental provenance in the west. Bello et al. (2019b) built
on this work by including teleseismic arrival time data from the same
temporary deployment as the current study to generate a detailed upper-mantle model of southeast Australia, which revealed that Bass Strait was
underlain by lower velocities, consistent with thinned lithosphere as a
result of failed rifting during the break-up of Australia and Antarctica.
Active source seismic profiling has also been widely used in southeast
Australia to characterize crustal velocity structure (e.g. Finlayson et al.,
1980; Collins, 1991; Finlayson et al., 2002; Drummond et al., 2006; Glen,
2013). This has largely focused on the transition from continental to
oceanic crust at passive margins, but it has also been used to image major
transition zones or faults between orogens (Glen, 2013) or within orogens
(Cayley et al., 2011a, b), the latter of which lead to the VanDieland
microcontinental model. Rawlinson and Urvoy (2006) jointly inverted
teleseismic arrival times and active source wide-angle travel times in
northern Tasmania to constrain crustal velocity, Moho geometry, and upper-mantle velocity structure, and they found that both northeastern and northwestern
Tasmania are characterized by thinner (< 28 km) and higher-velocity
crust compared with central Tasmania.
Potential field data have also been exploited to study the formation and
structure of the Tasmanides. Gunn et al. (1997) integrated potential field
data (magnetic and gravity), seismic reflection data, outcrop geology, and
well information to study the crustal structure of the Australian continent.
Their study found that the occurrence of tensional stress, oriented northeast–southwest
along basement structures in the Bass Basin, is able to explain the
formation of the three major sedimentary basins that overlie dense mafic
material, which in turn was formed by mantle decompression processes
associated with crustal stretching. From the interpretation of new
aeromagnetic data, Morse et al. (2009) delineated the architecture of the
Bass Strait basins and their supporting basement structure. Subsequent
studies by Moore et al. (2015, 2016) used gravity, magnetic, seismic
reflection, and outcrop data to support the hypothesis of a VanDieland
microcontinent. Their study showed that VanDieland comprises seven distinct
microcontinental ribbon terranes that appear to have amalgamated by the late
Cambrian, with major faults and suture zones bonding these ribbon terranes
together.
While the last few decades have seen important advances and insights made
into our understanding of the southern Tasmanides, there is still
limited data on the deep crustal structure beneath Bass Strait, which is our
region of interest. Therefore, it is timely that we can exploit, using the RF
technique, teleseismic data recorded by a collection of temporary and
permanent seismic stations in the region to study the structure of the
crust, Moho, and uppermost mantle beneath mainland Australia, Bass Strait, and
Tasmania.
Data
A collaboration involving five organizations (the University of Tasmania,
the Australian National University, Mineral Resources Tasmania, the Geological
Survey of Victoria, and FROGTECH) deployed the temporary BASS seismic array
from May 2011 to April 2013. It consisted of 24 broadband, three-component
seismic stations that spanned northern Tasmania as well as a selection of islands
in Bass Strait and southern Victoria. The instruments used were 23 Güralp 40T and 1 Güralp 3ESP sensors coupled to Earth Data PR6-24
data loggers. The permanent stations consist of eight broadband sensors
managed by IRIS, GEOSCOPE, and the Australian National Seismic Network
(ANSN). The distribution of all 32 seismic stations used in this
study is plotted in Fig. 2. Earthquakes with magnitudes mb> 5.5 at epicentral distances between 30 and 90∘ comprise the seismic sources used in this analysis (Fig. 3). This resulted
in an acceptable azimuthal coverage of earthquakes between the northwest and
east of the array, where active convergence of the Australian and Eurasian
plate coupled with westward motion of the Pacific plate has produced
extensive subduction zones. To the south and southwest of the array, the
absence of subduction zones in the required epicentral distance range means
that there are significantly fewer events available for analysis from these
regions.
Location of seismic stations used in this study
superimposed on a topographic and bathymetric map of southeast Australia
(modified from Bello et al., 2019a). The boundary of VanDieland is
delineated by a thick black dashed line. The thick red dashed line outlines the
boundary of the East Tasmania Terrane and Furneaux Islands. The thick white
dashed line highlights the eastern sector of the Lachlan Fold Belt.
The topography and bathymetry are based on the ETOPO1 dataset (Amante and
Eakins, 2009).
MethodsReceiver functions
The RF technique (Langston, 1979) uses earthquakes at teleseismic distances
to enable estimation of the Moho depth and shear wave velocity structure in the
vicinity of a seismic recorder. If this technique can be applied to a
network of stations with good spatial coverage, it represents an effective
way of mapping lateral variations in the Moho depth and crustal structure.
A recorded teleseismic wave field at a broadband station can be described by
the convolutional model in which operators that represent the source
radiation pattern, path effects, crustal structure below the station, and
instrument response are combined to describe the recorded waveform. By using
deconvolution to remove the effects of the source, path, and response of the
instrument (e.g. Langston, 1979), information on local crustal structure
beneath the station can be extracted from P–S wave conversions at
discontinuities in seismic velocity (Owens et al., 1987; Ammon, 1991).
