Carbonate Deposition Along the Paleo-pacific Passive Margin of East Antarctica

Published: 2021-06-29 13:35:05
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Extensive carbonate deposition along the paleo-Pacific passive margin of East Antarctica began as early as 525 Ma, during what is termed the Platform stage (Goodge, 1997; Goodge et al., 2004b). A period of moderate basin inversion followed carbonate sedimentation and occurred along the rifted margin, subaerially exposing portions of the platform (Withjack et al., 1995; Goodge, 1997). Basin inversion is generally linked to compression regimes, often associated with convergence (subduction) of an adjacent plate (Withjack et al., 1995). In the case of the Ross Orogen, basin inversion was related to spreading in the recently formed paleo-Pacific Ocean (Goodge et al., 1997). As the paleo-Pacific plate continued to spread normal to the rifted margin, resistance to subduction by the east Antarctic craton triggered a reactivation of normal faults associated with the rifting of Rodinia (Withjack et al., 1995; Goodge, 1997).
The reactivation of basement-propagating faults led to an inversion of the syn-rift stratigraphy associated with the fragmentation of Rodinia (Goodge, 1997). As compressional stresses continued to focus in the cratonic-oceanic crust transition of the continental margin, subduction initiated, and the early stages of arc magmatism commenced (Fig. 4) (Dewey, 1988; Bott, 1992; Goodge et al., 1997). Subduction was at an oblique angle with respect to the continental margin, likely due to a shift in plate-kinematics related to the amalgamation of the Gondwanan supercontinent (Goodge et al., 1997). Despite the onset of subduction and subsequent deformation on the previously quiescent continental margin, carbonate sedimentation continued (Goodge et al., 1997). Carbonate sedimentation was restricted to preserved sections of the carbonate platform (Goodge, 1997) and is recorded in the adjacent stratigraphy as the Shackleton Limestone of the Byrd Group, discussed below.Byrd Group
The Byrd group is comprised of the lower Shackleton Limestone and the overlying Starshot and Douglas Formations (Laird, 1963; Laird et al., 1971). The Byrd Group is a thick succession of carbonate facies that transition upsection into syn-tectonic siliciclastic successions (Laird et al., 1971). The Shackleton Limestone. The Shackleton Limestone is a 1000 – 2000 m thick lower Cambrian succession containing a basal, unfossiliferous quartz arenite member that underlies meter-scale siliciclastic-carbonate cycles (Laird et al., 1971; Burgess and Lammerink, 1979; Myrow et al., 2002b). Rees et al. (1989) constrained the depositional setting for the Shackleton Formation to a simple carbonate ramp with a localized facies change into higher-energy oolitic shoal deposits. The uppermost member of the Shackleton Limestone contain incised Archaeocyathan bioherms (Myrow et al., 2002b). The presence of Archaeocyathan bioherms restricts the depositional age of the uppermost Shackleton to Middle Cambrian time (Myrow et al., 2002b).
The basal member of the Shackleton Limestone is a weathered quartz arenite, with interbedded dolomitic grainstone (Myrow et al., 2002b; Goodge et al., 2004b). All upper contacts between sandstone and carbonate are gradational in nature containing calcareous sandstones and silica-rich carbonates (Laird et al., 1971; Myrow et al., 2002b). The relative abundance of carbonate increases upsection, where a siliciclastic component is absent (Laird et al., 1971; Myrow et al., 2002b). Sedimentary structures within the basal quartz arenite include bi-directional ripple marks and clay drapes, indicative of shallow water, tidal-influenced deposition (Laird et al., 1971; Myrow et al., 2002b). An overall upward-deepening sequence is supported by the upsection decrease in clastic components corresponding to a marine transgression that led to the establishment of a carbonate ramp along a quasi-stable cratonic margin (Myrow et al., 2002b). Brasier et al. (1994) indicated that this transgression was documented in many lower Cambrian rocks and likely indicates a eustatic event, corresponding to carbon isotope excursion IV (Myrow et al., 2002b). This global transgressive event likely corresponds to global changes in climate and sediment accommodation associated with the amalgamation of Gondwana (e.g. Hoffman, 1991; Knoll and Walter, 1992; Braiser, 1992; Goodge et al., 2004b).
