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Structural evolution of the Texas Orocline, eastern Australia
Pengfei Li a,⁎, Gideon Rosenbaum a, Paul J.T. Donchak b
a School of Earth Sciences, The University of Queensland, Brisbane 4072, Queensland, Australiab Geological Survey of Queensland, Department of Employment, Economic Development and Innovation, Brisbane 4000, Queensland, Australia
Article history:Received 4 March 2011Received in revised form 30 August 2011Accepted 11 September 2011Available online 1 October 2011
Handling Editor: M. Santosh
Keywords:Texas OroclineEastern AustraliaNew England OrogenTectonicsSubduction rollback
The Texas Orocline, with a half wavelength of ~120 km, is the largest and most obvious orocline in the south-ern New England Orogen and is clearly recognised in geological maps and geophysical images. In the area ofthe orocline, there is a major unconformity between Devonian–Carboniferous metasedimentary rocks (Texasbeds) and the overlying Early Permian rift-related basins. Detailed structural mapping shows that units bothabove and below the unconformity are folded around the orocline, indicating that at least part of the oroclinaldeformation has occurred after the deposition of the Early Permian rocks. In addition, Early Permian (298–290 Ma) granitoids are aligned parallel to the oroclinal structure, further indicating that deformation partlyoccurred during or after the Early Permian. Pre-oroclinal isoclinal folds (F1) and related axial plane cleavage(S1) are well developed in the Texas beds and are curved around the oroclinal structure. Syn-oroclinal struc-tures are characterised by minor kink folds and disharmonic folding (F2) in the core of the orocline. A laterphase of kink folding corresponding to a ~N–S shortening direction does not seem to be related to the oro-cline, and is interpreted as post-oroclinal deformation. A secondary penetrative fabric parallel to the axialplane of the orocline was not observed, indicating low contractional shortening across the orocline (b30%).We demonstrate that the observed strain is too low to account for oroclinal bending during dextral transpres-sion, as previously proposed. We suggest an alternative model involving an initial curved structure, probablyrelated to subduction rollback or a pre-existing curvature in the palaeomargin of eastern Australia, which wasamplified by dextral transpression and subsequent E–W contraction.
The term orocline was defined by Carey (1955) as a quasi-linearorogenic system that has been flexed to a horseshoe or elbow shapeduring subsequent deformation. Since then, numerous studies haveattempted to understand the mechanism of oroclinal bending (Mar-shak, 1988; Van der Voo, 2004; Weil, 2006; Johnston and Mazzoli,2009; Xiao et al., 2010). Some authors have shown that oroclinescan form in response to along-strike obstacles in the foreland of athin-skinned fold-and-thrust belt (Marshak, 1988, 2004), whereasothers suggested mechanisms involving orogenic-scale buckling inresponse to plate convergence (Johnston and Mazzoli, 2009). Alterna-tively, the formation of many oroclines has been attributed to subduc-tion curvature controlled by changes in rollback velocity along thesubduction segment (Schellart et al., 2007). The formation of suchoroclines is not restricted to deformation in shallow crustal levels,but can involve lithospheric-scale deformation (Lucente and Sper-anza, 2001; Rosenbaum and Lister, 2004). The actual subduction cur-vature could be promoted by torodial asthenospheric flow at the slabedges (Schellart, 2004; Loiselet et al., 2009), and/or differential
ssociation for Gondwana Research.
rollback velocities associated with along-strike variations in the den-sity of the subducting lithosphere (Morra et al., 2006; Rosenbaum andMo, 2011). In order to understand the fundamental mechanisms con-trolling oroclinal bending, it is important first to identify internalstructures within the orocline and to relate these structures to thecorresponding strain associated with oroclinal bending.
A series of oroclines is found in the southern New England Orogen ineastern Australia (Korsch and Harrington, 1987) (Fig. 1a). The Texas Oro-cline is the largest andmost obvious one and is clearly recognised in geo-physical images (Fig. 1b, c). The eastern limb of the Texas Orocline isconnected to another curved structure, the Coffs Harbour Orocline, form-ing a large (wavelength of ~250 km) Z-shaped structure. Farther south,an additional oroclinal structure (Manning Orocline, Fig. 1a) has been de-scribed (Korsch and Harrington, 1987). Recently a fourth orocline (Nam-buccaOrocline) has beenproposedbased on the curved geometry of EarlyPermian granitoids (Fig. 1a) (Rosenbaum, 2010; Rosenbaum, in press).