P-wave RFs were determined from teleseismic P-wave forms using FuncLab
software (Eagar and Fouch, 2012; Porritt and Miller, 2018), following
preprocessing using the seismic analysis code (SAC) (Goldstein et al.,
2003). RFs were computed by applying an iterative time-domain deconvolution
scheme developed by Ligorria and Ammon (1999) with a 2.5 s Gaussian filter
width. This is achieved by deconvolution of the vertical component waveform
from the radial and transverse waveforms with a central frequency of
∼ 1 Hz. This frequency was selected on account of significant
source energy detected in the ∼ 1 Hz range of teleseismic P arrivals, which are sensitive to crustal-scale anomalies. It also provides a
favourable lateral sensitivity with respect to Fresnel zone width (∼ 15 km at Moho depth) when the conversions from P to S are mapped as velocity and
crustal thickness variations.
The complete set of 1765 events (Fig. 3) and 32 stations produced 21 671
preliminary RFs. These RFs were manually inspected using the FuncLab trace
editor, and a subset of 9674 RFs were selected for further analysis using
the visual clarity of the direct arrivals as an acceptance criterion. Due to
high noise levels and fewer events associated with the temporary BASS array
dataset, a modest number of good-quality RFs resulted from the above
selection method; therefore, different selection criteria were applied that assessed
the P-arrival, Moho conversion, and later amplitudes in conjunction with
overall noise levels exhibited by the transverse component RFs. This enabled
the temporary BASS stations to yield between 2 and 30 good-quality receiver
functions, and increased the number of stations where H–κ stacking
and NA inversion could be applied from 13 to 20.
Distribution of distant earthquakes (teleseisms) used in
this study. The locations of events that are ultimately used for RF analysis
are denoted by yellow dots. Concentric circles are plotted at
30∘ intervals from the centre of Bass Strait.
Topography and bathymetry colours are based on the ETOPO1 dataset
(Amante and Eakins, 2009).
H–κ stacking
Having obtained reliable P-wave RFs, the H–κ stacking technique is used
to estimate crustal thickness and bulk Vp/Vs for individual stations. We apply
the method of Zhu and Kanamori (2000) to stations where the direct Ps (Moho
P to S conversion) phase and its multiples are observed. This technique makes
use of a grid search to determine the crustal thickness (H) and Vp/Vs (κ)
values that correspond to the peak amplitude of the stacked phases. A clear
maximum requires a contribution from both the primary phase (Ps) and the
associated multiples (PpPs). In the absence of multiples, the maximum
becomes smeared out due to the inherent trade-off between crustal thickness
(H) and average crustal velocity properties (κ) (Ammon et al., 1990;
Zhu and Kanamori, 2000). The H–κ stacking algorithm reduces the
aforementioned ambiguity by summing RF amplitudes for Ps and its multiples – PpPs and PpSs+PsPs – at arrival times corresponding to a range of H and Vp/Vs values. In the
H–κ domain, the equation for stacking amplitude is
sH,κ=∑j=1Nw1rjt1+w2rjt2+w3rjt3,
where rj(ti); i=1, 2, 3 are the RF amplitude values at the expected
arrival times t1, t2, and t3 of the Ps, PpPs, and PpSs+PsPs phases
respectively for the jth RF; w1, w2, and w3 are weights based on
the signal-to-noise ratio (w1+w2+w3=1); and N is the total
number of radial RFs for the station. s(H,κ) achieves its maximum value
when all three phases stack constructively, thereby producing estimates for
H and Vp/Vs beneath the station (see Figs. 5 and S1–S4). In
this study, the weighting factors used are w1=0.6, w2=0.3, and
w3=0.1. The H–κ approach requires an estimate of the mean
crustal P-wave velocity, which is used as an initial value. Based on the
results of a previous seismic refraction study (Drummond and Collins, 1986),
we use an average crustal velocity of Vp=6.65 km/s to obtain our estimates
of H and κ in the study area, noting that H–κ stacking results
are much more dependent on Vp/Vs than Vp (Zhu and Kanamori, 2000). To estimate the
uncertainties in the H–κ stacking results, we compute the standard
deviation of the H and κ values at each station. When only a small
number of RFs are available at a station (e.g. four in the case of MILA), the
estimates are unlikely to be particularly robust, and in such instances, they are
perhaps best viewed as a lower bounds on uncertainty.
While simple to implement, the Zhu and Kanamori (2000) method can suffer
from large uncertainties due to its assumption of a simple flat-laying layer
over a half-space with constant crustal and upper-mantle properties.
Consequently, there are only two search parameters (H and κ) plus a priori
information (Vp, weightings), and this method does not account for variation with
back-azimuth. These problems can cause non-unique and inaccurate estimates,
which can lead to potentially misleading interpretations; for instance, a
low-velocity upper-crustal layer can appear as a very shallow Moho in an
H–κ stacking search space diagram. Also, a dipping Moho and/or
anisotropic layers within the crust can contribute to uncertainty.
