Introduction
Australo-Antarctica comprises several geophysically distinct Precambrian crustal domains. Various attempts were made to reconstruct the Precambrian basement geology of Antarctica and Australia based on the continuity of comparable geological terranes, major shear zones and geophysical signatures (e.g. Aitken et al. Reference Aitken, Betts, Young, Blankenship, Roberts and Siegert2016). However, the paucity of knowledge regarding regional variations in sub-ice geology within the composite East Antarctic Craton (e.g. Harley et al. Reference Harley, Fitzsimons and Zhao2013) continues to hamper determinations of the extent of Australian geological entities in the interior of the Antarctic Wilkes Land terrain.
Gleaning an improved understanding of the subglacial geology in East Antarctica is also critical for the development of new models and reconstructions to assess the stability of the East Antarctic Ice Sheet (EAIS) within future warmer climate scenarios (e.g. DeConto et al. Reference DeConto, Pollard and Harwood2007, Pollard & DeConto Reference Pollard and DeConto2009, Reference DeConto and Pollard2016, Pollard et al. Reference Pollard, DeConto and Alley2015). The Wilkes Subglacial Basin (WSB) sector of East Antarctica is of particular importance for assessments of past and future EAIS stability (Fig. 1; e.g. Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009, Pandey et al. Reference Pandey, Pant, Arora and Gupta2021). The bedrock beneath the WSB lies mostly below sea level, and in several regions the bedrock deepens inland, making this sector of East Antarctica potentially more prone to marine ice-sheet instability processes. However, the development of robust models of past EAIS behaviour requires validation through both geophysical and geological datasets.
Fig. 1. Geological map of exposed bedrock around East Antarctica in proximity to Wilkes Land (inferred by Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009, Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013) displayed over a subglacial bedrock topography map (from Bedmap2; Fretwell et al. Reference Fretwell, Pritchard, Vaughan, Bamber, Barrand and Bell2013) to indicate probable sediment provenance areas for the sediments we investigated at site U1359 of Integrated Oceanic Drilling Program Expedition 318. CB = Central Basin; EB = Eastern Basin; MSZ = Mertz Shear Zone; WB = Western Basin.
Our knowledge of the subglacial geology and topography in Wilkes Land has improved significantly due to aero geophysical surveys such as the International Collaborative Exploration of the Cryosphere through Airborne Profiling Project (ICECAP; e.g. Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014, Frederick et al. Reference Frederick, Young, Blankenship, Richter, Kempf, Ferraccioli and Siegert2016) and Icehouse Earth: Stability or Dynamism/Wilkes Basin/Transantarctic Mountains System Exploration (ISODYN/WISE; e.g. Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009, Jordan et al. Reference Jordan, Ferraccioli, Vaughan, Holt, Corr, Blankenship and Diehl2010). However, the inferences from geophysical methods still require further confirmation and verification through direct geological evidence to directly probe and identify hidden geological terranes.
Linkage of the Wilkes Land sector of East Antarctica and its Australian counterpart based on the geophysical and geological inferences from coastal exposures suggest the presence of large-scale Palaeoproterozoic basements accreted during supercontinent reassembly (Aitken et al. Reference Aitken, Betts, Young, Blankenship, Roberts and Siegert2016). The Wilkes Land sector itself can be divided into three components: 1) the west Wilkes province occupies the region west of the Totten Glacier from Windmill Islands to the west of Law Dome comprising Sabrina Subglacial Basin (Morrissey et al. Reference Morrissey, Payne, Hand, Clark, Taylor, Kirkland and Kylander-Clark2017), 2) central Wilkes Land from the Totten Glacier to the Mertz Suture Zone comprising Frost Subglacial Basin and Astrolabe Subglacial Basin (Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014) and 3) eastern Wilkes Land with WSB, which predominantly comprises Archaean to Palaeoproterozoic rocks exposed in Terre Adélie Land and George V Land and has geological affinities with the Gawler Craton of South Australia (Fitzsimons Reference Fitzsimons2003, Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014). However, these inferences are mostly from the unconstrained projection of the Australian geological terrain in the interior of the Wilkes Land terrain. These projections are unable to infer the presence of Ferrar Large Igneous Province (FLIP) in the WSB, which is suggested to be Jurassic in age purely constrained from the projection of Ferrar dolerites present in the Transantarctic Mountains (Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013).
Direct evidence of the regional geological constraints from the Wilkes sector is missing, which would be significant for understanding its thermo-tectonic history and sub-ice geological provinces. In Australo-Antarctic reconstruction, the Mertz Shear Zone (MSZ; Fig. 1) is a ~1500 Ma old mid-crustal shear zone (Lamarque et al. Reference Lamarque, Bascou, Maurice, Cottin, Riel and Ménot2016, Pierce et al. Reference Pierce, van de Flierdt, Williams, Hemming, Cook and Passchier2017) that has been inferred to align with the Kalinjala Shear Zone (KSZ) in South Australia (e.g. Reid & Hand Reference Reid and Hand2012). However, alternative Gondwana reconstructions favour linking the MSZ to the Coorong Shear Zone further to the east (Gibson et al. Reference Gibson, Totterdell, White, Mitchell, Stacey, Morse and Whitaker2013). There is a limited rock outcrop exposed east of the MSZ, which limits our direct knowledge of the sub-ice geology of this sector.
Marine sedimentary records recovered from the Antarctic continental margin provide the most accessible proxy information on the palaeoclimatic and tectono-thermal evolution of East Antarctica, as these sediments represent source rock domains (Pant et al. Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013, Gupta et al. Reference Gupta, Pandey, Arora, Pant and Rao2022). Heavy mineral fractions of the marine sediments reflect provenance characteristics, and the utility of heavy minerals in provenance and process interpretation is widely recognized (e.g. Bateman & Catt Reference Bateman and Catt2007, Mange & Wright Reference Mange and Wright2007). In addition, the influence of chemical weathering is relatively limited in Antarctica, especially post-formation of the EAIS since the Oligocene (e.g. Scher et al. Reference Scher, Bohaty, Zachos and Delaney2011). Therefore, specific minerals and mineral groups can be directly linked to particular domains of sub-ice provenance, and their chemical traits can imply the geology of these domains. The chemistry of various minerals, especially magmatic and metamorphic minerals, is predominantly governed by their genetic processes. Mineral compositions can thus be used to estimate physical conditions such as pressure and temperature for the rocks in the provenance. Our results provide new geological insights into correlations between East Antarctica and Australia, in particular during the Neoproterozoic, through heavy mineral fractions and rock fragments recovered from Integrated Oceanic Drilling Program (IODP) Expedition 318 by characterizing the lithologies of the Wilkes Land sector and comparing them to their Australian counterparts.