Detrital zircon U-Pb age populations from the basal Shackleton Limestone reflect a mix of cratonic sources. Sample SLB (Fig. 4) contains lesser contributions from early Paleoproterozoic and Archean sources and prominent peaks at 970 Ma, 1135 Ma, 1515 Ma, 1750 – 1935 Ma, and 2500 Ma peaks, indicating primarily Grenville and Mesoproterozoic sources (Goodge et al., 2004b). Based on U-Pb detrital zircon geochronology from sample SLB, the Shackleton Limestone is interpreted to be an autochthonous deposit with respect to the craton margin, with likely sediment sources in the adjacent East Antarctic shield (Goodge et al., 2004b). Sediment was likely transported across an actively subsiding continental margin prior to a major eustatic transgression resulting in the formation of a stable carbonate platform (Goodge, 1993b; Goodge et al., 2004b).
Syn-Orogenic/Late-Orogenic Stage
By latest-Early Cambrian time, the paleo-Pacific plate had subducted sufficiently to generate volumetrically significant arc-magmatism and significant arc denudation was underway (Goodge et al., 1997; Goodge et al., 2002; Goodge et al., 2004b). A abundant siliciclastic, fine-grained sedimentation across the platform limited the extent of the photic zone, and effectively ended carbonate sedimentation by the end of Early Cambrian time (Myrow et al., 2002; Goodge et al., 2004b). Goodge et al. (2004b) indicated that crustal flexure associated with Ross convergence caused rapid erosion, and progradation of conglomeratic alluvial-fan sediments across the carbonate platform, marking the initiation of the synorogenic stage (Myrow et al., 2002b; Goodge et al., 2004b).
Crustal flexure forced rapid subsidence of the carbonate platform, indicated by deep incision upon the upper bioherm unit of the Shackleton Limestone (Goodge et al., 2004b). Rapid subsidence resulting from compressive forces followed and is supported by the presence of a surface of non-deposition, a phosphatic hardground overlying the Shackleton Limestone (Myrow et al., 2002b). Schlager (1989, 1998) noted that hardgrounds form near the sediment-water contact in a deep-marine, sediment starved settings. Subsequently, phosphate replaces carbonate as seas transgressed across a carbonate ramp, and a hardground forms (Myrow et al., 2002b). As deformation progressed, syntectonic sedimentation continued and is recorded in the stratigraphic section as the Holyoake, Starshot, and Douglas Formations (e.g. Myrow et al., 2002b). These syntectonic deposits formed in response to rapid uplift and denudation of the proximal Ross Magmatic Arc (Goodge et al., 2004b).
The Holyoake Formation. The Holyoake Formation is middle-Lower Cambrian in age and corresponds to an additional shift in the tectonic style within the Ross Orogen (Myrow et al., 2002b). Lithologically, the Holyoake formation is a thin, organic-rich shale with centimeter-scale interbedded nodular carbonate (Myrow et al., 2002b). The basal contact of the Holyoake formation juxtaposes organic-rich shale with the underlying archaeocyathan bioherms of the Shackleton Limestone (Myrow et al., 2002b). This organic-rich shale represents a regional flooding surface that formed in response to rapid transgression as the initiation of deformation associated with the Ross Orogen forced rapid subsidence of the adjacent ocean basin (Myrow et al., 2002b, Goodge et al., 2004b).
The Starshot Formation. The Starshot Formation was first proposed by Laird (1963) for a suite of syn-tectonic facies consisting of interbedded calcareous siltstone, sandstone, and conglomerates (Laird, 1963; Laird et al., 1971; Myrow et al., 2002b). The Starshot Formation is a Middle to Late Cambrian deposit consisting of fluvial and near-shore facies, consisting of turbidites with interbedded conglomerates (Laird, 1963; Laird et al., 1971; Myrow et al., 2002b). While sedimentary structures are rare, faunal assemblages reflect primarily a marine depositional setting for the Starshot Formation (Myrow et al., 2002b). Conglomerate-bearing beds are generally graded near the base, transitioning upsection into convolute, plane and cross-stratified bedding with a laminated siltstone to mudstone cap, typical of the classic turbidite-derived Bouma sequence described by Bouma (1962) (Laird, 1963; Laird et al., 1971; Myrow et al., 2002b; Goodge et al., 2004b). Clasts within the conglomerate beds of the Starshot formation are composed of quartzite, granitoids, felsic volcanic rocks, and archaeocyathan-bearing limestone (Myrow et al., 2002b; Goodge et al., 2004b).