Non-linear waveform inversion
In an effort to refine the crustal model, we invert a stack of the radial
RFs by adopting the workflow described by Shibutani et al. (1996). We divide
the waveform data (RFs) into four 90∘ quadrants based on the
back-azimuth of their incoming energy. The first quadrant back-azimuth
range is from 0 to 90∘, and an equivalent range in a clockwise
direction defines the consecutive quadrants. The second and third
quadrants (southeastern and southwestern back-azimuths) have very small
numbers of RFs. Data from the first and fourth quadrants are of better
quality, with the first quadrant showing more coherency than the fourth
quadrant, which is likely due to the orientation of surrounding tectonic
plate boundaries; hence, the pattern of P-wave energy radiated towards
Australia. Kennett and Furumura (2008) showed that seismic waves arriving in
Australia from the northern azimuths undergo multiple scattering but low
intrinsic attenuation due to heterogeneity in the lower crust and mantle;
this tends to produce prolonged high-frequency coda. An important assumption
in our inversion is that we neglect anisotropy and possible Moho dip, which
we assume have a second-order influence on the waveforms that we use to constrain
1-D models of the crust and upper mantle.
Visual examination of coherency in P to S conversions allows us to select a
subset of RF waveforms for subsequent stacking. This resulted in groups of
mutually coherent waveforms after which a move-out correction is then applied
to remove the kinematic effect of different earthquake distances prior to
stacking using a cross-correlation matrix approach described in Chen et al. (2010) and Tkalčić et al. (2011). Our visual acceptance criteria
yields RFs at only 14 out of the 32 stations used for this study. An example
of some stacked RFs is given in Fig. 4.
Stacked receiver functions from the Australian National
Seismic Network (ANSN) stations TOO, YNG, and MOO as well as the Global Seismographic Network (GSN) station TAU. Small
arrows indicate the arrival of the Ps (black),
PpPs (red), and PpPs+PsPs (blue) phases from the Moho.
Results from the H–κ
stacking analysis for RFs (Zhu and Kanamori, 2000) at MOO, CAN, and
TOO stations. (a) Normalized amplitudes of the stack over all
back-azimuths along the travel time curves corresponding to the
Ps and PpPs phases for each case. (b) The corresponding stacked receiver function for each station.
We invert RFs for 1-D seismic velocity structure beneath selected seismic
stations using the neighbourhood algorithm (NA) (Sambridge, 1999a, b) in
order to better understand the internal structure of the crust and the
nature of the transition to the upper mantle. NA makes use of Voronoi cells
to help construct a searchable parameter space, with the aim of
preferentially sampling regions of low data misfit. In the inversion
process, a Thomson–Haskell matrix method (Thomson, 1950; Haskell, 1953)
was used to calculate a synthetic radial RF for a given 1-D (layered)
structure. During the inversion, as in Shibutani et al. (1996) and Clitheroe
et al. (2000), each model is described by six layers: a layer of sediment; a
basement layer; an upper crust, middle crust and lower crust; and an
underlying mantle layer, all of which feature velocity gradients and,
potentially, velocity jumps across boundaries. The inversion involves
constraining 24 parameters: Vs values at the top and bottom of each layer,
layer thickness, and the Vp/Vs ratio in each layer (Table 1). The inclusion of
Vp/Vs ratio as an unknown primarily aims to accommodate the effects of a sediment
layer with limited prior constraints (Bannister et al., 2003). There are two
important controlling parameters required by NA: (1) the number of models
produced per iteration (ns); and (2) the number of neighbourhoods
re-sampled per iteration (nr). After a number of trials, we chose the
maximum number of iterations to be 5500, with ns=13 and nr=13
for all iterations. We employ a chi-square χ2 metric (see
Sambridge, 1999a, for more details) to compute the misfit function, which is a
measure of the inconsistency between the true (ϕiobs) and
predicted (ϕiprem) waveforms for a given model
m:
χν2m=1ν∑i=1Ndϕiobs-ϕipremσi,
where σi represents the noise standard deviation determined
from ϕiobs, following the method described in Gouveia and
Scales (1998); and ν represents the number of degrees of freedom (the
difference between the number of observations and the number of parameters
being inverted for). Using the above-mentioned parameters, the inversion
targets the 1-D structure that produces the best fit between the predicted
and observed RF. Figures 7–9 and S5–S9 present example
results of inversions via density plots of the best 1000 data-fitting
S-wave velocity models produced by the NA. The optimum data-fitting model is
plotted in red.
Model parameter bounds used in the neighbourhood algorithm
receiver function inversion. Vsupper and Vslower represent
the S- wave velocity at the top and bottom of a layer respectively.
Vp/Vs represents the P- and S-wave velocity ratio within a layer.