This paper focuses on the structure of the Texas Orocline withthe aim of understanding the mechanism of oroclinal bending andthe way that strain was distributed during deformation. Previousstudies have shown that the oroclinal structure is delineated by thecurvature of the bedding and fabric orientations (Lucas, 1960; Len-nox and Flood, 1997). Here we provide new structural data support-ing the existence of the orocline and analyse overprintingrelationships between mesoscopic folds and fabric elements. We
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show that regardless of the remarkable shape of the Texas Orocline,oroclinal-related deformation is characterised by relatively lowstrain.
2. Geological setting
The New England Orogen is the easternmost and youngest compo-nent of the Tasmanide orogenic collage of eastern Australia (Glen,2005; Offler and Murray, 2011). It comprises a Devonian to Carbonif-erous volcanic arc, forearc basin and accretionary wedge terranes, at-tributed to the convergent plate boundary of eastern Gondwana(Terra Australis, Cawood, 2005). The orogen can be subdivided intoa northern segment and a southern segment, separated by Mesozoicsedimentary rocks of the Clarence–Moreton Basin (Fig. 1a). Thesouthern New England Orogen constitutes a weakly deformed forearcbasin in the west (Tamworth Belt) and accretionary wedge metasedi-mentary rocks in the east (Tablelands Complex) (Fig. 1a). These are
Fig. 1. (a) Geological map of the southern New England Orogen based on 1:250,000 map sheand Grafton) and 1:100,000 map sheets (Texas, Stanthorpe, Inglewood and Allora). NB, NaNO, Nambucca Orocline. The blue dashed line indicates the general trend of the oroclinal strshows the trend of the orocline under the Mesozoic Surat Basin as inferred from the total maB, Bondonga; PMF, Peel-Manning Fault. Ages of granitoids are after Donchak et al. (2007) adata by Rosenbaum, Li and Rubatto). (c) Total magnetic intensity image showing the Texas
separated by the Peel-Manning Fault System (Voisey, 1959; Leitch,1974; Korsch, 1977).
In the area of the Texas Orocline, the Tamworth Belt is predomi-nantly covered by younger sedimentary rocks (Wartenberg et al.,2003), and is exposed only in the eastern limb of the orocline (EmuCreek Block, Fig. 1b). The accretionary wedge rocks (Tablelands Com-plex) in this area are widely exposed and are represented by theTexas beds (Olgers et al., 1974), which consist of a sequence of volca-niclastic turbidite with minor chert and jasper, altered mafic volcanicrocks and limestone (Flood and Fergusson, 1982; Fergusson andFlood, 1984; Donchak et al., 2007).
The Texas beds are overlain by Early Permian clastic sedimentaryand volcanic rocks (Roberts et al., 1996; Donchak et al., 2007). Theserocks have been interpreted to be deposited in rift basins, associatedwith widespread extension that affected eastern Australia during theEarly Permian (Korsch et al., 2009a). A rhyolitic layer from the baseof one of these basins (Alum Rock; Fig. 1b) yielded a zircon age of
ets (Singleton, Newcastle, Tamworth, Hastings, Manilla, Dorrigo-Coffs Harbour, Inverellmbucca Block; TO, Texas Orocline; CO, Coffs Harbour Orocline; MO, Manning Orocline;ucture (after Rosenbaum 2010). (b) Geological map of the Texas Orocline. Dashed linegnetic intensity image. SS, Silver Spur; W,Waroo; T, Terrica; AR, Alum Rock; P, Pikedale;nd Cross et al. (2009), except of Mt You You and Bundarra (unpublished U-Pb SHRIMPOrocline. Dashed line indicates the trend of the axial plane of the orocline (~345°).
281P. Li et al. / Gondwana Research 22 (2012) 279–289
~291 Ma (Roberts et al., 1996). The youngest sedimentary rocks inthese basins are Artinskian strata (~285–275 Ma) (Briggs, 1993,1998).