The Starshot represents a definitive change in sediment provenance, recording the shift from Neoproterozoic and older grains, to a significant Ross-aged peak (Myrow et al., 2002b; Goodge et al., 2004b). Throughout the study area the Starshot is deformed and is therefore interpreted to have been deposited prior to the termination of the Ross Orogen (Goodge et al., 2004b). The presence of carbonate clasts within the Starshot supports the interpretation that it is age-equivalent, or younger, than the Shackleton Limestone (Goodge et al., 2004b). Based on the presence of bioherm-bearing clasts within the Starshot Formation, it is interpreted that the Starshot Formation likely represents the initiation and continuation of an erosional response to deformation associated with Ross Orogen convergence (Goodge et al., 2004b). Goodge et al. (2004b) conducted U/Pb detrital zircon analyses on three samples from the Starshot Formation, described below.
Sample HRS, from the lowermost Starshot Formation, contained a heterogeneous distribution of grains with a prominent Ross-age peak (Goodge et al., 2004b). Ages range from 580 – 490 Ma, with a prominent peak at 545 Ma (Fig. 5) (Goodge et al., 2004b). Deposition is restricted to middle Cambrian time or later, as reflected by the youngest distinct population at 510 Ma (Goodge et al., 2004b). The oldest grains from sample HSF are ca. 2.5 Ga but are not statistically significant with respect to the younger populations (Goodge et al., 2004b). The provenance of sample HSF is markedly different from that of underlying Beardmore Group (Goodge et al., 2004b). Sedimentary petrography and detrital zircon geochronology indicate that sample HSF was sourced from a proximal, young, plutonic source (Goodge et al., 2004b).
Samples DSG and SRG were collected from a feldspathic arenite within the Starshot Formation (Goodge et al., 2002, 2004b). Zircon populations are dominated by latest Neoproterozoic and earliest Cambrian constituents including peaks at ca. 560, and 1,100 Ma, reflective of Ross Orogen and Grenville contributions, along with minor older Paleoproterozoic and Archean components (Goodge et al., 2002, 2004b). The Ross-age peak comprised approximately half of the zircons contained within the samples, reflective of the proximal Ross magmatic source (Goodge et al., 2002, 2004b). The youngest zircon population from samples DSG and SRG indicate that the maximum depositional age for the Starshot Formation was during early Cambrian time (Goodge et al., 2002, 2004b).
Goodge et al. (2004b) collected sample USF from a sandstone within the Starshot formation near Mount Ubique. Zircon populations reflect a range of sources, unlike what is observed from stratigraphically lower Starshot samples (Goodge et al., 2004b). Age distributions are primarily Ross-aged, reflected by a population at ca. 520 Ma, but Grenville, Paleoproterozoic and Archean inputs are present to a lesser extent (Goodge et al., 2004b). The maximum depositional age for syn-tectonic sample USF is during the late Cambrian time, indicated by the youngest zircon population at ca. 500 Ma (Goodge et al., 2004b).
Sample DIF is a moderately sorted, moderately mature sandstone comprised of quartz and plagioclase with accessory tourmaline, and titanite (Goodge et al., 2004b). Sample DIF contains a wide range of zircon populations with prominent Ross and Grenville-aged peaks between 580-490 Ma and 1200-950 Ma, both reflecting Ross and Grenville contributions (Fig. 6) (Goodge et al., 2004b). In addition to Ross and Grenville contributions, minor Paleoproterozoic and Archean components also exist (Goodge et al., 2004b). The Ross and Grenville signatures are interpreted to represent first-cycle sedimentation of a proximal Grenville and Ross source; and the Archean and Paleoproterozoic components are likely sourced from distal, crystalline exposures (Goodge et al., 2004b).
The Douglas Conglomerate. The Douglas Conglomerate is a middle Cambrian syn-tectonic unit composed of a range of facies but is predominantly conglomeratic in composition (Laird et al., 1971; Rees and Rowell, 1991; Myrow et al., 2002b). Facies within the Douglas Conglomerate include alluvial, marine, marginal marine, and fluvial deposits with grain sizes ranging from very fine sand to cobble and boulder size grains (Rees and Rowell, 1991). Conglomeratic beds are generally poorly sorted and texturally immature (Rees and Rowell, 1991; Goodge et al., 2004b). Many of the boulder-sized clasts contained within Douglas beds are calcareous in composition and are in contact with archaeocyathan bioherms, reflecting erosion of the underlying Shackleton Limestone during Ross deformation (Rees et al., 1988). Lenticular sandstone beds with a scoured base are common within the Douglas Conglomerate (Myrow et al., 2002b). Abundant siltstone and minor shale components occur and are on average 10 cm thick (Myrow et al., 2002b). Myrow et al. (2002b) concluded based on the range of grain sizes and the general lack of textural maturity that the Douglas Conglomerate reflect deposition immediately outboard of the magmatic arc.

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