Maps of crustal thicknesses and average Vp/Vs from H–κ stacking in
southeast Australia from 16 stations are shown in Fig. 6. At the remaining
stations, we could not detect any clear multiples or Moho conversions in the
RFs from any direction. A previous study by Chevrot and van der Hilst (2000)
has noted that this region is devoid of clear multiples. The crustal
thickness for all analysed stations in the study area varies from 23.2 ± 5.0 km (BA02) beneath northwestern Tasmania to 39.1 ± 0.5 km (CAN)
beneath the Lachlan Fold Belt, and the variation is strongly correlated with
topography. Crust beneath VanDieland (Fig. 6a) is thin in the north
(∼ 37.5 km) and south (∼ 33 km), but it appears to
be considerably thinner beneath the Victorian and Tasmanian margin Bass
Strait (∼ 25 km). The mountainous region of the Lachlan Fold
Belt has the deepest Moho at 39.1 ± 0.5 km (CAN) and a corresponding
Vp/Vs value of 1.73 ± 0.02. Crust that is consistently between ∼ 31
and 33 km thick lies beneath the East Tasmania Terrane and eastern Bass
Strait (ETT + EB). The Vp/Vs ratio varies between ∼ 1.65 beneath
station BA11, which also exhibits the thinnest crust, and ∼ 1.93 beneath stations BA19 and BA20 in southern Victoria. There is no
obvious correlation between the number of RFs used in the H–κ
stacking and the size of the uncertainty in either the Moho depth or Vp/Vs, but as
mentioned previously, the uncertainty estimates for stations with a low
number of RFs are likely to be less robust. Table 2 shows a summary of the
H–κ stacking results for the stations that have been analysed.
(a) Variations in the crustal thickness and (b)Vp/Vs ratio taken from the linear
(H–κ) stacking results (Table 2). Crustal
thickness varies between ∼ 23 and 39 km.
The Vp/Vs ratios vary from ∼ 1.65 to
1.93. The thick black dashed line denotes the boundary of VanDieland. The thick
magenta dashed line outlines the boundary of the East Tasmania Terrane and
eastern Bass Strait (ETT + EB). The thick cyan dashed line highlights
the eastern part of the Lachlan Fold Belt. The illuminated
topography and bathymetry are based on the ETOPO1 dataset (Amante and
Eakins, 2009).
Summary of H–κ stacking and NA inversion results
for the current study.
Results of the NA inversion were successfully obtained for a selection of
permanent and temporary stations, as shown in Table 2 and Fig. 10. If the
Moho is defined by a gentle velocity gradient, the base of the velocity
gradient is used as a proxy for the Moho depth, as done in previous RF (e.g.
Clitheroe et al., 2000; Fontaine et al., 2013a) and seismic refraction
(Collins, 1991; Collins et al., 2003) studies. We also adopt an upper-mantle
velocity of Vp=7.6 km/s (i.e. Vs=4.3–4.4 km/s for Vp/Vs ratios of 1.73–1.77 at the base of the Moho gradient) following Clitheroe et al. (2000), who used
this value for RF studies, and Collins et al. (2003), who used Vp> 7.8 km/s for their summary of both seismic refraction and RF results; these
Vp values are consistent with global Earth models (e.g. Kennett et al., 1995).
Therefore, we also require the S-wave velocity to be >∼ 4.4 km/s beneath the Moho. We present the S-wave velocity
profiles from the NA inversion for stations CAN, MOO, TOO, YNG, BA13, and
BA17 in Figs. 7–9 as well as the observed and predicted RFs. The S-wave
velocity inversion results for the remaining stations are included in the Supplement (Figs. S5–S8). In assigning the
Moho depth, we consider three criteria to examine the quality of the
inversion result: (1) misfit value χ2; (2) the quality of the RF
stack (which is based on our ability to pick the direct and multiple
phases); and (3) the visual fit between the synthetic and observed RF.
Models that fail to fit significant arrivals in the observed RF are
rejected. Based on these criteria, the inversion results are classified as follows:
very good – very low χ2 (typically < 0.4), very good
visual fit to direct and multiple phases;
good – low χ2 (typically 0.4–0.8), direct phases clearly visible,
multiple phases less clear, and a good visual fit to all major identifiable
phases;
poor – medium to high χ2 (in the range 0.8–1.2), direct phases
visible, multiple phases unclear, and moderate visual fit to some
identifiable phases.
(a, c) Seismic velocity models for the CAN and MOO stations
obtained from the neighbourhood algorithm (Sambridge, 1999a). The grey area
indicates all the models searched by the algorithm. The best 1000 models are
indicated by the yellow to green colours; the best model (smallest misfit)
is shown using the red line, for both the S-wave
velocity (top x axis) and Vp/Vs ratio
(bottom x axis), and the white line is the average velocity model.
Small black arrows denote the estimated depth of the Moho. (b, d) Waveform
matches between the observed stacked receiver functions (black) and
predictions (blue) based on the best models. “Misfit” refers to the
chi-square estimate as defined by Eq. (2).
In general, the optimum χ2 value is normally considered to be 1,
as below this value, the tendency is to fit noise rather than signal.
However, this is for the ideal case when the number of degrees of freedom
and the absolute values of the data uncertainty are well known (e.g. in the
case of a synthetic test). In the case of observational data, these values
are often poorly constrained, so using the relative χ2 values coupled with visual assessment of the data fit appears to be
reasonable. With regard to the character of the crust–mantle transition,
this study classifies the transition zone as sharp ≤ 2 km, intermediate
2–10 km, or broad ≥ 10 km, as initially proposed by Shibutani et al. (1996) and modified by Clitheroe et al. (2000).