Contemporaneously with the deposition of the Early Permian riftbasins, S-type granitoids have been emplaced (Shaw and Flood,1981). A series of such granitic plutons in the eastern limb (Bullaga-nang, Mt You You, Ballandean, and Jibbinbar; Fig. 1b), together withthe Bundarra Granite in the western limb, mimics the shape of theTexas Orocline (Rosenbaum, 2010, in press). A later phase of magma-tism, involving voluminous I-type granitoids and arc-related volcanic
Fig. 2.Map of representative bedding orientations (S0) in the study area. Stratigraphic boundarea is divided into five sub-domains. Inset stereographic plots (lower hemisphere, equal abedding orientation in Domain 2. (c) E–Wbedding orientation in the hinge zone around thein Domains 4 and 5. (f) and (g) Bedding orientations in Alum Rock and the Waroo beds motations in Terrica beds moderately dipping to the NNE and SW.
rocks, intruded at 260–220 Ma (Shaw and Flood, 1981; Bryant et al.,1997; Shaw and Flood, 2009). The general trend of this magmaticfield is NE–SW, crosscutting the oroclinal structure.
3. Structural observations
Structural mapping was conducted in the area of the Texas Oro-cline, in both the Texas beds and overlying Early Permian sedimenta-ry rocks. In the following section, we present observations of primary
aries are after the 1:100,000 map sheets (Texas, Stanthorpe, Inglewood and Allora). Therea) show poles to bedding. (a) NW–SE bedding orientation in Domain 1. (b) NE–SWarea of Mosquito Creek (Domain 3). (d) and (e) Local variations in bedding orientationsderately to shallowly dipping to southwest and south, respectively. (h) Bedding orien-
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structures, pre-oroclinal secondary structures, syn-oroclinal struc-tures and post-oroclinal structures.
3.1. Primary structures
Bedding (S0) in the Texas beds is steeply dipping or sub-verticaland shows variable strike orientations parallel to the oroclinal struc-ture (Fig. 2). In the eastern limb, the bedding strike orientation is
S0
S1
S1
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(a) (
(c) (
((e)
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W
E W
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SE NW
N
Fig. 3. Photographs of pre-oroclinal structures. (a) Field photograph showing shallo28.5156°S/151.3620°E); (b) Slaty cleavage (S1) in the Texas beds (GPS coordinates 28.1945(GPS coordinates 28.7040°S/151.2020°E); (d) Photomicrograph (crossed polarised light) shcoordinates 28.1783°S/151.7778°E); (e) Angular relationships between cleavage (S1) andSteeply plunging F1 fold in chert (GPS coordinates 28.1965°S/151.3577°E). (g) F1 fold in a qturbidite (GPS coordinates 28.6297°S/151.6678°E).
dominantly NW–SE (Fig. 2a), whereas in the western limb the orien-tation is dominantly NE–SW (Fig. 2b). In the hinge zone around thearea of Mosquito Creek (Domain 3; Fig. 2c), the bedding strike orien-tation is E–W. However, the hinge zone farther south (Domains 4 and5) shows local variations in bedding orientations (Fig. 2d, e) attribut-ed to syn-oroclinal folding (see below).
Within the overlying Early Permian sedimentary rocks, three areashave been investigated (Alum Rock, Terrica, and Waroo; Fig. 2).
µm001
S1
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S1
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b)
d)
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E
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wly dipping bedding (S0) in the Early Permian Waroo beds (GPS coordinates°S/151.3500°E); (c) Dominant cleavage (S1) defined by pressure solution in limestoneowing dominant cleavage (S1) defined by the preferred orientation of mica grains (GPSbedding (S0) in fine grained sandstone (GPS coordinates 28.2004°S/151.3554°E). (f)uartz vein (GPS coordinates 28.2272°S/151.3497°E). (h) Shallowly plunging F1 fold in
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Bedding in the Alum Rock Formation, characterised by conglomeratelayers and interbedded arenite, siltstone and volcaniclastic layers,dips moderately to the southwest (Fig. 2f), which coincides withdata reported by Lennox and Flood (1997). The Waroo beds (newname) mainly comprise thick arenite layers interlayered with thinsiltstone bands. Bedding strikes E–W and is shallowly dipping to thesouth (Figs. 2g, 3a). The variations in Early Permian bedding strikeorientations from NW–SE in Alum Rock to E–W in the Waroo bedsseem to coincide with the bedding changes in the Texas beds aroundthe oroclinal structure. It seems likely, therefore, that Early Permianrocks have also been affected by the oroclinal deformation, as previ-ously suggested by Olgers et al. (1974). In the Terrica beds, inter-bedded conglomerate, arenite, mudstone and siltstone layers aremoderately dipping to the NNE and SW (Fig. 2h). These variations
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Fig. 4.Map of representative orientations of the dominant fabric (S1). Stratigraphic boundaridivided into five sub-domains. Inset stereographic plots (lower hemisphere, equal area) shomain 2; (c) and (d) E–W cleavage in Domains 3 and 4. (e) Variable cleavage orientations in
possibly correspond to folding with a moderately NW plunging foldaxis.