Geological overview of the area
Investigating the behaviour of the EAIS at the eastern margin of Wilkes Land was one of the major geoscientific objectives of IODP Expedition 318 (Escutia et al. Reference Escutia, Cooper, Eittreim, Tanahashi, Ishihara and O'Brien2011, Verma et al. Reference Verma, Bhattacharya, Biswas, Shrivastava, Pandey and Pant2014, Pandey et al. Reference Pandey, Pant, Biswas, Shrivastava, Joshi and Nagi2018). Recent geological interpretations derived from IODP Expedition 318 (e.g. Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013, Reference Cook, Hemming, van de Flierdt, Davis, Williams and Galindo2017, Verma et al. Reference Verma, Bhattacharya, Biswas, Shrivastava, Pandey and Pant2014, Pandey et al. Reference Pandey, Pant, Biswas, Shrivastava, Joshi and Nagi2018, Biswas et al. Reference Biswas, Pant, Pandey, Shrivastava, Panday, Ravichandran and Nair2020) augment earlier records from other Antarctic continental areas and margins indicating a relatively dynamic EAIS, especially during Late Miocene to Early/Middle Pliocene times (Naish et al. Reference Naish, Powell, Levy, Wilson, Scherer and Talarico2009, Verma et al. Reference Verma, Bhattacharya, Biswas, Shrivastava, Pandey and Pant2014, Pandey et al. Reference Pandey, Pant, Biswas, Shrivastava, Joshi and Nagi2018, Reference Pandey, Pant, Arora and Gupta2021, Biswas et al. Reference Biswas, Pant, Pandey, Shrivastava, Panday, Ravichandran and Nair2020). The general circulation model experiments of the Wilkes Land sector suggested that east wind drifts prevailed north of 60°S under Miocene boundary conditions, similar to today (Herold et al. Reference Herold, Huber and Müller2011, Bijl et al. Reference Bijl, Houben, Hartman, Pross, Salabarnada, Escutia and Sangiorgi2018). Hence, we could get sediments from as far as the Transantarctic Mountains (Pant et al. Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013, Cook et al. Reference Cook, Hemming, van de Flierdt, Davis, Williams and Galindo2017). The proximal interior gets sediments transported mainly through Mertz and Ninnis glaciers contributing from the interior of Wilkes Land (McMullen et al. Reference McMullen, Domack, Leventer, Olson, Dunbar and Brachfeld2006, Orejola et al. Reference Orejola and Passchier2014). All of the sediments and detritals further transported to the continental rise through the Jussieu submarine channel. Sites U1359 and U1361 were drilled on the eastern levee of this channel. The present work provides an account of the heavy phases and rock fragments of site U1359.
A marine sediment core was drilled from site U1359 during Expedition 318 at ~4003 m below sea level (mbsl) and ~100 km from the Antarctic shelf break. This site lies south of the Antarctic Polar Front and gets sediment supply through Jussieu submarine channel streams (Escutia et al. Reference Escutia, Cooper, Eittreim, Tanahashi, Ishihara and O'Brien2011). Because of the deepwater setting of the sedimentary archive, its stratigraphic record is relatively complete (Tauxe et al. Reference Tauxe, Stickley, Sugisaki, Bijl, Bohaty and Brinkhuis2012, Hansen et al. Reference Hansen, Passchier, Khim, Song and Williams2015).
Geology of site U1359
Four boreholes (U1359A–U1359D) were drilled at depths of 193.5, 252.0, 168.7 and 602.2 m below sea floor (mbsf), respectively, at site U1359. A composite lithology of site U1359 shows the dominance of silty clay with dispersed clasts throughout the core (Fig. 2). Five distinct lithofacies and three lithostratigraphic units were defined (Escutia et al. Reference Escutia, Cooper, Eittreim, Tanahashi, Ishihara and O'Brien2011) based on variations in biogenic components, bioturbation and sedimentary structures, packages of silt-fine sand laminae and diatom abundance. Details of the samples under investigation are also mentioned in Fig. 2. Lithostratigraphic Unit I (0–43.5 mbsf) consists of diatom-rich silty clays with dispersed clasts including occasional foraminifer-bearing clayey silt and sandy silt. Unit II (43.5–247.1 mbsf) is composed of interbedded bioturbated diatom-bearing silty clays and olive-grey diatom-bearing silty clays (Fig. 2). The lowermost part of Unit II (~200.0–247.1 mbsf) represents sediments of 6–9 Ma age. Unit III extends from 247.1 mbsf to the bottom of the cored section at 596.3 mbsf (~9–13 Ma; Fig. 2) and comprises interbedded bioturbated diatom-bearing silty clays and laminated silty clays (Fig. 2). Clasts > 2 mm in size occur throughout all lithostratigraphic units and are mostly dispersed in nature (1% in abundance; Escutia et al. Reference Escutia, Cooper, Eittreim, Tanahashi, Ishihara and O'Brien2011). A study from site U1358 suggested a stable Mertz Glacier system during the Late Pleistocene and sand size supply through ice rafting (Orejola et al. Reference Orejola and Passchier2014). The presence of extensive ice-rafted debris (IRD) in some portions of sites U1359 and U1361 implies an occasional weakening of the westwards-flowing Polar Current or the collapse of the continental shelf ice sheet due to warming (Cook et al. Reference Cook, Hemming, van de Flierdt, Davis, Williams and Galindo2017). Sites U1359 and U1361 are drilled close to each other, hence their sharing of various sedimentological characteristics, and they are suggested to have a multi-provenance sediment supply (Pant et al. Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013, Cook et al. Reference Cook, Hemming, van de Flierdt, Davis, Williams and Galindo2017, Pandey et al. Reference Pandey, Pant, Biswas, Shrivastava, Joshi and Nagi2018).
Fig. 2. A composite lithology of site U1359 (modified after Tauxe et al. Reference Tauxe, Stickley, Sugisaki, Bijl, Bohaty and Brinkhuis2012) along with age constraints based on magnetostratigraphy and biostratigraphy. Lithostratigraphic units IIc and III representing Miocene sediments of the studied cores are also mentioned. The lithostratigraphic units are based on the observed change in facies association. WL-U7 and WL-U6 are two seismic unconformities. Locations of samples selected for the present work are shown at different depths. FOT = first occurrence of Thalassiosira inura (Tauxe et al. Reference Tauxe, Stickley, Sugisaki, Bijl, Bohaty and Brinkhuis2012); GPTS = geomagnetic polarity timescale; mcd = measured compensated depth.
Geology of the inland sectors
The composite Terre Adélie Craton (TAC; Terre Adélie Land and George V Land) extends along the Antarctic coast between longitudes 136.4°E and 144.5°E within the Mawson craton and represents the easternmost area of the East Antarctic Shield (Fig. 1). The eastern boundary is marked by the prominent MSZ (144.3°E; Talarico & Kleinschmidt Reference Talarico and Kleinschmidt2003), which separates the Archaean and Proterozoic basement of the TAC from the Ross-Delamerian crystalline basement including granitoids and metasediments from the Cape Webb area (Di Vincenzo et al. Reference Di Vincenzo, Talarico and Kleinschmidt2007). The Ross-Delamerian rocks mark the western boundary of the Ross orogen that extends eastwards into Oates Land. The western boundary is not properly defined due to a lack of rock exposures and geophysical information. Rocher X (136.4°E) is the westernmost outcrop related to the TAC. Further west, outcrops in the Windmill Islands display Archaean and Proterozoic material strongly overprinted by late Mesoproterozoic events (between 1340 and 1130 Ma; Fitzsimons Reference Fitzsimons2003). The TAC exposure is restricted to hectometre to kilometre square islands and capes scattered along the coastline and to a few nunataks, mainly in the inland area of Commonwealth Bay. Farther east is Victoria Land, which is a high-grade metamorphic terrain except that the timing of peak metamorphism of terranes in the northern part has slightly younger U-Pb ages (460–500 Ma; e.g. Schüssler et al. Reference Schüssler, Bröcker, Henjes-Kunst and Will1999) than the southern portion (c. 480–550 Ma; e.g. Goodge Reference Goodge2007). The ages of the central Transantarctic Mountains are close to 480–545 Ma (Goodge Reference Goodge2007).
