Records from the Antarctic Margin

The Oligocene-Miocene boundary interval was first sampled in Antarctica at DSDP Site 270 in the Eastern Ross Sea (Fig. 9.1). Drilling recovered a succession of silty mudstones including glaciomarine sediment, which spans the Oligocene-Miocene boundary (Hayes and Frakes, 1975; Leckie and Webb, 1983). Although an abrupt lithological change with a potential hiatus at the Oligocene-Miocene boundary is noted by Leckie and Webb (1983), poor chronological resolution prevents unambiguous correlation with the Mi1 event and the earliest Miocene. The only exposed Oligocene-Miocene boundary strata reported from Antarctica crop out on King George Island (Fig. 9.1) and include the Destruction Bay Formation (Latest Oligocene) and Cape Melville Formation (earliest Miocene; Birkenmajer et al., 1985; Birkenmajer, 1987). In a recent summary of stratigraphy and facies of the succession, Troedson and Riding (2002) concluded that a significant glacial advance occurred at the boundary and that chronological control was good enough to suggest a correlation with the Mi1 event. The facies indicate a significant regional ice grounding event across Bransfield Strait and beyond the South Shetland Islands (Troedson and Riding, 2002). Unfortunately, drilling on the shelf and slope south of the South Shetland Islands (ODP Leg 178; Barker and Carmerlenghi, 2002) did not yield any more definitive records of the Oligocene-Miocene boundary.

Late Oligocene/early Miocene strata reported from the East Antarctic margin include Maud Rise (ODP Leg 113 sites 689 and 690; Barker et al., 1988a), the Weddell Sea margin (ODP Leg 113 Site 693; Barker et al., 1988a) and Kerguelen Plateau (ODP Leg 120 sites 747 and 748; Schlich and Wise, 1992; Fig. 9.1). The record at Maud Rise is relatively thin and comprises exclusively siliceous and carbonate ooze, although rare glacial drop stones are reported in strata from Site 689 (Barker et al., 1988a). At the Weddell Sea margin Oligocene-Miocene sediments are also fine grained and include diatom mud, clay and ooze (Barker & Carmerlenghi, 2002). Oligocene-Miocene boundary sediments at Kerguelen Plateau are also carbonate ooze (Schlich and Wise, 1992), however, foraminifera preservation and age resolution were good enough at ODP Site 747 to yield a benthic oxygen isotope stratigraphy across the boundary at that site (Wright and Miller, 1992; Billups and Schrag, 2002; Fig. 9.5) and the amplitude of the Mi1 event was much reduced compared to equatorial values, with a 818O shift of only 0.3 m across the Oligocene-Miocene boundary. A strontium isotope stratigraphy has also been assembled using planktic foraminifera from ODP Site 747 (Oslick et al., 1994). Oslick et al. (1994) reported significant increases in 87Sr/86Sr following the early Miocene Mi isotope events with a ~ 1 myr lag. They suggested that this increase and similar subsequent stepwise increases in early-middle Miocene oceanic 87Sr resulted from changes in the glacial state of East Antarctica.

Drilling in Prydz Bay (ODP Legs 119 and 188; Hambrey et al., 1991; Cooper and O'Brien, 2004) did not yield any Oligocene-Miocene age strata

Figure 9.5: Oxygen isotope and Mg/Ca data from the Oligocene-Miocene boundary interval in ODP Site 747 from Kerguelen Plateau. Data from Billups and Schrag (2002) adjusted to time scale of Billups et al. (2004).

and Hambrey et al. (1991) concluded that this was due to erosion beneath an expanded middle-Miocene ice sheet.