In summary, our data show that bedding in the Early Permian rift ba-sins is considerably shallower than in the underlying Texas beds, in agree-ment with Lennox and Flood (1997). These relationships indicate anangular unconformity or a fault contact between the two successions.
3.2. Pre-oroclinal secondary structures
A dominant axial plane foliation (S1) is developed throughout theTexas beds. The foliation is characterised by slaty cleavage associatedwith pressure solution (Fig. 3b, c, d). It is normally steeply dippingand parallel or sub-parallel to bedding. In some places, low angle ver-gence relationships are recognised (Fig. 3e). Similarly to the bedding,
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granitoids
S trace1
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e)
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es are after 1:100,000 map sheets (Texas, Stanthorpe, Inglewood and Allora). The area isw poles to cleavage (S1). (a) NW–SE cleavage in Domain 1; (b) NE–SW cleavage in Do-Domain 5.
284 P. Li et al. / Gondwana Research 22 (2012) 279–289
the general orientation of the dominant foliation (S1) is parallel to theshape of the orocline (Fig. 4). In the eastern limb of the orocline (Do-main 1), the cleavage (S1) is consistently striking NW–SE (Fig. 4a);whereas in the western limb (Domain 2), S1 is predominantly strikingNE–SW (Fig. 4b). Similarly to the behaviour of S0, S1 in the hinge zoneof the orocline (Domain 3) strikes E–W (Fig. 4c), whereas in Domains4 and 5 it varies due to syn-oroclinal smaller-wavelength folding(Fig. 4d, e).
F1 folds are recognised by the changes in bedding/cleavage ver-gence relationships (Fig. 3e), by minor folds in the bedding (Fig. 3f)and folding of earlier quartz veins and veinlets (Fig. 3g). The foldsare typically tight to isoclinal. They are non-cylindrical, with foldaxes predominantly plunging steeply, but locally becoming shallower(Fig. 3f, h). The axial planes of these folds are steeply dipping, and areconsistently parallel to the dominant cleavage (S1).
The S1 foliation is not recognised in the overlying Early Permian rocks,indicating that this generation of folding had occurred prior to the depo-sition of the Early Permian strata. The S1 fabric ismost likely related to de-formation within the accretionary wedge during the Carboniferous.
3.3. Syn-oroclinal structures
The form of the orocline is demonstrated by the distribution of bed-ding (S0) and dominant cleavage (S1) as discussed above. The axis of the
151°E
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‘ 01°8
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(Axi
la
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no
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fht
ol
orc
ine
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Fig. 5. (a) Mean directions of poles to S1 cleavage in Domains 1–4. The pole to girdle (β1) iscalculated from the axis of the orocline (β1=80°–255°) and its strike orientation (345°) takeroughly consistent with this axial plane orientation. (c) Map of S1 cleavage in the core of thedisharmonic folding (F2) with axial plane trending approximately NNW. The inset stereog(within the area of Fig. 5c) and the corresponding β-axis; (d) Map showing axial planes tothe syn-oroclinal minor folds (F2), striking NNE in the western limb (d1) and NW in the eaaxial planes of post-oroclinal minor folds (F3).
orocline can be calculated by considering the poles to the mean princi-pal orientations of the dominant cleavage (S1) in Domains 1, 2, 3 and 4(Fig. 5a). These poles define a girdle, and the β-axis that corresponds tothis girdle is 80°–255°, indicating a steeply plunging fold hinge for theorocline (β1 in Fig. 5a, b). Together with the strike orientation of theaxial plane, as seen in amap view (345°, Figs. 1c and 4),we can calculatethe attitude of the axial plane as 80°–255°.