We note that for the seven permanent stations for which we produce receiver
function inversion and H–κ stacking results, five have estimates of the Moho
depth from previous receiver function studies. Clitheroe et al. (2000)
estimated the Moho depth at 49 km beneath CAN based on a non-linear inversion,
which is ∼ 10 km greater than the results we obtain for both
NA inversion and H–κ stacking (see Sect. 7.1 for further
discussion of this discrepancy). Ford et al. (2010) determined the Moho depth
beneath the MOO, TOO, TAU, and YNG stations using H–κ stacking and found
values (compared with our H–κ stacking results) of 33 ± 3 km
(33.0 ± 1.2 km), 34 ± 3 km (37.5 ± 1.2 km), 32 ± 3 km
(33.5 ± 1.9 km), and 33 ± 2 km (37.3 ± 0.5 km) respectively.
These are all within error, with the slight exception of YNG station,
located in Young, on the western flanks of the Great Dividing Range, where
we might expect the crust to be slightly thicker than average. Overall,
however, these similarities suggest that our results are likely to be
robust.
(a, c) Seismic velocity models for TOO and YNG stations
obtained from the neighbourhood algorithm. (b, d) Comparison between the
observed stacked and the predicted receiver functions from the NA inversion.
See the caption of Fig. 7 for more details.
Seismic velocity models for temporary stations BA13 (a) and BA17 (b) obtained from the neighbourhood algorithm. See Figs. S6–S9 in the
Supplement for all receiver function inversion results for
the temporary BASS network, including waveform fits (Supplement Fig. S7
includes the waveform fit for stations BA13 and BA17).
Discussion
For convenience, the seismic stations are separated into three groups (Fig. 2) based on tectonic setting and the results obtained. Stations YNG, CAN,
CNB, MILA, and BA13 are located in the Lachlan Fold Belt; stations BA02,
BA11, BA19, BA20, TAU, MOO, and TOO sit above the VanDieland microcontinental
block; and stations BA07, BA08, BA09, and BA17 lie in the East Tasmania
Terrane and eastern Bass Strait (ETT + EB). Stations BA22 and BA24 lie to
the west of VanDieland. This discussion focuses on crustal thickness, the
nature of the Moho, and crustal velocity and velocity ratio variations from
H–κ stacking and the 1-D S-wave velocity models. Overall, the agreement
between Moho depths obtained from the H–κ stacking results and
NA inversion is generally within error (Table 2), which makes a joint
interpretation more straightforward. Comparison is also made to other
studies that have examined crustal seismic properties in southeast
Australia, and we attempt to integrate our new findings with previous
results from teleseismic tomography, SKS splitting, and ambient noise
tomography in order to better understand the crust and upper-mantle
structure and dynamics beneath this region.
Lateral variation in crustal thickness and the nature of the Moho
The RF analysis clearly reveals the presence of lateral changes in crustal
thickness that span mainland Australia through Bass Strait to Tasmania
(Figs. 6 and 10; in the latter case, RF depths from previous studies are
also included for reference). The stations located in the Palaeozoic Lachlan
Fold Belt reveal a generally thick crust that ranges between ∼ 37 and 40 km. Although the Moho was picked as a velocity jump for stations
YNG, CAN, and CNB, the velocity nonetheless tends to continue to increase
with depth below the discontinuity. This, coupled with the fact that
Clitheroe et al. (2000) estimate the Moho to be almost 10 km deeper beneath
CAN, is consistent with the presence of mafic underplating (e.g. Drummond
and Collins, 1986; Shibutani et al., 1996; Clitheroe et al., 2000), sourced
from the ambient convecting mantle. The top and bottom of such a layer could
feature a velocity step with depth, and its internal structure is likely to
be layered and/or gradational, resulting in uncertainty in the true
Moho depth. Based on deep crustal reflection profiling, Glen et al. (2002)
suggested that the deep Moho underlying the Lachlan Orogen results from
magmatic underplating that added a thick Ordovician mafic layer at the base
of the crust coupled with a thick sequence of Ordovician mafic rocks that
can be found in the mid and lower crust. Finlayson et al. (2002) and Glen et
al. (2002) also inferred the presence of underplating near CNB and CAN from
seismic refraction data. Collins (2002) postulated that the underplating
might have occurred in the back-arc region of a subduction zone due to
pronounced adiabatic decompression melting in the asthenosphere. The seismic
tomography model of Rawlinson et al. (2010, 2011) exhibits an increase in
P-wave speed at 50 km depth beneath CAN, CNB, and YNG, and the authors suggest
that magmatic underplating may be the cause of the high-velocity anomaly. A
recent study by Davies et al. (2015) identified the longest continental
hotspot track in the world (over 2000 km total length), which began in north
Queensland at ∼ 33 Ma, and propagated southward underneath the
present-day Lachlan Fold Belt and Bass Strait. The magmatic underplating
could therefore be a consequence of the passage of the continent above a
mantle upwelling leading to a more diffuse crust–mantle transition zone. The
thickened crust and a transitional Moho observed in the Lachlan Fold Belt
are consistent with the proposed delamination models of Collins and Vernon (1994).
Map showing crustal thickness variations based on the
S-wave velocity inversion results of this study (circles) and previous
studies (triangles) (Clitheroe et al., 2000; Fontaine et al., 2013a, b;
Shibutani, 1996; Tkalcic et al., 2012). The topography and bathymetry are
based on the ETOPO1 dataset (Amante and Eakins, 2009).