Methods
The examined sediments represent the ~200–600 measured compensated depth (mcd) interval corresponding to the ~6–13 Ma period (Fig. 2; Tauxe et al. Reference Tauxe, Stickley, Sugisaki, Bijl, Bohaty and Brinkhuis2012. The dominant grain size lies within the field of coarse silt to fine silt; hence, whole samples were taken for chemical treatment (Biswas et al. Reference Biswas, Pant, Pandey, Shrivastava, Panday, Ravichandran and Nair2020) and entire detritals > 53 μm were considered for heavy media separation. Grain sizes ~1 mm were studied as rock fragments separately, which were derived as IRD. The heavy mineral grains were separated via Bromoform and drained from the bottom of the flasks into funnels lined with WHATMAN #40 filter paper. The grains were thoroughly rinsed with distilled water and dried for 12 h. One extraction produced an average of ~0.3–0.5 g of heavy mineral grains, yielding ~100 grains on average. However, the population of heavy minerals varies throughout the depth. The entire separated portion of the heavy fraction was mounted on a glass slide. Approximately 60% of the total mounted grains were analysed. The mounted slides were analysed using energy-dispersive spectrometry (EDS) attached to a scanning electron microscope (SEM) and an electron probe micro-analyser (EPMA).
SEM-EDS analyses were carried out at the Department of Geology, University of Delhi, India, on a Zeiss SEM fitted with a Cambridge EDS and at the Department of Geology, Banaras Hindu University, Varanasi, India, on a Zeiss SEM fitted with a CAMECA EDAX EDS. The EPMA analyses were carried out at the Geological Survey of India, Faridabad, India, on a four-spectrometer CAMECA SX100 electron microprobe. The operating conditions for silicate analysis were 15 kV of high voltage and 10 nA beam current with a beam diameter of 1 μm or less. Natural mineral standards were used for calibrations. Calibrations for monazite were carried out at 20 kV and 20 nA (Pant et al. Reference Pant, Kundu, Joshi, Dey, Bhandari and Joshi2009). For monazite, Pb Mα was measured on an large pentaerythritol (LPET) crystal and the peak counting time was 240 s, with the background measured on both sides. For better statistics and counting, sub-counting was done four times with the background measured each time for Pb. For uranium, measured using an LPET crystal, the U Mβ line was used to avoid interference with the Th Mβ line, with a peak counting time of 120 s and a background time of 60 s on both sides of the peak. The thorium Mα peak was also counted for 240 s and the background for 120 s on a pentaerythritol (PET) crystal. Rare Earth elements measured and the X-ray lines used included La (Lα), Ce (Lα), Nd (Lα), Pr (Lβ), Sm (Lα), Ho (Lβ), Dy (Lα) and Gd (Lβ). Positions for background measurements were derived using the simulation software Virtual WDS (wavelength-dispersive spectrometry; Reed & Buckley Reference Reed, Buckley, Benoit, Bresse, Van't dack, Werner and Wernisch1996). Counting time at the peak varied from 40 to 60 s for these elements. The Y Lα line was used for yttrium, and it was counted for 240 s on peak and 120 s for background on both sides. The interferences of Y Lλ on Pb Mα and Th Mβ on Pb Mα were corrected after measuring the interfering lines during calibration and thereafter applying the overlap correction. The standards used for Pb, U and Th were crocite, U-glass and Th-glass. Synthetic silica-aluminium glass containing 4% rare Earth elements were used as standards for La, Ce, Nd, Pr, Sm, Ho, Dy and Gd. Phosphorus was calibrated with apatite. For Y, yttrium aluminium garnet was used. Al and Fe were calibrated using corundum and hematite, whereas, for silica content, orthoclase was used. Further details of the analytical procedure are given in Pant et al. (Reference Pant, Kundu, Joshi, Dey, Bhandari and Joshi2009).
Results
Mineral chemistry of the heavy minerals
Pyroxene dominates the total heavy mineral population followed by amphiboles, garnet, biotite magnetite and other phases. In the following sections, we present the mineral-wise description of this chemistry, which is significant for inferring the lithology of the provenance regions and subsequent igneous and metamorphic subglacial terrain interpretations. Based on the textures and chemical composition of the heavy minerals observed from the backscattered electron (BSE) image and microprobe point analysis, two affinities of igneous and metamorphic origin can be easily distinguished. Figure 3a shows the BSE images of major metamorphic minerals observed in the sediments of site U1359, such as pyroxene, ilmenite, almandine, pyrope, spessartine, biotite, amphibole, zircon, etc. Figure 3b shows the igneous affinity minerals, such as pigeonite, calcic plagioclase, apatite, titanite and pyroxenes. Overall, the populations of zircon and monazite were rare, as the generalized sub-ice Wilkes Land terrain is considered to be an ultramafic/mafic igneous terrain (FLIP).
Fig. 3. a. Selected back-scattered electron images of metamorphic minerals observed at various depths. Green dots represent the positions of electron probe micro-analyser point analyses of respective grains. A. Fused texture of two different sub-calcic amphiboles, indicating the metamorphic nature of these minerals. B. Diopside of metamorphic affinity. C. Smooth grain boundaries of zircon, indicating a metamorphic origin. D. Subhedral grains of pyrope and augite with smooth and equilibrated grain boundaries, indicating a metamorphic origin. E. Detrital biotite grain. F. Tourmaline and chlorite. b. Selected back-scattered electron images of igneous minerals observed at various depths. Green dots represent the locations of spot analyses in the respective grains. A. Anhedral grain of pigeonite and calcic plagioclase. B. Orthopyroxene of igneous origin associated with plagioclase. C. Detrital grains of apatite, titanite and orthopyroxene. D. Phenocryst of augite coexisting with plagioclase. E. Dissociated orthopyroxene grain; the outer texture of the grain indicates the effects of hydrothermal activity. F. Anhedral altered grain of pigeonite.
Pyroxene
The pyroxene classification of Morimoto et al. (Reference Morimoto1988) is used to differentiate pyroxene based on its chemical composition. All pyroxene analyses conform to the Ca-Mg-Fe or Quad chemical group (Morimoto et al. Reference Morimoto1988) as shown in Fig. 4a. Pyroxenes with Ca < 5% are considered orthopyroxene in this work as the grains are too fine to resolve optically as orthorhombic or monoclinic crystals.
Fig. 4. Classification diagrams of pyroxene a. after Morimoto et al. (Reference Morimoto1988), showing plots of detrital pyroxene of site U1359. The abundance of augite and pigeonite is clearly visible. Isotherms are plotted to explain the stability temperature field for the formation of pigeonites and augite ranging from 800°C to 1200°C (Lindsley & Andersen Reference Lindsley and Andersen1983). b. Plot of Al2O3 vs FeO + MgO for distinguishing between igneous and metamorphic orthopyroxenes from the heavy fraction of U1359 (fields after Rietmeijer et al. Reference Rietmeijer1983). Opx = orthopyroxene.
Orthopyroxene ranges from En46Fs53Wo01 at 200–203 mcd to En83Fs14Wo3 observed at 322–327 mcd. Mg-rich orthopyroxenes (enstatite) are present throughout the depth up to ~508 mcd (Table S1). No orthopyroxene grains were identified below this depth. Fe-rich orthopyroxene is reported from ~200 and ~280 mcd. Figure 4b shows that the orthopyroxenes present are of predominantly igneous affinities, but certain grains align towards the metamorphic field based on their FeO + MgO and Al2O3 contents (Rietmeijer et al. Reference Rietmeijer1983), corroborated texturally by the annealed nature of metamorphic grains (Fig. 3a, panel B). Orthopyroxene of igneous origin is shown in Fig. 3b (panel E), where it is associated with plagioclase. The grain is subhedral and shows slight weathering at the outer boundaries, indicating chemical alteration. The transition-zone orthopyroxene can be of igneous as well as metamorphic origin (Rietmeijer et al. Reference Rietmeijer1983).