Folds related to the orocline (F2) are not common and there is nopenetrative axial plane foliation associated with them. However,within the core of the orocline (Domain 5) we mapped local varia-tions in the orientation of S1 (Figs. 4 and 5c). These variations seemto correspond to large-scale (10–20 km) folds, with β-axis of 81°–239°, which roughly lies on the axial plane of the orocline (β2 inFig. 5b, c). In addition, minor F2 folds are locally found in the Texasbeds. These are characterised by kink folds and local crenulations,which overprint the dominant foliation (Fig. 6a, b). The axial planesof F2 minor folds are predominantly striking NW–SE on the easternlimb of the orocline, approximately parallel to orientation of theaxial plane of the orocline (Fig. 5d). However, two observations ofminor folds on the western limb of the orocline show an axial planeorientation of NNE–SSW (Fig. 5d). We interpret the inconsistent ori-entation of F2 folds and the lack of a penetrative axial plane foliationas indicators for a relatively low strain (b30% shortening) duringoroclinal deformation. An alternative explanation for the inconsistent
151°10’ 151°20’ 151°30’ 151°40’ 151°50’
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3
rocks GranitoidsAxial plane of F2 fold
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interpreted as the axis of the orocline (80°–255°); (b) The axial plane of the orocline,n from the map pattern of foliations (see Fig. 4). The hinge of F2 folds (β2=81°–239°) isorocline (see location in Fig. 4). The local variations in S1 orientations are interpreted asraphic diagram (lower hemisphere, equal area) shows poles to dominant fabric (S1)minor folds around the Texas Orocline. Inset stereographic plots show axial planes ofstern limb (d2) of the orocline respectively. Inset d3 shows stereographic plots of the
S0
S 84°-281°2:
S1
S2
S3
(c) (d)
S2
S1
(b) 200 mµ
E W N S
N
(a)
Fig. 6. Photographs of syn- and post-oroclinal structures. (a) F2 kink fold overprinting the dominant fabric (S1) in slate (Texas beds, GPS coordinates 28.4090°S/151.4451°E). (b)Photomicrograph (crossed polarised light) of a crenulation cleavage (S2) overprinting the dominant fabric (S1) (Texas beds, GPS coordinates 28.1993°S/151.2843°E). (c) F2 foldsin Early Permian rocks (Silver Spur basin, GPS coordinates 28.8508°S/151.2316°E). (d) Post-oroclinal spaced fabric (S3) in Early Permian granite (Mt You You, GPS coordinates28.6752°S/151.5810°E).
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orientations of the kink folds is that some of these folds had in factformed prior to oroclinal bending.
Minor F2 folds are rare in the Early Permian rocks. In the SilverSpur basin (Fig. 2), we found one mesoscopic fold trending ~N–S(Fig. 6c). Similar folding of Early Permian rocks is expressed in thecurvature of bedding orientation in the Waroo and Alum Rock ba-sins. This demonstrates that Early Permian sedimentary rocks wereaffected by E–W shortening, which could probably be attributed toa late stage in the development of the orocline. Lennox and Flood(1997) reported that macroscopic folds in the Silver Spur basin hadgently plunging fold axes, indicating that folding occurred after de-position of the Early Permian rocks, consistently with our interpreta-tion. The contrast between steeply plunging folds in the Texas bedsand gently plunging folds in the Early Permian rocks indicates thattilting of the former occurred prior to the deposition of the EarlyPermian basins.
3.4. Post-oroclinal structures
A number of minor kink folds have axial planes oriented perpendic-ular to the axial plane of the orocline (Fig. 5d). These folds cannot be re-lated to the orocline and are interpreted to represent post-oroclinaldeformation. An E–W structural fabric (S3), defined by spaced foliationand no mica growth, was recognised in the Early Permian granitoids(Fig. 6d), and could also be related to this deformation.