9.3.1. McMurdo Sound, South Western Ross Sea

The Antarctic Oligocene-Miocene record is most complete in the Victoria Land Basin as recovered in the CIROS-1 and CRP-2A drill holes (Barrett, 1989, Cape Roberts Science Team, 1999; Figs. 9.1 and 9.6). As with Prydz Bay, much of the Oligocene record of the Victoria Land Basin is marked by significant hiatuses (Wilson et al., 1998, 2000a, b; Florindo et al.. 2005), however, the latest Oligocene-early Miocene is preserved in both records (Naish et al., 2001; Wilson et al., 2002; Roberts et al., 2003). On the basis of radiometric, biostratigraphic and magnetostratigraphic data, Wilson et al. (2002) placed the Oligocene-Miocene boundary at 183.7 m in the CRP-2A core at the base of a normal polarity interval correlated with Polarity Subchron C6Cn.2n using Berggren et al.'s (1995) time scale. However, following the astronomical revision of the late Oligocene through early Miocene time scale (Billups et al., 2004; Gradstein et al, 2004; Pälike et al., 2004), Naish et al. (2008) placed the boundary at 130.27 m in an unconformity in the CRP-2A core and revised the age of strata underlying the unconformity to encompass Polarity Chron C7n. Roberts et al. (2003) placed the Oligocene-Miocene boundary at 247 m in the CIROS-1 core ~35km south of CRP-2A. However, following the revision of Naish et al. (2008), the Oligocene-Miocene boundary in the CIROS-1 core more likely occurs in an unconformity at 92m (Fig. 9.7). The boundary immediately overlies an unconformity at 248.71 m, which might represent as much as 1 myr following the age revision of Antarctic shelf diatom zones (Scherer et al., 2000) implied by Naish et al. (2008).

The strata recovered in both the CRP-2A and CIROS-1 drill holes are cyclic in nature and interpreted to represent periodic advance and retreat of ice across the Antarctic margin concomitant with sea-level fall and rise, respectively (Fielding et al., 1997; Naish et al., 2001). Each sequence is organised into a vertical succession, which begins with an erosion surface and is followed by a diamictite and sandstone, which gives way to sparsely fossiliferous bioturbated mudstone representing a cycle of glacial advance and retreat followed by open water conditions across the site of deposition (Naish et al., 2001) in concert with changes in relative sea-level (Dunbar et al., 2008; Fig. 9.6). Naish et al. (2008) estimated the glacioeustatic influence on relative water depth changes by deconvolving the tectonic, isostatic and palaeobathy-metric components of water depth. These results are consistent with the 818O

Depositions] environment

Inner shelf to shoreface Outer shelf under iceberg influa

Ice proximal, Ice contact.

Deltaic Grounding Ihe fen

Depositions] environment lea proximal, Ice contact.

Ice proximal, ica contact, subglacial

Inner shelf, under glacial Influence

Ice proximal, marina Inner shelf Ice proximal, grounding line, subglacial

Shoaling to shoreface with iceberg influence Outer shelf under Iceberg Influence

Ice proximal, grounding line fan Ice proximal, Ice contact, subglacial

Shorafacafinnar-shelf under glacial influence Ice proximal, Ice contact, subglacial Inner shelf Grounding line fan delta

Ice proximal, ice contact, subglacial

Outer shelf under minor iceberg influence Ice proximal, ice contact, subglacial

Inner shelf to shoreface Outer shelf under iceberg influa

Ice proximal, Ice contact.

Deltaic Grounding Ihe fen lea proximal, Ice contact.

Ice proximal, ice contact, subglacial

Outer shelf under minor iceberg influence Ice proximal, ice contact, subglacial

Cooler Temperatures indicated by pollen and reduced metwater indicated by loss of Cymatiosphaera

Miocene

Equivalent position of Nothofagus Leaf in CIROS-1 core

Cooler Temperatures indicated by pollen and reduced metwater indicated by loss of Cymatiosphaera

Miocene

Ollgocene

Equivalent position of Nothofagus Leaf in CIROS-1 core

Figure 9.6: Environmental proxy data for the upper part of the CRP-2A drill core. Grain size and clast data, and sequence stratigraphic, palaeobathymetric, depositional environment and ice margin interpretations are from Cape Roberts Science Team (1999). CIA, chemical index of alteration (data from Passchier and Krissek, 2008). ARM, anhysteretic remanent magnetization (data from Verosub et al., 2000). Temperature and meltwater indicators are discussed in Barrett (2007). Nothofagus leaf in CIROS-1 was identified at 215.5mbsf immediately underlying the Oligocene-Miocene boundary as identified from the revised age model presented in this paper (Fig. 9.7).