Pigeonite, augite and diopside are major components and clinopyroxene is present throughout the investigated depth range. Augite dominates the heavy mineral assemblage followed by pigeonite and diopside. Pigeonite ranges in composition from En77Fs08Wo15 to En41Fs42Wo17 (Table S2). The MgO/MgO + FeO ratio varies from 0.34 to 0.92. Pigeonite is typically of igneous volcanic origin. It usually occurs as a phenocryst in mafic volcanic rocks. Since it is an unstable mineral at lower temperatures and it breaks down to orthopyroxene and augite, its presence suggests a shallow-level crystallization or volcanic activity associated with a relatively fast-ascending magma, resulting in the preservation of pigeonite phenocryst. The Fe/Mg ratio is controlled by the temperature of the magma: higher Mg values indicate a higher temperature (Robinson et al. Reference Robinson, Ross, Nord, Smyth and Jaffe1977). The clinopyroxene classification diagram also shows the contours of the isotherms. None of the clinopyroxenes fall in the forbidden zone and hence represent a population of stable clinopyroxenes. The isotherms suggest higher formation temperatures for the pigeonites and augites, ranging from 800°C to 1200°C (Lindsley & Andersen Reference Lindsley and Andersen1983). Compositionally, augite varies from En57Fs14Wo29 to En26Fs32Wo42. XMgO ranges from 1.00 to 0.39 (Table S3), which is a significant variance, indicating that they might have evolved from different sources (e.g. magmas). Augites of Ca affinity are typically of igneous origin (basalt or andesite).
There is no distinct way to differentiate compositions of metamorphic augite from igneous augite, except based on their texture. We observe both types of augites in our samples. Diopside is also present as a major phase (Table S4), suggesting the possibility of ultramafic igneous rocks such as kimberlites or peridotites and/or contact-metamorphosed skarns developed from high-silica dolomites in the provenance.
We characterize the type of mafic source of pyroxenes based on the geochemical criteria defined by Leterrier et al. (Reference Leterrier, Maury, Thonon, Girard and Marchal1982). Geochemically, the clinopyroxene derived from intermediate to acidic rocks has an Aliv content ranging from 0.03 to 0.05 atoms per formula unit on a six O basis (a.p.f.u.), while values of Aliv content higher than 0.05 are found to belong to mafic volcanic rocks (Rodriguez-Gonzalez et al. Reference Rodriguez-Gonzalez, Fernandez-Turiel, Perez-Torrado, Paris, Gimeno, Carracedo and Aulinas2012). Therefore, clinopyroxene with a Aliv content higher than 0.05 is considered for geochemical discriminant diagrams. The results show that most of the samples indicate a tholeiitic basalt and calc-alkaline basalt as the prime source rocks (Fig. 5a,b). A few grains show affinity towards calc-alkaline basalt as well (Fig. 5b).
Fig. 5. Discrimination diagrams of Leterrier et al. (Reference Leterrier, Maury, Thonon, Girard and Marchal1982) for clinopyroxenes present in the heavy media of sediments of site U1359: a. Ti vs Na + Ca (in atoms per formula unit); b. Ti vs Al (in atoms per formula unit).
Amphibole
The chemical composition of the analysed grains reveals the occurrence of chiefly four types of amphiboles: sub-calcic amphiboles, calcic amphiboles, sodic-calcic amphiboles and sodic amphiboles (classification based on Leake et al. Reference Leake, Woolley, Arps, Birch, Gilbert and Grice1997). Since lithium is not discernible from EPMA, magnesium-iron-manganese-lithium amphiboles (also termed sub-calcic amphiboles) were inferred by the very low value of calcium and sodium, calculated for site B, as compared to the other groups of amphiboles (Table S5).
Figure 6a shows the presence of actinolite, magnesio-hornblende, ferro-hornblende, tschermekite and ferro-tschermekite. In Fig. 6b, the grains are classified as edenite, pargasite, ferro-edenite, ferro-pargasite, kaersutite and ferro-kaersutite. Magnesio-hornblende, actinolite, edenite and pargasite are present as dominant phases, while ferro-hornblende, ferro-actinolite, tschermekite, ferro-tschermekite, ferro-edenite, ferro-pargasite, kaersutite and ferro-kaersutite are less abundant. Their representative analyses are given in Table S6. Actinolite and magnesio-hornblende with high Si a.p.f.u. values are related to metamorphic sources, while magnesio-hornblende with medium to low Si a.p.f.u. values and other calcic amphiboles such as ferro-hornblende and ferro-tschermakite belong to the igneous plutonic source (Rodriguez-Gonzalez et al. Reference Rodriguez-Gonzalez, Fernandez-Turiel, Perez-Torrado, Paris, Gimeno, Carracedo and Aulinas2012).
Fig. 6. Composition plots of calcic amphiboles from the sediments of site U1359 on the classification diagram devised by Leake et al. (Reference Leake, Woolley, Arps, Birch, Gilbert and Grice1997). a. (Na + K)A ≥ 0.5, showing the dominance of actinolite and magnesio-hornblende. b. (Na + K)A < 0.5, showing the dominance of edenite and pargasite. c. Composition of calcic amphiboles from the sediments of site U1359 on the Alvi vs Aliv discrimination graph of Leake (Reference Leake1965).
The Alvi and Aliv values (in a.p.f.u.) of calcic amphiboles are used to chemically distinguish igneous and metamorphic amphiboles (Fig. 6c). The grey field in Fig. 6c indicates a plutonic igneous source and the white portion outlined by a dotted line indicates the field of metamorphic calcic amphiboles (Leake Reference Leake1965, Rodriguez-Gonzalez et al. Reference Rodriguez-Gonzalez, Fernandez-Turiel, Perez-Torrado, Paris, Gimeno, Carracedo and Aulinas2012).
Detrital calcic amphiboles of site U1359 cover almost all of the fields with variable abundances. There is almost an equal amount of igneous as well as metamorphic sourcing of amphiboles. Amphiboles of metamorphic affinity predominantly belong to a high-pressure metamorphic belt, but a few samples also have an affinity to the low-pressure metamorphosed belt (Fig. 6c). The presence of sodic-calcic and sodic amphiboles indicates the possibility of the presence of blueschist facies rocks within a subglacial metamorphic terrain in the WSB region. However, this inference would require confirmation in the form of additional high-pressure minerals.
Garnet
The representative analyses of garnet are given in Table S7. Garnet is a characteristic mineral of metamorphic origin. However, it is also reported from acidic igneous rock (Deer et al. Reference Deer, Howie and Zussman1992). Chemically, garnet varies from Alm80Py15Sps3Grs2 to Alm45Py36Sps1Grs18. The classification of garnet is carried out on the trilinear plot of Fe-Mn-Mg-Ca formulated by Mange & Morton (Reference Mange and Morton2007) (Fig. 7a). The diagram classifies detrital garnet identified from sediments of site U1359 in five divisions and subdivisions based on their chemical characteristics. It also aids in describing the most probable source for the garnet.
Fig. 7. Classification of garnet, biotite and feldspar. a. Subdivision of garnet in the Fe + MFn-Mg-Ca ternary plot devised by Mange & Morton (Reference Mange and Morton2007), showing the presence of type A, B and C classes of garnet in the detrital sediments of site U1359. b. Classification of biotite based on the Fe/(Fe + Mg) vs Al content. Grey circles represent biotite data from the present study while black circles represent the biotite data published in Pant et al. (Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013). c. Ternary phase diagram of the feldspars obtained from Greenwood & Earnshaw (Reference Greenwood and Earnshaw1998). Diamonds represent the feldspars present in the sediments of site U1359, which are associated with the pyroxenes and amphiboles.