4. Discussion
4.1. Timing of oroclinal bending
Our data show that the Early Permian rift-related sedimentaryrocks were affected by E–W contraction, which we attribute to alate stage in the development of the orocline. These rocks were de-posited contemporaneously with the emplacement of S-type granit-oids, which are aligned parallel to the orocline. The age of thegranitoids and sedimentary rocks is ~300–275 Ma (Robert et al.,
1996; Donchak et al., 2007; Briggs 1993, 1998), indicating that atleast part of the oroclinal structure developed after ~275 Ma. Howev-er, based on palaeomagnetic data (Aubourg et al., 2004), it seems thata component of the oroclinal deformation had already occurred priorto ~291 Ma, which is the age of interstratified volcanic rocks in theEarly Permian Alum Rock basin (Fig. 1b).
The oroclinal structure is crosscut by Late Permian to Triassic(260–220 Ma) granitoids and volcanic rocks of a calc-alkaline affinity.The orocline, therefore, had already existed prior to the emplacementof these rocks. Further evidence for the termination of oroclinal defor-mation prior to the Late Permian has been demonstrated by palaeo-magnetic data, showing that Late Permian rocks (GilgurryMudstone) have not been affected by the orocline (Aubourg et al.,2004).
The relationship between the oroclinal deformation and theHunter-Bowen orogeny remains enigmatic. Our observations showthat at least part of the oroclinal deformation took place after~275 Ma. Contractional deformation associated with the Hunter-Bowen orogeny commenced at ~265 Ma, and was active until theMiddle Triassic (~230 Ma) (Holcombe et al., 1997). The bulk of thisdeformation occurred contemporaneously with calc-alkaline magma-tism, which, as mentioned above, overprinted the oroclinal structure.It is possible, however, that the initiation of the Hunter-Bowen orog-eny at ca. 265 Ma was responsible for further contraction in the re-gion, leading to a later stage of tightening of the oroclinal structure.
Offler and Foster (2008) have recently proposed that the develop-ment of the Texas and Coffs Harbour Oroclines took place from 273 to260 Ma. This time bracket is generally in agreement with our inter-pretation for the late stage of oroclinal bending. However, accordingto Offler and Foster (2008), oroclinal bending was controlled by asouthward tectonic transport of the Coffs Harbour Orocline that ledto the development of E–W fabric in the Early Permian NambuccaBlock (Fig. 1a). Based on 39Ar/40Ar geochronology, the timing of thisfabric development was estimated at 264–260 Ma (Offler and Foster,2008). Whether this fabric is directly related to the oroclinal deforma-tion is an open question, which we discuss in the following section.
0 100km0 100km
Pin Pin
(a) (b)
~500km
TexasOrocline
CoffsHarbourOrocline
NambuccaBlock
X
Z
Fig. 7. Schematic illustration of a model that is based on oroclinal bending in response to a strike-slip faulting offshore eastern Australia (e.g. Offler and Foster, 2008). (a) A possiblepre-oroclinal shape of the Texas Orocline. (b) Post-oroclinal structure showing the required dextral displacement (~500 km) and the imposed strain (blue ellipse). The calculationof the principal stretching axes (X and Z) is based on the assumption of plane strain and no volumetric strain. From the finite strain ellipse, we measure X/Z=5.61, and given thatX·Z=1 (area conservation), we calculate that Z=0.42. This Z value corresponds to maximum shortening of 58%. Yellow area shows the position of the Early Permian NambuccaBlock and its dominant ~E–W fabric (red dashed lines).
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4.2. Style of oroclinal deformation and geodynamic implications
Our data clearly support the existence of the orocline, as indicatedby the orientations of S0 and S1 and the occurrence of F2 folds. However,fabric parallel to the axial plane of the orocline was not observed, nei-ther havewe observed high strain zones that might reflect strain locali-sation during oroclinal bending. These observations suggest that theoroclinal deformation has incorporated relatively low strain. We havenot found evidence for widespread patchy S2 foliations fanning aroundthe orocline, as described by Lennox and Flood (1997).