Strong lateral changes in crustal seismic structure (Figs. 6, 10)
beneath VanDieland appear to be a reflection of the region's complex
tectonic history. The thick crust (∼ 37 km) beneath the Selwyn
Block (see Fig. 1 for its location) – within the northern margin of
VanDieland in southern Victoria – thins dramatically to ∼ 26 km as it enters Bass Strait, increases to ∼ 30 km beneath King
Island (BA11), and then thins to ∼ 23 km beneath northwestern Tasmania,
before increasing to ∼ 33 km in southern Tasmania. The results
in southern Tasmania agree with those of Korsch et al. (2002) from a seismic
reflection profile adjacent to the TAU and MOO seismic stations. The thinner
crust beneath Bass Strait and its margins may be a consequence of
lithospheric thinning and/or delamination associated with failed rifting
that accompanied the break-up of Australia and Antarctica (Gaina et al.,
1998). Stations BA07, BA08, BA09, and BA17 (ETT + EB) collectively indicate
crust of relative uniform thickness (∼ 31–32 km; Fig. 10a, b).
Relative to western Bass Strait, the crust is slightly thicker in this part
of the study area, which may suggest underplating associated with a
Palaeozoic subduction system (e.g. Drummond and Collins, 1986; Gray and
Foster, 2004).
In general, our understanding of crustal thickness variations are limited by
station separation, so it is difficult to determine whether smooth
variations in thickness or step-like transitions explain the observations.
Vp/Vs and bulk crustal composition
Vp/Vs can constrain chemical composition and mineralogy more robustly than P- or
S-wave velocity in isolation (Christensen and Fountain, 1975). We observe
variations in Vp/Vs across the study region, which we can largely equate with
variations in composition or melt. Studies in mineral physics and field
observations show (1) an increase in Vp/Vs with decreasing SiO2 content in
the continental crust (Christensen, 1996), and (2) partial melt is revealed
by elevated Vp/Vs, especially if the anomaly is localized to an intra-crustal
layer (Owens and Zandt, 1997). A more felsic (SiO2) composition in the
lower crust is represented by a lower Vp/Vs, which reflects removal of an
intermediate-mafic zone by delamination, whereas a more mafic lower crust is
revealed by higher Vp/Vs (> 1.75) which may be due to underplated
material (Pan and Niu, 2011). However, lower crustal delamination can also
result in decompression melting, which can yield elevated Vp/Vs (He et al.,
2015). We interpret the variation of observed Vp/Vs in the southern Tasmanides to
be a consequence of compositionally heterogeneous crust and localized
partial melt that may likely be sourced from recent intraplate volcanism
(Rawlinson et al., 2017).
Figure 6b shows the distribution of bulk Vp/Vs across the study area. The pattern
of Vp/Vs ratios appears to delineate three distinct zones of crust. Beneath the
Lachlan Orogen, values are ∼ 1.75, which is consistent with
the presence of a mafic lower crust, as suggested by a number of other
studies (Drummond and Collins, 1986; Shibutani et al., 1996; Clitheroe et
al., 2000; Finlayson et al., 2002). Beneath eastern Bass Strait, the Vp/Vs ratios
are slightly lower, with BA07, BA08, and BA09 exhibiting values of 1.70, 1.70,
and 1.71 respectively. These values are in agreement with constraints from
seismic reflection and refraction studies (Finlayson et al., 2002; Collins
et al., 2003) and may indicate a felsic to intermediate crustal composition.
The geology of Flinders Island, which hosts both BA07 and BA08, is dominated
by Devonian granites, which is consistent with this observation. Beneath
VanDieland, Vp/Vs is highly variable, with the greatest contrast between BA11
(∼ 1.65) and BA19/20 (∼ 1.93), and BA19/20 and
TOO (1.68). BA11 is located on King Island, which is characterized by
Precambrian and Devonian granite outcrops, which may help explain the low
Vp/Vs. The high Vp/Vs beneath BA19/20 is harder to explain but could be caused by
melt in the crust associated with the Newer Volcanics Province, which sits
along the Cosgrove intraplate volcanic track and last erupted only
∼ 4.6 ka (Rawlinson et al., 2017). The return to lower Vp/Vs
beneath TOO over a relatively short distance (∼ 100 km) is
also difficult to explain, but we note that this region of Victoria is
underlain by granite intrusions.
In summary, the crust beneath VanDieland exhibits the greatest lateral
heterogeneity in Vp/Vs, which likely reflects considerable variations in
composition and the presence of melt. This can partially be explained by the
tectonic history of the region, which includes failed rifting in Bass Strait
accompanied by widespread magma intrusion and granite emplacement as well as, more
recently, the passage of a plume (Rawlinson et al., 2017). Furthermore,
Moore et al. (2015) used reflection transects and potential field data to
infer that VanDieland is comprised of up to seven continental ribbon
terranes that are bounded by major faults and suture zones, which were
likely amalgamated by the end of the Proterozoic. Hence, considerable
variations in composition and, in turn, Vp/Vs ratio are to be expected.