Type A garnets typically have high Mg and low Ca. These garnets are generally considered to be derived from high-grade granulite facies metasedimentary sequences (Sabeen et al. Reference Sabeen, Ramanujam and Morton2002, Morton et al. Reference Morton, Hallsworth and Chalton2004), but they can also be derived from intermediate silicic igneous rock (Green & Ringwood Reference Green and Ringwood1968, Fitton Reference Fitton1972, Hamer & Moyes Reference Hamer and Moyes1982). In the proximal Wilkes Land terrain, the most probable source for type A garnet is granulite facies metasedimentary rocks, as we also found metamorphic orthopyroxene, which is typical of granulite facies metamorphism.
Type B garnet has low Mg and variable Ca. These garnets are derived from amphibolite facies metasediments. Two subgroups are distinguished based on the XCa%: Bi, which has XCa < 10%, and Bii, which has XCa > 10%. The population within the Bi field is from either intermediate acidic rocks or low-grade metasedimentary rocks.
Type C garnets are rich in Mg and Fe. Two subclasses are distinguished here based on the XMg content: Ci has XMg < 40% and Cii has XMg > 40%. Almost 50% of the garnet grains fall in the field of Ci and are interpreted to be derived from a metabasic sequence.
Biotite
Representative biotite analyses are listed in Table S8, and the structural formula is calculated based on 11 oxygen atoms. High Ti biotite was reported previously from the Plio-Pleistocene sediments of the same core (Pant et al. Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013); however, Miocene sediments show a predominance of low to moderate Ti and Fe contents (Fig. 7b) and have higher MgO (TiO2 < 1.5%; MgO ~11–13%) in biotite. The chemical composition of biotite without associated assemblages is insignificant for distinguishing the mode of genesis. However, a moderate to low Ti content is suggestive of medium-grade metamorphic and/or acidic-intermediate magmatic sourcing.
Feldspars
The feldspars found in the heavy mineral population of the sediments of site U1359 are associated with pyroxenes and amphiboles. Their structural formula is calculated based on eight oxygen atom and the representative analyses are given in Table S9. A classification of feldspar was also carried out (Fig. 7c). Plagioclases are common feldspar grains in the analysed fractions.
There are two chemical variants present, with one being high in calcium (An > 50) and the other of low-intermediate composition (An < 50; Fig. 7c); both show considerable variation. High-calcium plagioclase also belongs to two categories, with one representing a narrow range (An50–70) and the other of bytownite (An70–90) composition. The low-intermediate calcium plagioclase (An0–50) represents at least two metamorphic sources, one rich in albite (XAb > 90) and the other with a range of values from An20 to An40. The narrow range of anorthite content of high Ca represents a basaltic magmatic source. The ternary plot of the feldspar classification shows the presence of K-feldspar in the sediments. Its presence is common in felsic volcanic and plutonic rocks. However, it has also been reported from ultrapotassic mafic, high-temperature contact metamorphic rocks and hydrothermally altered rocks.
Ilmenite
The representative ilmenite analyses are reported in Table S10. Chemically, Basu & Molinaroli (Reference Basu and Molinaroli1991) found that ilmenite grains with TiO2 contents between 50 and 60 wt.% are more prevalent in metamorphic rocks, but they also occur in igneous rocks ranging between 40 and 50 wt.%.
Schneiderman (Reference Schneiderman1995) noted that ilmenite from metamorphic sources is richer in TiO2 than ilmenite in igneous rocks. Here, we found both types of ilmenites, with a predominance of ilmenite with TiO2 ranging from 40% to 50%, which is taken here as being indicative of an igneous origin; however, 17 grains out of 40 show values that suggest derivation from metamorphic rocks instead.
Epidote, sphene, scapolite, cordierite, Ba-silicates, magnetites, apatite, rutile, Fe-hydroxides, sulphides, titanomagnetite, zircon, etc., are the other heavy minerals that were observed in trace amounts. A wide range of chemically distinguishable mineral species indicate the presence of rather a complex terrain comprising a low-grade as well as a high-grade metamorphic provenance and a mafic igneous source.
Mineral chemistry of the rock fragments
Detrital rock fragments provide the most valuable tool for investigating potential source rocks. Rock fragments are the detritals derived from certain lithologies as IRD and consist of more than one texturally constrained coexisting phase (Fig. 8a, left inset). These facilitate the interpretation of the genesis of the rocks and the tectonic setting of the provenance region as they preserve coexisting phases and textural features. In this study, we focused on three different rock types identified based on the texture and mineral composition of the associated phases: basalt fragment, garnet biotite schist and two-mica schist (Fig. 8). One representative rock fragment is of igneous origin, while the other two rock fragments are metamorphic in origin and are used here to further constrain the metamorphic conditions of the source rocks
Fig. 8. Various selected rock fragments. a. Representative of rock fragments collected from various depths of site U1359 in the left inlet and back-scattered electron image of the igneous-origin rock fragment 1 (RF-1) with clinopyroxene (cpx) and anorthite (an) identified as major phases and sphene present as an accessory phase. b. Back-scattered electron image of RF-2: muscovite-biotite schist (green dot represents the location of spot analyses carried out by electron probe micro-analyser). The presence of schistosity confirms its metamorphic origin. c. Back-scattered electron image of RF-3: garnet biotite schist represented by a porphyroblast of garnet with inclusions of biotite and quartz.
Basalt/dolerite (rock fragment 1)
Rock fragment 1 (RF-1) is retrieved from 218.9 mcd (Fig. 8a), and the presence of a hypidiomorphic equigranular texture denotes it to be igneous in origin. Phenocrysts of clinopyroxene and calcic plagioclase are present as the major constituents and show a subophitic intergranular texture (Fig. 8a). EDS analyses of RF-1 classify clinopyroxene as augite (CaO ~25%, FeO and MgO ~12–13%), while plagioclase is anorthitic (CaO ~22%). Sphene is identified as an accessory phase. The analyses reaffirm the presence of igneous-origin augite in these sediments.
Two-mica schist (rock fragment 2)
This rock fragment is recovered from the sediment depths of 230–233 mcd and represents a fine-grained schistose metamorphic rock with an assemblage of muscovite + biotite + quartz + ilmenite + monazite. Coexisting muscovite and biotite define the fine-grained schistosity (Fig. 8b). Compositionally, muscovite is present as the dominant phase, indicating the low-grade metamorphism of the pelitic source.
Representative analyses of the two-mica schist are given in Table S11. The a.p.f.u. values in the table are based on 11 oxygen atoms for biotite and muscovite and 5 oxygen atoms for ilmenite. The XMg/(XFe + XMg) values of biotite are 0.51–0.57, which signifies enrichment of Mg in biotite, and the Ti value ranges from 0.17 to 0.35 a.p.f.u., indicating a low-grade metamorphic terrain of the rock fragment. Ilmenite has Ti ~1.8 a.p.f.u and Fe ~1.2 a.p.f.u. Monazite is found to be associated with the two-mica schist (shown in Fig. 8b), and this is discussed in detail separately.
Garnet biotite schist (rock fragment 3)
Porphyroblast of subhedral garnet associated with biotite and quartz is the third rock fragment we investigated to deduce metamorphic conditions. This rock fragment is recovered from the sediments in the depth range ~227–230 mcd.
Based on the observed features and mineral composition, the possible mineral assemblage for the garnet biotite schist is garnet + biotite + quartz. The typical metapelitic assemblage and the presence of internal schistosity in the porphyroblast of garnet attest to its metamorphic origin (Fig. 8c). The internal schistosity is defined by biotite and quartz. The size of the porphyroblast is ~400 μm. Garnet is subhedral with a smooth grain boundary, indicating the absence of any significant alteration. Garnet may have developed from the consumption of biotite and quartz during a possible second stage of metamorphism.