The geodynamic environment responsible for the oroclinal defor-mation is still an unsolved problem. Previous models (e.g. Murray etal., 1987; Offler and Foster, 2008) suggested that oroclinal bendingoccurred in response to oblique plate convergence. According tothese models, the Z-shaped structure of the Texas and Coffs HarbourOroclines represents a mega-drag fold associated with dextralstrike-slip faulting. This scenario, which is analogous to the origin ofthe New Zealand Orocline (Kamp, 1987), is schematically illustratedin Fig. 7, assuming a dextral-strike slip fault located offshore as pro-posed by Offler and Foster (2008). The E–W fabric in the NambuccaBlock was interpreted by Offler and Foster (2008) as the result ofthis southward tectonic transport (Fig. 7b). In this scenario, approxi-mately 500 km of dextral displacement is required in order to achieve
the oroclinal structure (Fig. 7b). If we assume plane strain and no vol-umetric strain, we can calculate the predicted strain ellipsoid associ-ated with this deformation (Fig. 7). Our calculation shows that suchdeformation would result in shortening of more than 50%. Such a de-gree of shortening deformation would certainly result in the develop-ment of axial plane cleavage. However, our observations show a lackof penetrative cleavage parallel to the axial plane of the orocline.
An alternative mechanism for the development of the New En-gland Orocline has recently been suggested by Cawood et al. (2011).These authors suggested a model, in which fore-arc fragmentsmoved northward with respect to cratonic Gondwana, and progres-sively buckled around a pinning point located between the Texasand Coffs Harbour Oroclines. The model of Cawood et al. (2011) isconsistent with available palaeomagnetic data, but similarly to thedextral-strike slip model, is implying relatively high strain in thearea of the Texas Orocline, which is in disagreement with ourobservations.
It appears, therefore, that buckling alone cannot account for thedevelopment of the Texas Orocline. We therefore propose a possiblescenario in which orocline developed in multiple stages (Fig. 8). Theinitial curvature could have formed in a subduction environment, ina similar fashion to the curvature of modern subduction zones(Morra et al., 2006; Rosenbaum and Mo, 2011). These processes are
0 100km
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(b) ~280-265Ma
Bowenbasin
Fig. 8. Schematic diagrams showing a possible scenario for oroclinal bending in three stages. (a) Early curvature of the Texas and Coffs Harbour oroclines is acquired by along-strikevariations in rollback velocities, accompanying the formation of Early Permian rift basins and exhumation of metamorphic complexes. The blue line represents inferred ~E–W tearfaults. T &W, Tia andWongwibinda metamorphic complexes; MM, Mt Mee metamorphic complex; (b) Further curvature in response to dextral strike slip faulting; (c) Final stage oforoclinal deformation in response to ~E–W shortening.
287P. Li et al. / Gondwana Research 22 (2012) 279–289
particularly common in subduction segments undergoing subductionrollback, and are commonly accompanied by back-arc extension(Royden, 1993; Lonergan and White, 1997; Rosenbaum and Lister,2004). Indeed, there is widespread evidence for extensional tectonicsin eastern Australia during the Early Permian (Korsch et al., 2009a),and at least part of this extension has been attributed to subductionrollback (Little et al., 1992; Jenkins et al., 2002). It is therefore possi-ble that the initial curvature of the orocline was acquired during theEarly Permian by along-strike variations in rollback velocities(Fig. 8a). According to this model, this early stage of subduction cur-vature occurred contemporaneously with the formation of the EarlyPermian rift basins, the exhumation of metamorphic complexes, andthe emplacement of S-type granitoids (Fig. 8a).
Alternatively, it is possible that an initial curvature was inheritedfrom the palaeomargin of the Australian continent prior to the develop-ment of New England Orogen. Geophysical data from northwesternNew SouthWales (Fig. 9a), west of the Texas Orocline, show prominent~E–W lineaments, which are essentially perpendicular to the ~N–S ori-entation of the dominant subduction-related features in eastern Austra-lia (Glen, 2005; Burton, 2010). These E–W structural elements,attributed by some authors to the boundary between the Thomsonand Lachlan Orogens (Glen, 2005), could have led to the primary curva-ture of the continentalmargin duringNewEngland orogenesis (Fig. 9b).It is also possible that this primary curvature was further overprintedand amplified during the Early Permian by along-strike variations inrollback velocities as discussed before.