Depth Lithology fmbsf)

Antarctic shelf

^40Ar/39Ar

Antarctic shelf

^40Ar/39Ar

Depth (mbsf)

90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350

CIROS-1

Antarctic shelf

~ . . diatom Lithology Polarity zones

CIROS-1

Antarctic shelf

~ . . diatom Lithology Polarity zones

(Ma)

Figure 9.7: Revised age models for the CRP-2A and CIROS-1 drill cores from McMurdo Sound using the astronomical time scale of Billups et al. (2004) following the arguments in Naish et al. (2008). Magnetic polarity and diatom biostratigraphy data are from Wilson et al. (2000a, b, 2002) and Roberts et al. (2003). CIROS-1 age model is constrained by diatom biostratigraphic zones with revised ages from Naish et al. (2008).

sea-level calibration of Pekar et al. (2002) from the New Jersey margin. Each cycle is interpreted to represent between 10 and 40 m of eustatic variation in the late Oligocene with perhaps 50 m of sea-level fall concomitant with an ice sheet some 20% larger than present coincident with the unconformity, which is correlated with the Oligocene-Miocene boundary by Naish et al. (2008). Grounding of ice across the site at the Oligocene-Miocene boundary is also confirmed by macro- and micro-structures indicative of glacio-tectonic deformation (Passchier, 2000; van der Meer, 2000).

The Oligocene-Miocene boundary is marked by some significant changes in physical properties in the CRP-2A core. Late Oligocene sedimentary cycles underlying the boundary are 55-60 m thick and relatively complete, whereas early Miocene sedimentary cycles are 10-20 m thick and truncated (Cape Roberts Science Team, 1999; Fig. 9.6). The clay mineralogy of the strata across the Oligocene-Miocene boundary in CRP-2A records stable physical weathering conditions (Ehrmann et al., 2005). Major element ratios derived from mudrock geochemistry for the same strata show significant shifts in the chemical index of alteration (CIA) across the Oligocene-Miocene boundary, which indicate periods of increased physical weathering and mechanical erosion associated with glacial advance (Passchier and Krissek, 2008). The CIA data reported by Passchier and Krissek (2008; Fig. 9.6) was corrected for the presence of primary volcanic detritus in order to reflect the palaeoclimatic record within the mudrock geochemistry. Short-lived glacial events at ~23, ~21 and ~19Ma indicated by the CIA data are correlated with the Mi events by Passchier and Krissek (2008) and interpreted to represent significant climatic and ice-sheet events in East Antarctica. Magnetic properties also show a marked change across the Oligocene-Miocene boundary at 130.27 m in the CRP-2A core (Verosub et al., 2000). An earlier change in magnetic properties at 270 m (late Oligocene) is attributed to inception of the McMurdo Volcanic Province (Verosub et al., 2000).

Despite these changes in physical properties across the Oligocene-Miocene boundary in the CRP-2A core, palynological data although sparse due to low concentrations of organic matter, indicate a partially open landscape dominated by small Nothofagus (Southern Beech) stands or sparse tundra vegetation persisting through the late Oligocene-early Miocene (Askin and Raine, 2000). Oligocene-Miocene strata in the CIROS-1 core also contain similar amounts of pollen (Mildenhall, 1989) and a Nothofagus leaf fossil was preserved in latest Oligocene strata of the CIROS-1 Core (Hill, 1989; Figure 9.7). Mean summer temperature records derived from the (K+Na)/Al ratios of the CRP cores (Passchier and Krissek, 2008) indicate relatively constant mean summer temperatures of ~10°C in the latest Oligocene dropping to ~6°C in the early Miocene. Marine palynomorphs, however, suggest a more significant change following the Oligocene-Miocene boundary with a significant reduction in the occurrence of prasinophyte algae, which is taken to indicate a reduction in offshore meltwater influence and hence cooler climates (Barrett, 2007).

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