The representative mineral chemistry of garnet and biotite is given in Table S11. Garnet is compositionally homogeneous and almandine-spessartine rich with XFe ~0.57, XMg ~0.14, XMn ~0.23 and XCa ~0.04. The XMg of biotite is ~0.55 and Ti is 0.38 a.p.f.u., which are comparable to the values of biotite obtained from the two-mica schist. This indicates a similar genesis of the biotite of both rock types. The ~23% spessartine content also suggests that the garnet is of low-grade metamorphic origin.
Geothermometry
The mineral assemblages and mineral compositions of any type of rock are sensitive to the pressure and temperature conditions of formation if events subsequent to mineral equilibration have not significantly modified the original composition of these minerals. Hence, minerals in equilibrium can be calibrated for the estimation of the pressure and temperature conditions at which a geological material experienced a major event such as metamorphism. Considering Wilkes Land as the dominant provenance for the detritals of site U1359, there is a possibility of the associated minerals of rock fragments being represented in thermodynamic equilibrium and hence being used to decipher the limiting metamorphic conditions existing in the provenance. For garnet-biotite schist rock fragments, the garnet-biotite formulations of Ferry & Spear (Reference Ferry and Spear1978), Dasgupta et al. (Reference Dasgupta, Sengupta, Guha and Fukuoka1991) and Bhattacharya et al. (Reference Bhattacharya, Mohanty, Maji, Sen and Raith1992) are used. The estimated temperature is ~587 ± 17°C at 7 kbar pressure, which indicates amphibolite facies metamorphic conditions.
Monazite geochronology
Site U1359 is the easternmost site drilled during IODP Expedition 318 on the continental rise. The chances of getting any geochronological signatures are reduced significantly due to its remoteness from the provenance. Moreover, the sub-ice terrain in Wilkes Land is dominantly an igneous province (Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013), which further diminishes the chances of getting any radioactive minerals as detritals. Despite these limitations, a monazite grain was found to be associated with the muscovite-biotite schist rock fragment. The grain is subhedral with a maximum dimension of ~40 μm. The monazite is Ce rich with values ranging from 0.406 to 0.441 a.p.f.u. Th varies from 0.050 to 0.089 a.p.f.u. (all a.p.f.u. values are on a four oxygen atom basis), while P ranges from 0.883 to 0.937 a.p.f.u. The Ca value ranges from 0.055 to 0.080 a.p.f.u., indicating the presence of solid solutions of brabantite and huttonite (Table S13). For individual spot ages, the formulation of Montel et al. (Reference Montel, Foret, Veschambre, Nicollet and Provost1996) was followed, whereas the age probability plots and unmixing of ages were obtained using the software Isoplot 3.00 (v.3.71.09.06.19nx; Ludwig Reference Ludwig2003). Uncertainties in individual analyses in the data table and weighted mean ages are quoted at the 95% confidence level (2σ). The -distribution of age points is shown in Fig. 9a. The monazite provides a well-constrained age of 799 ± 13 Ma (Fig. 9b). The weighted average indicates 799 ± 25 Ma (Fig. 9c), whereas the chemical Th-U-total Pb isochron method (CHIME) age extracted from the monazite grain indicates an isochron age of 810 ± 20 Ma (Fig. 9d), confirming an event at ~800 Ma.
Fig. 9. a. Back-scattered electron image of monazite present in rock fragment 2 extracted from heavy mineral fractions at of depth 230–233 measured compensated depth marked with ages. b, Frequency diagram and weighted average of the age obtained from monazite. c. Weighted average indicates 799 ± 25 Ma. d. CHIME age extracted from the monazite grain, indicating an isochron age of 810 ± 20 Ma. CHIME = chemical Th-U-total Pb isochron method; MSWD = mean square of weighted deviation.
Discussion
Igneous and metamorphic origins of detritals are inferred based on the textures and chemical composition of the heavy mineral assemblage. Our classification of pyroxene shows the presence of orthopyroxene, pigeonite, augite and diopside. Further classification enables us to differentiate between metamorphic orthopyroxene and igneous orthopyroxene. Considering that clinopyroxene with Aliv > 1.5 belongs to igneous sources, the analysis was used to infer the geochemical behaviour of the source basalt. The results show that most of the samples indicate a tholeiitic to calc-alkaline basalt as a prime source rock (Fig. 5a). Four different species of amphiboles present in the sediments are sub-calcic amphiboles or calcic amphiboles (Fig. 6a,b). The metamorphic and igneous calcic amphiboles are interpreted based on the discrimination graph in Fig. 6c. The calcic amphiboles and biotite of metamorphic affinity show characteristics of both high-grade as well as low-grade rocks, indicating the presence of both types of metamorphic belt in the sediment provenance region (Figs 6c & 7b). The classification of garnet defined by Mange & Morton (Reference Mange and Morton2007) shows the presence of garnet belonging to granulite facies as well as amphibolite facies of metasediments (Fig. 7a). Minor sourcing from metabasic and ultramafic rocks is also indicated from the classification (Fig. 7a). The compositional variation of garnet corroborates our interpretation of two contrasting grades of metamorphism in the provenance region.
Two-mica schist (230–233 mcd) and garnet-biotite schist (227–230 mcd) are two different rock fragments that were studied to deduce the metamorphic conditions. Both of the rock fragments are metamorphic in origin, with the mineral assemblage for RF-2 (Fig. 8b) given as muscovite + biotite + quartz + ilmenite + monazite and that for RF-3 given as garnet + biotite + quartz (Fig. 8c). The presence of two-mica schist indicates low- to medium-grade metamorphic conditions. Temperature and pressure deduced as limited conditions agree with the inferred granulite facies and greenschist to amphibolite facies of metamorphism, respectively. Signatures of greenschist to a granulite grade of metamorphism are evident from the mineral chemistry as well as pressure-temperature estimations, indicating the presence of a mobile belt in the source domains. Chemical dating of monazite provides a well-constrained age of 799 ± 13 Ma (Fig. 9) from the low-grade assemblage, evincing the presence of a low-grade metamorphic event during Neoproterozoic (Tonian) times.
The geology of the Wilkes Land sector is predominantly inferred through 1) unconstrained projections of its Australian counterpart (Boger Reference Boger2011), 2) geophysical signatures suggesting the Proterozoic basement provinces accreted during supercontinental assemblies (e.g. Golynsky et al. Reference Golynsky, Ferraccioli, Hong, Golynsky, von Frese and Young2018) and 3) geochronological dating of detrital zircons and monazites recovered from the offshore sediments and exposed sedimentary rocks (e.g. Pierce et al. Reference Pierce, van de Flierdt, Williams, Hemming, Cook and Passchier2017, Daczko et al. Reference Daczko, Halpin, Fitzsimons and Whittaker2018).
The WSB is almost 600 km wide along the Coats Land coast and is > 1400 km long (Bedmap2; Fretwell et al. Reference Fretwell, Pritchard, Vaughan, Bamber, Barrand and Bell2013). It includes a variety of sub-basins identified from radar data and defined as the Eastern Basin, Central Basin and Western Basin (Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009) separated by highlands (Frederick et al. Reference Frederick, Young, Blankenship, Richter, Kempf, Ferraccioli and Siegert2016; shown in Fig. 10). The main source of sediment supply from the WSB is predominantly along Mertz Glacier through ice streams, which is the close to the drill site (Rignot et al. Reference Rignot, Mouginot and Scheuchl2011, Pant et al. Reference Pant, Biswas, Shrivastava, Bhattacharya, Verma and Pandey2013). Other more distal ice streams capable of delivering sediments to the IODP site include Ninnis Glacier, which follows the Western Basin, and the underlying Central Basin delivers its sediment to the western and eastern Cook ice streams (Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009, Frederick et al. Reference Frederick, Young, Blankenship, Richter, Kempf, Ferraccioli and Siegert2016). However, the present-day configuration of the EAIS is unlikely to have remained unchanged during the 6–13 Ma period, at least according to sheet and climate modelling predictions (e.g. DeConto & Pollard Reference DeConto and Pollard2016) and previous marine sediment interpretations (Verma et al. Reference Verma, Bhattacharya, Biswas, Shrivastava, Pandey and Pant2014). In end-member deglaciated scenarios, for example, subglacial highlands will become a more probable sediment source region, while in a glaciated setting the subglacial basins will be a more probable source region due to enhanced erosion along the fast-flowing ice streams (e.g. Pandey et al. Reference Pandey, Pant, Biswas, Shrivastava, Joshi and Nagi2018).