The initial curvature is consistent with palaeomagnetic results fromAlum Rock , which indicate that rotation of ~40° had occurred prior to~291 Ma, and that further rotation of ~80° occurred by subsequent de-formation (Aubourg et al., 2004). The earlier rotation of ~40°was possi-bly associated with an early stage of arc/subduction zone curvature,while the subsequent rotation of ~80° could be attributed to oroclinalamplification by dextral strike-slip faulting and E–W shortening duringthe early stage of theHunter-Bowen orogeny (Figs. 8b, c and 9c, d) (Col-lins, 1991; Holcombe et al., 1997; Korsch et al., 2009b).
We note that the rollback model requires ~E–W crustal-scale tearfaults (Fig. 8), which have not been identified in the current study.Such structures are difficult to recognise, because the northern part ofthe Texas Orocline is concealed by the Mesozoic Surat and Clarence–
Moreton Basin (Fig. 1a). The existence of such possible structurescould be tested in future reconstruction models, which will require fur-ther constraints from geophysical data on the structure of the TexasOrocline under cover, complemented by information on the timingand geochemistry of magmatism and palaeomagnetic data on verticalaxis block rotations.
5. Conclusions
The Texas Orocline is defined by the curvature of bedding (S0) anddominant sub-parallel fabric (S1). Syn-oroclinal structures are charac-terised by minor crenulations and minor folds (F2), particularly in thecore of the orocline, with varying axial plane orientations. A penetra-tive axial plane cleavage was not developed during oroclinal defor-mation. These features demonstrate that the oroclinal deformationinvolved relatively low strain (b30% shortening). A post-oroclinalphase of kink folding occurred in response to mild ~N–S shortening.
The tectonic processes responsible for oroclinal bending remainenigmatic. A model purely based on mega-folding in response to dex-tral strike-slip faulting, previously proposed in the literature (Murrayet al., 1987; Offler and Foster, 2008), is not supported by the observedlow strain structures. We therefore argue that the oroclinal deforma-tion has involved multiple stages. The initial curvature may haveformed in response to subduction rollback, which was accompaniedby the formation of Early Permian rift basins and the emplacementof S-type granitoids. Alternatively, it is possible that the initial curva-ture was partly inherited from an arcuate palaeomargin of easternAustralia. The initial curvature reduces the required strain of the sub-sequent deformation events, whichmay have involved drag folding inresponse to dextral strike-slip faulting (Murray et al., 1987; Offler andFoster, 2008) and ~E–W shortening during the early stage of theHunter-Bowen orogeny.
Acknowledgements
Funded by the Australian Research Council (DP0986762). Themanuscript benefited from constructive comments by two anony-mous reviewers. We thank Ben Wruck and Adrienna Brown for assis-tance in the field, and Rod Holcombe and Hongyuan Zhang for
Bow
enn
,yS
Gu
nn
edah
&d
eybas
sin
ThomsonOrogen
LachlanOrogen
Te
xas Orocline
C
offs Harbour
Oro
c lin
e
NewEnglandOrogen
Queensland
New South Wales
8 S2 °
30°S
32°S
6 S2 °
°8 E41°E441 °6 E41 1 E50° 1 E52° 1 4 E5 °
Brisbane
Newcastle
ThomsonOrogen
LachlanOrogen
ThomsonOrogen
LachlanOrogen
NewEnglandOrogen
NewEnglandOrogen
(a)
(b) (c) (d)
ThomsonOrogen
LachlanOrogen
NewEnglandOrogen
Fig. 9. (a) Total magnetic intensity image showing a prominent ~E–W lineament to the west of the New England Orogen.White bold lines represent the boundaries of tectonic elements(after Glen, 2005).White fine lines illustrate structural lineaments to the west of New England Orogen (after Gray and Foster, 2004); (b) A schematic diagram showing a primary curvedsubduction zone in response to a pre-existing E–Wboundary between the Lachlan and Thomson Orogens; (c) Further curvature in response to dextral strike faulting; (d) Final phase oforoclinal bending in response to ~E–W contraction.
288 P. Li et al. / Gondwana Research 22 (2012) 279–289
discussions on the structural interpretation and tectonic model. Wealso wish to thank local landowners for permitting us to visit andmap in their properties.
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