Moho depth comparison
Prior to this study, a variety of seismic methods have been used to
constrain the Moho depth in southeast Australia, including receiver functions,
reflection profiling, and wide-angle refection and refraction experiments. In
an effort to combine the results from all of these studies into a single
synthesis, Kennett et al. (2011) developed the AusMoho model; this included
Moho depth estimates from over 11 000 km of reflection transects across the
continent, numerous refraction studies, and 150 portable and temporary
stations. Due to irregular sampling, the detail of this model is highly
variable; for example, the region beneath Bass Strait is constrained by only
five measurements, whereas the central Lachlan Fold Belt around Canberra
(see Fig. 1 for location) features relatively dense sampling at
∼ 50 km intervals or less.
AusMoho includes previous receiver function results from Shibutani et al. (1996), Clitheroe et al. (2000), Fontaine et al. (2013a), and Tkalcic et al. (2012), as well as reflection and refraction transects in Tasmania, parts of
the Lachlan Orogen, and western Victoria. Figure 11 illustrates AusMoho for
our study region, which exhibits large variations in the Moho depth (from
∼ 10 to > 50 km). These extremes are due to the
presence of oceanic crust outboard of the passive margin of the Australian
continent, and the root beneath the Southern Highlands, which represent the
southern extension of the Great Dividing Range in New South Wales.
Superimposed on Fig. 11 are Moho depths from the four previous receiver
function studies cited above as well as NA inversion and H–κ depth
estimates from this study. As expected, the correlation between the previous
RF results and AusMoho is generally good, as they were part of the
dataset used to build this model. In places where they do not match, this can
be attributed to the presence of seismic refraction or reflection lines
that were also used to constrain AusMoho.
Comparison between the AusMoho model (background colour
map) and the Moho depths determined through RF analysis in this and previous
studies. Small coloured circles denote the Moho depths determined from
H–κ stacking, whereas large coloured circles correspond to
receiver function estimates. When both H–κ- and NA-derived
depths are available at a single station, the smaller H–κ
circle is superimposed on the larger NA circle so that both depths can be
observed on the one plot. Moho depths determined from previous RF studies
are denoted by triangles.
In general, the agreement between the results from this study and AusMoho is
good, but there are exceptions. For instance, CAN, CNB, YNG, and MILA tend to
be somewhat shallower than AusMoho. However, this can be attributed to the
likely presence of mafic underplating alluded to earlier, which can
effectively yield two options for the Moho transition due to an expected
high Vp/Vs (> 1.85) in the underplate layer (e.g. Cornwell et al.,
2010). AusMoho Moho depths beneath BA07 and BA08 are considerably shallower
than our estimates, which we attribute to a lack of data coverage in this
region. Sizable discrepancies also exist beneath BA02, BA19, and BA20: in
the former case, the uncertainty in our H–κ stacking estimate is 5 km, which may be a factor here; in the latter case, we also note that there
is sparse data coverage southeast of Melbourne to constrain AusMoho, so it
would appear that our new Moho depths are more likely to be correct.
Overall, while there is good consistency between AusMoho and our new
results, any updated version of AusMoho should incorporate the Moho depth
estimates from this study.
Composite result of teleseismic tomography (mantle
velocity anomalies), ambient noise (crustal velocity anomalies), receiver
functions (Moho), and shear wave splitting (inferred mantle flow relative to
over-riding plate). Velocity slices are taken at
148∘ E. Note that the crustal model produced from
ambient noise tomography is defined in terms of Vsv, but it was converted to Vp
in the study of Bello et al. (2019b) to permit its inclusion in the starting
model for the inversion of teleseismic P-wave arrival time residuals. In
this figure, the crustal Vp anomalies are shown.
Although AusMoho did make use of results from a 3-D wide-angle reflection
and refraction survey of Tasmania (offshore shots and onshore stations), it
only used a few sample points for the final Moho model (Kennett et al.,
2011); therefore, the resolution of AusMoho is considerably less than the
Moho model produced by Rawlinson et al. (2001). Consequently, we plot our
three RF results on top of this model in the Supplement (Fig. S10). The
agreement between the Moho model and RF depths beneath MOO and TAU is good,
but RF estimates beneath BA02 are shallower than the Moho model by about 4 km. However, this is within the margin of error for the H–κ stacking
result.
Synthesis
In this final section, we present a synthesis of results for southeast
Australia that are based on (1) our new receiver function results, (2) teleseismic SKS splitting results from Bello et al. (2019a), (3) teleseismic
tomography undertaken by Bello et al. (2019b), (4) ambient noise crustal
imaging results from Young et al. (2013), and (5) AusMoho (Kennett et al.,
2011). This synthesis is encapsulated in the plot shown in Fig. 12, which
is a representative transect through the Lachlan Orogen south through Bass
Strait and into Tasmania. Moho depths are taken from AusMoho and are refined
where additional information is available from our new RF results; crustal
P-wave velocity is taken from the ambient noise results (following
conversion from S-wave velocity; see Bello et al., 2019b, for more details);
and mantle P-wave velocities are taken from Bello et al. (2019b). Arrows are
based on interpreted mantle flow patterns undertaken as part of the shear
wave splitting study. This previous study used approximately the same
temporary and broadband station network as that used in the current study
and found that fast axis orientations of
anisotropy were aligned with contemporary plate motion (north-northeast) beneath the Lachlan Orogen, but beneath
Bass Strait, a radial pattern was observed that is consistent with an
upwelling mantle that impinges on the lithosphere and spreads out in all
directions. Interestingly, the location of this phenomenon corresponds
approximately to the predicted location of the Cosgrove hotspot track source
(Davies et al., 2015) and may be caused by an upwelling mantle plume. Thus,
the low velocities in the upper mantle beneath Bass Strait may be due to
elevated temperatures and melt, although it is not straightforward to
explain the higher velocities below 200 km depth in this context.