Fig. 10. Reconstructed map of geology and linkage of Australia and East Antarctica. a. The base map of Australia is from Gerner et al. (Reference Gerner, Holgate and Budd2010) and that of East Antarctica is from An et al. (Reference An, Wiens, Zhao, Feng, Nyblade and Kanao2015), made using Quantarctica (Matsuoka et al. Reference Matsuoka, Skoglund, Roth, de Pomereu, Griffiths and Headland2021). The Centralian Superbasin and its constituent basins in South Australia are inferred from Walter et al. (Reference Walter, Veevers, Calver and Grey1995). Geologies of Adélie Land and North Victoria Land are as reported by Ménot et al. (Reference Ménot, Duclaux, Peucat, Rolland, Guillot and Fanning2007) and Weaver et al. (Reference Weaver, Bradshaw and Laird1984), respectively. Locations of subglacial Eastern Basin (EB), Central Basin (CB) and Western Basin (WB) are from Ferraccioli et al. (Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009). The probable area of extension of the Gairdner Dyke Swarm is shown by blue dashed lines in the interior of the Wilkes Subglacial Basin. b. Known geology of the study area overlaid on the ADMAP-2 aeromagnetic anomaly map (modified from Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013, Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014, Golynsky et al. Reference Golynsky, Ferraccioli, Hong, Golynsky, von Frese and Young2018). MSZ = Mertz Shear Zone.
The dominance of clinopyroxenes (especially pigeonite) and igneous amphiboles we identified in U1359 heavy minerals supports the presence of mafic rocks in the hinterland of the George V Coast region inferred as FLIP by Cook et al. (Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013; Fig. 4). However, based on aeromagnetic interpretations, FLIP is compared to the Jurassic Ferrar Group, which is relatively well-exposed mostly as sills and rare dykes intruding into the Beacon Supergroup throughout the Transantarctic Mountains (e.g. Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013). We infer an extension of mafic rocks that are well-known in the formerly adjacent Gawler Craton of South Australia, at least within the northern WSB. In addition, diverse geochemical characteristics of clinopyroxene (Fig. 6) suggest multiple events of volcanic eruptions in the WSB. Hence, the FLIP evolution seems to be much more complex than previously inferred (e.g. Cook et al. Reference Cook, van de Flierdt, Williams, Hemming, Iwai and Kobayashi2013).
The Jurassic Ferrar Group is relatively well-exposed mostly as sills and rare dykes throughout the Transantarctic Mountains (e.g. Roland & Wöerner Reference Roland and Woerner1996). It intrudes into the Beacon Supergroup rocks of Devonian to Jurassic age (Barrett et al. Reference Barrett, Elliot and Lindsay1986) and is in places overlain by thick Jurassic Kirkpatrick Basalt (Elliot et al. Reference Elliot, Fleming, Haban and Siders1995).
Considering the MSZ-KSZ correlation, Mawson Craton in East Antarctica exhibits clear affinities with South Australia (e.g. Pierce et al. Reference Pierce, van de Flierdt, Williams, Hemming, Cook and Passchier2017). Coastal exposures of Terre Adélie Craton (e.g. Duclaux et al. Reference Duclaux, Rolland, Ruffet, Ménot, Guillot and Peucat2008 and references therein) preserve key geological evidence for several orogenic and magmatic events that resemble the events recognized in Gawler Craton (Fitzsimons Reference Fitzsimons2003, Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014). The WSB may also contain Neoproterozoic red beds, potentially correlative to those identified in the Adelaide Rift Complex in Australia (Schmidt & Williams Reference Schmidt and Williams2013). Mafic rift-related Neoproterozoic rocks also explain the distinct linear north-west to south-east-orientated aeromagnetic anomalies identified along the coastal regions of George V Land (Damaske et al. Reference Damaske, Ferraccioli and Bozzo2003). These anomalies lie at an angle to the more north to south-orientated Ferrar-related magnetic anomalies seen further south (Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009). Although these magnetic anomalies have previously been interpreted as Jurassic rift-related tholeiites along the coast of George V Land (Finn et al. Reference Finn, Goodge, Damaske, Fanning, Fütterer, Damaske, Kleinschmidt, Miller and Tessensohn2006), we contend here that an alternative interpretation that these arise from older Neoproterozoic mafic sources is more probable. High-resolution aeromagnetic data over Gawler Craton and also the Coompana block in Australia (e.g. Wise et al. Reference Wise, Dutch, Pawley, Foss and Thiel2018) image linear north-west to south-east- to west-north-west to south-south-east-orientated magnetic anomalies over the 830–750 Ma Gairdner-Amata Dyke Swarm and inferred correlative rocks. Notably, these anomalies lie on strike, with the anomalies detected from much lower-resolution aeromagnetic data along the George V Coast (Damaske et al. Reference Damaske, Ferraccioli and Bozzo2003). Based on this correlation to the Australian counterpart, the possibility of extension of the Gairdner-Amata Dyke Swarm in the interior of the WSB is undeniable. The detrital pigeonite plot suggests the more complex variable composition of pigeonite if we compare it to the Ferrar dolerites, suggesting more than one igneous volcanic event recorded in the interior of the WSB.
Crustal sagging at c. 800 Ma initiated the Centralian Superbasin followed by the deposition of marine and fluvial sands (Walter et al. Reference Walter, Veevers, Calver and Grey1995). A similar event would have caused the presence of low-grade metamorphic rocks in the interior of the WSB. The Tonian-aged monazite present in low-grade schist at U1359 of c. 830–750 Ma might represent the basement age for an older mafic volcanic activity that we propose to be the continuation of the Gairdner-Amata Dyke Swarm and the formation of the Adelaide Rift Complex in formerly adjacent Australia (Zhao et al. Reference Zhao, McCulloch and Korsch1994). This hypothesis is further supported by the clinopyroxenes of U1359 that suggest sourcing from a putative tholeiitic magmatic source, which would be similar to that of the Gairdner giant dyke swarm (Zhao et al. Reference Zhao, McCulloch and Korsch1994). The WSB region was not only affected by previously proposed early Cambrian basin formation and Jurassic tholeiitic magmatism (Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009), but also by older Neoproterozoic crustal extension and magmatism, which in turn heralded the break-up of the Rodinia supercontinent (e.g. Li et al. Reference Li, Bogdanova, Collins, Davidson, De Waele and Ernst2008). We suggest that the mafic rocks exposed in the WSB are much older than Jurassic, and their association with the Australian Gairdner-Amata Dyke Swarm is more probable than the earlier proposed association with the Jurassic Ferrar Group of the Transantarctic Mountains. This is further supported by the presence of low-geothermal-gradient signatures reported from the area (Fig. 10a; An et al. Reference An, Wiens, Zhao, Feng, Nyblade and Kanao2015). Younger magmatic rocks exposed in the Beacon Supergroup and the Ferrar Group in the Transantarctic Mountains exhibit high geothermal signatures (Fig. 10a, inset image).
Figure 10b shows the geology superimposed on the ADMAP-2 aeromagnetic data of the terrane, exhibiting high magnetic anomalies in the WSB (polygons with dashed black outline in Fig. 10b). Our new data suggest that the WSB includes low-grade Neoproterozoic metasedimentary basement rocks (c. 800 Ma) that are significantly older compared to the aeromagnetically inferred early Cambrian rocks, which have in turn been interpreted as reflecting a broad back-arc basin region affected by the Ross Orogen (Fig. 10b; Ferraccioli et al. Reference Ferraccioli, Armadillo, Jordan, Bozzo and Corr2009). This corroborates the orthopyroxene-garnet thermobarometer thermal conditions of 869 ± 24°C at 8 kbar and, using the garnet and biotite thermometer, of ~587 ± 17°C at 7 kbar. This supports a largely unexposed and at least low- to medium-grade metamorphosed Neoproterozoic terrane in the WSB region. The inferred low-grade metasediments could potentially resemble exposed low-grade metasedimentary rocks assigned to the western Wilson Terrane in Oates Land (Tessensohn & Henjes-Kunst Reference Tessensohn and Henjes-Kunst2005); however, the inferred c. 550–530 Ma depositional ages of these rocks are too young to be related to the c. 800 Ma metasedimentary rocks inferred in this study. The unexposed metasediments in the WSB region are also too old to represent a continuation of the c. 550–520 Ma Kanmantoo Trough that was subsequently inverted during the Delamerian Orogen (Aitken et al. Reference Aitken, Young, Ferraccioli, Betts, Greenbaum and Richter2014, Reference Aitken, Betts, Young, Blankenship, Roberts and Siegert2016). Whether these Precambrian metamorphic rocks are confined to the linear mountain ranges that flank the western WSB or also partially underlie the glacially over-deepened Western Basin and Central Basin (Fig. 10; Frederick et al. Reference Frederick, Young, Blankenship, Richter, Kempf, Ferraccioli and Siegert2016), especially where Ninnis Glacier presently flows, is less clear, however.
The presence of 800–750 Ma signatures is further supported by detrital zircon geochronology in the sediments investigated from nearby Mertz Shelf-slope erratic detritals (e.g. Goodge & Fanning Reference Goodge and Fanning2010, Veevers & Saeed Reference Veevers and Saeed2011). Ages of 800 Ma are also reported from the Section Peak Formation and Ross Supergroup of adjacent North Victoria Land, which were clustered with either early Proterozoic ages c. 1200 Ma or with the Pan-African c. 500 Ma event (e.g. Paulsen et al. Reference Paulsen, Deering, Sliwinski, Bachmann and Guillong2016). From detrital zircon records, c. 800 Ma ages ascribed usually either to Pb loss or not being part of the basin (e.g. Daczko et al. Reference Daczko, Halpin, Fitzsimons and Whittaker2018, Tooze et al. Reference Tooze, Halpin, Noble, Chase, O'Brien and Armand2020) merit more consideration. The limitations of the chemical geochronology of monazite smooth grain boundaries and the lack of cracks suggest the metamorphic origin of the monazite and not the detrital origin in the protolith, as well as indicating that the monazite has not experienced any mobility of REE phases and hence will provide accurate ages. The present work suggests the possibility of a 1100–500 Ma timeframe to preserve a complex history of crustal evolution, registering multiple thermo-tectonic events in the interior of the WSB. Globally, this is manifested in secular trends shown by a peak in zircon frequency at c. 700–800 Ma (Puetz et al. Reference Puetz, Ganade, Zimmermann and Borchardt2018), the collision ages of greenstone belts (Condie Reference Condie1994) and the age distribution of eclogites and ultra-high-temperature metamorphic zones (Brown Reference Brown2007). This is further supported by the newly emerged evidence of Tonian-Cryogenian (c. 700–800 Ma) metamorphic/magmatic events from Neoproterozoic shields worldwide (e.g. Africa: Itano et al. Reference Itano, Iizuka, Chang, Kimura and Maruyama2016, Sommer et al. Reference Sommer, Kröner and Lowry2017; China: Ge et al. Reference Ge, Shen, Selby, Deng and Mei2016, Haozheng et al. Reference Haozheng, Oscar, Huafeng, Hong, Xi and Mingguo2020; East Antarctica: Jacobs et al. Reference Jacobs, Elburg, Läufer, Kleinhanns, Henjes-Kunst and Estrada2015, Arora et al. Reference Arora, Pant, Pandey, Chattopadhyay, Greenbaum and Siegert2020, Sadiq et al. Reference Sadiq, Dharwadkar, Roy, Arora, Shah and Bhandari2021 Gupta et al. Reference Gupta, Pandey, Arora, Pant and Rao2022, and concordant zircon ages of Mikhalsky et al. Reference Mikhalsky, Andronikov, Leitchenkov and Belyatsky2020, Zong et al. Reference Zong, Ren and Wu2020; India: Bose et al. Reference Bose, Das, Torimoto, Arima and Dunkley2016, Arora et al. Reference Arora, Pant, Sharma and Sadiq2017; southern Kazakhstan: Tretyakov et al. Reference Tretyakov, Pilitsyna, Degtyarev, Salnikova, Kovach and Lee2019). Although these ages are reported from almost every Gondwanan component, the events corresponding to c. 700–800 Ma (Tonian-Cryogenian) essentially remain enigmatic in the global orogenic context. Constraining the evolutionary history of Wilkes Land during the Neoproterozoic era is beyond the scope of our study.
Conclusions
1) Our sediment provenance study reveals previously unrecognized Precambrian metamorphic subglacial basement terrains. An upper amphibolite to granulite facies metamorphic belt is inferred to exist inland of correlative rock exposures in the Terre Adélie Craton, and it is linked based on the similarity in metamorphic grade to the Archaean and Palaeoproterozoic metamorphic rocks that are well-known in the formerly adjacent Gawler Craton in South Australia.
2) A unimodal age of 799 ± 13 Ma (Tonian) was estimated from a monazite grain found in the low-grade two-mica schist rock fragment, indicating a less recognized event in this sector.
3) The possibility of the presence of a Neoproterozoic metasedimentary basin and a younger magmatic terrane defines a complex geological set-up of the WSB.
Acknowledgements
This work is a result of analyses of sediments of the drill core U1359 provided by the IODP Expedition 318. The National Centre for Polar and Ocean Research funded this project. Dr Prakash Srivastava, Geological Survey of India, is acknowledged for providing sediment samples of site U1359 of IODP Expedition 318. We would also like to thank the Geological Survey of India, Faridabad, for providing access to the EPMA facility. N.C. Mehra, Department of Geology, University of Delhi, is acknowledged for his support during the SEM-EDS analysis. We thank Professor N.V. Chalapathi Rao for allowing access to the SEM-EDS facility at Banaras Hindu University, India. The work acknowledges the use of data produced during the International Collaboration for Exploration of the Cryosphere through Aerogeophysical Profiling (ICECAP) project. MP and DA acknowledge University Grants Commission (UGC) India for research fellowships during their doctorates. MP acknowledges the IOE seed grant of Banaras Hindu University, Varanasi, India. We thank Professor Sandra Passchier and an anonymous reviewer for their insightful comments. We also acknowledge Dr Thamban Meloth for editorial handling.
Financial support
The project is funded by National Centre for Polar and Ocean Research (NCPOR), India.
Competing interests
The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.
Availability of data and material (data transparency)
The data are included in this publication in the form of Supplemental Tables S1–S13.
Author contributions
MP: execution, interpretation and writing - original draft preparation; NCP: conception, design and writing - review and editing; DA: interpretation and writing - review and editing; FF: interpretation and writing - review and editing; RG: writing, figures and editing; SJ: data analysis.
Supplemental material
Thirteen supplemental tables will be found at https://doi.org/10.1017/S0954102023000123.
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