The thicker Moho boundary beneath the Lachlan Orogen (Fig. 12) reflects
the likely presence of underplating, which makes the base of the crust
harder to discern seismically. However, the crust is clearly thicker here
than beneath Bass Strait or Tasmania. The Moho depth beneath the northern part
of the Fig. 12 is not constrained by our RF results, but according to
AusMoho, it is relatively flat, which is consistent with Precambrian crust,
and there is a faster mantle lithosphere. The strong variations in crustal
velocity beneath Bass Strait can be attributed to failed rifting resulting
in the formation of thick (> 10 km) sedimentary basins and
elevated temperatures (lower velocities) as well as the intrusion of mafic-rich
materials into the lower and mid crust (higher velocities).
Conclusions
We used H–κ stacking of teleseismic RFs to determine the crustal thickness
and Vp/Vs ratio and generate 1-D S-wave velocity profiles of the crust from RF
inversion in order to investigate the internal crustal velocity structure
beneath the southern Tasmanides in southeast Australia. Our main findings
are summarized below.
The thick crust and broad crust–mantle transition beneath the Lachlan Fold
Belt may be caused by magmatic underplating of mafic materials beneath the
crust, which is consistent with an elevated Vp/Vs ratio (relative to ak135) of
∼ 1.73. Thicker crust is also to be expected from the elevated
topography of the eastern Lachlan Fold Belt.
The crustal structure is complex beneath VanDieland. It thins considerably
from the northern tip of the microcontinent (∼ 37 km) into
Bass Strait (∼ 26 km) and northern Tasmania (∼ 23 km), yet the crust is somewhat thicker in southern Tasmania
(∼ 33 km) compared with Bass Strait. This may in part be due to
the complex origins of the microcontinent, which appears to be comprised of
multiple Precambrian continental ribbons, but is also likely due to failed
rifting in Bass Strait before and during the separation of Australia and
Antarctica. This resulted in lithospheric stretching and delamination, magmatic
intrusion, and the deposition of thick sedimentary sequences. Recent
intraplate volcanism and the possible progression of a mantle plume beneath
VanDieland in the last few thousand years may also have produced
compositional heterogeneity and melt in the crust. Such events are likely to
contribute significantly to variations in crustal thickness and the
pronounced changes in Vp/Vs that we observe.
Stations within the ETT + EB collectively indicate crust of uniform
thickness (∼ 31–32 km) and uniform Vp/Vs (∼ 1.70), which
clearly distinguishes it from VanDieland. This region of the crust likely
represents a southern continuation of the Lachlan Orogen and is, therefore,
underpinned by crust of oceanic origin.
Comparison of our new Moho depth results with the AusMoho model reveals an
overall consistency, although at some of our station locations where AusMoho
has few constraints, there are noticeable differences, such as southern
Victoria and beneath Flinders Island. The discrepancies beneath the Lachlan
Orogen are attributed to the presence of underplated mafic material, which
can obfuscate the location of the Moho.
A synthesis of our new RF results with pre-existing teleseismic tomography,
shear wave splitting, and ambient noise studies reveals a complex lithosphere
that has clearly been impacted by orogeny (thickened crust), failed rifting
beneath Bass Strait (thinned crust and complex crustal velocities), and
recent intraplate volcanism (high Vp/Vs ratios and a radial pattern of fast
anisotropy patterns above a presumed zone of mantle upwelling).
Data availability
The dataset is available at 10.6084/m9.figshare.12233723 (Bello et al., 2020).
The supplement related to this article is available online at: https://doi.org/10.5194/se-12-463-2021-supplement.
Author contributions
MB performed the data analysis and wrote the draft of the paper. NR and
DGC guided the study and assisted with data interpretation. MB, DGC, and NR
discussed the results and revised the paper. AMR and OKL revised the
paper and assisted with the data interpretation.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
The work in this paper was performed as part of a PhD study and has been
jointly funded by Abubakar Tafawa Balewa University (ATBU), Bauchi, Nigeria,
and the University of Aberdeen, UK. The authors acknowledge the efforts of
staff, students, and fieldwork technicians from the Australian National
University and the University of Tasmania, who deployed the temporary BASS array
used in this study. We also thank Qi Li and Armando Arcidiaco for their
efforts in BASS data preprocessing and archiving. We are grateful to
IRIS and Geoscience Australia for providing data from several stations on
mainland Australia and in Tasmania. Figure 1 was made using Inkscape software
(The Inkscape Project, 2020), and Figs. 2, 3, 6, and 9 were produced using
the Generic Mapping Tools (Wessel et al., 2013).
Financial support
This research has been supported by the Australian Research Council (grant no. LP110100256).
Review statement
This paper was edited by Irene Bianchi and reviewed by two anonymous referees.
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