Identification of the Oligocene Miocene Boundary

Early definition of the Oligocene-Miocene boundary relied on the identification of the last occurrence of the calcareous nannofossil Dictyococcites

Figure 9.1: Polar stereographic projection showing the location of key Oligocene-Miocene boundary sections discussed in the text. Abbreviations of locations: numbers are DSDP/ODP drill sites; AP, Antarctic Peninsula; BS, Bransfield Strait; CIROS, Cenozoic Investigations in the Ross Sea Drilling Project; CR, Ceara Rise; CRP, Cape Roberts Drilling Project; DP, Drake Passage; EAIS, East Antarctic Ice Sheet; KGI, King George Island; KP, Kerguelen Plateau; MR, Maud Rise; PB, Prydz Bay; RS, Ross Sea; SSI, South Shetland Islands; VLB, Victoria Land Basin; WAIS, West Antarctic

Ice Sheet; WS, Weddell Sea.

Figure 9.1: Polar stereographic projection showing the location of key Oligocene-Miocene boundary sections discussed in the text. Abbreviations of locations: numbers are DSDP/ODP drill sites; AP, Antarctic Peninsula; BS, Bransfield Strait; CIROS, Cenozoic Investigations in the Ross Sea Drilling Project; CR, Ceara Rise; CRP, Cape Roberts Drilling Project; DP, Drake Passage; EAIS, East Antarctic Ice Sheet; KGI, King George Island; KP, Kerguelen Plateau; MR, Maud Rise; PB, Prydz Bay; RS, Ross Sea; SSI, South Shetland Islands; VLB, Victoria Land Basin; WAIS, West Antarctic

Ice Sheet; WS, Weddell Sea.

C6Cr C6Cn.3nC6Cn.2r C6Cn.2n OC6Cn.1n C6Br

C6Cr C6Cn.3nC6Cn.2r C6Cn.2n OC6Cn.1n C6Br

1995)

Figure 9.2: History of the calibration of the Geomagnetic Polarity Time Scale in the vicinity of the Oligocene-Miocene boundary. Boundary position in Ness et al. (1980) and Berggren et al. (1985) is fixed to the last occurrence of the nannofossil Reticulofenestra bisectus and in Cande and Kent (1992, 1995), Gradstein et al. (2004) and Billups et al. (2004) is fixed to the base of magnetic polarity subchron C6Cn.2n. Ness et al. (1980), Berggren et al. (1985), Cande and Kent (1992, 1995) are calibrated by various mid-ocean ridge spreading rate models.

Billups et al. (2004) and Gradstein et al. (2004) are calibrated by astronomical tuning at ODP Site 1090.

1995)

Figure 9.2: History of the calibration of the Geomagnetic Polarity Time Scale in the vicinity of the Oligocene-Miocene boundary. Boundary position in Ness et al. (1980) and Berggren et al. (1985) is fixed to the last occurrence of the nannofossil Reticulofenestra bisectus and in Cande and Kent (1992, 1995), Gradstein et al. (2004) and Billups et al. (2004) is fixed to the base of magnetic polarity subchron C6Cn.2n. Ness et al. (1980), Berggren et al. (1985), Cande and Kent (1992, 1995) are calibrated by various mid-ocean ridge spreading rate models.

Billups et al. (2004) and Gradstein et al. (2004) are calibrated by astronomical tuning at ODP Site 1090.

bisectus (23.7 Ma; Berggren et al., 1985). However, this has proved problematic in the colder waters, coarser sediments and hiatus prone strata of the Antarctic and Southern Ocean. The reassignment of the boundary by Cande and Kent (1992, 1995) to the slightly older base of Magnetic Polarity subchron C6Cn.2n (23.8 Ma; Fig. 9.2) has made its identification more straightforward in Antarctica and the Southern Ocean but only in relatively complete and continuous stratigraphic successions (Wilson et al., 2002; Roberts et al., 2003). More recently, the recognition of astronomically influenced cyclical physical properties and 818O records in continuously deposited deep successions has enabled astronomical calibration of late Oligocene through early Miocene time. The astronomical calibration suggested that, while still coincident with the base of subchron C6Cn.2n, the boundary was in fact nearly a million years younger (22.97 0.1 Ma, Shackleton et al., 2000, Palike et al., 2004; 23.03 Ma, Billups et al., 2004, Gradstein et al., 2004; Fig. 9.2). The climatic significance of this was outlined by Zachos et al. (2001b) and Palike et al. (2006) who recognised the coincidence of the Oligocene-Miocene boundary and the Mi1 isotope excursion with an unusual coincidence of low eccentricity and low-amplitude variability in obliquity of the Earth's orbit (Fig. 9.3). This would have placed the Earth in a sustained period of unusually low seasonality (cold summers), which Zachos et al. (2001b) claimed would have limited polar summer warmth and encouraged ice growth at the poles. Equally, within a few hundred thousand years, the coincidence of increased eccentricity and highamplitude variability in obliquity would have resulted in warmer polar summers and increased summer melt.

9.2. Proxy Records 9.2.1. The Isotopic Record

Oxygen isotope ratios (d18O) in foraminiferal tests from deep-sea sedimentary records have long been recognised to represent the Cenozoic climatic (temperature, sea level and ice volume) history of the Earth (e.g. Shackleton et al., 1977 and references therein). However, deciphering the climatic history of Antarctica from 818O values alone in deep-sea records has always proven difficult due to the ambiguity of influence on the signal from the volume of ice on land versus isotopic fractionation, which is related to the water temperature during the precipitation of calcite (Miller et al., 1991). Early studies attempted to separate the two influences by focusing their analyses on

Eccentricity Mg Temperature ('C)

Eccentricity Mg Temperature ('C)

Figure 9.3: Oxygen isotope and Magnesium temperature data across the Oligocene-Miocene boundary interval (data from Lear et al., 2004). Obliquity and eccentricity orbital target data from Laskar et al. (2004). All data are plotted against the astronomical time scale presented by Billups et al. (2004).

Figure 9.3: Oxygen isotope and Magnesium temperature data across the Oligocene-Miocene boundary interval (data from Lear et al., 2004). Obliquity and eccentricity orbital target data from Laskar et al. (2004). All data are plotted against the astronomical time scale presented by Billups et al. (2004).

foraminiferal species, such as benthic forms, known to live in water masses thought to be relatively stable in their temperature history (Shackleton and Kennett, 1975; Kennett, 1977), hence deducing that any shorter-term fluctuation was due to ice volume rather than temperature fractionation. While these studies recognised major threshold changes in the climatic deterioration of Antarctica, assuming no Northern Hemisphere ice sheets at the time, their resolution was limited and the record incomplete across the Oligocene-Miocene boundary. In a higher-resolution compilation of Oligocene-Miocene 818O records, Miller et al. (1991) recognised a relatively transient ca. 2myr period and 1 m amplitude cyclicity in Miocene isotope values (isotope events Mi1-Mi7). The 1 m shifts, they suggested were of the same order as the threshold shifts identified by earlier studies (Kennett and Shackleton, 1976) and represent similar volumes of ice accumulation on the Antarctic craton. The most significant of these shorter order isotopic events, Mi1, was coincident with the Oligocene-Miocene boundary (Fig. 9.3). Originally defined by Miller et al. (1991) from DSDP Site 522 (Figs. 9.1 and 9.3), the event has subsequently been confirmed at numerous locations and the timing, duration and magnitude refined (Zachos et al., 1997; Paul et al., 2000; Zachos et al., 2001a, b; Billups et al., 2002, 2004, Billups and Schrag, 2002).

For at least 2myr prior to the Oligocene-Miocene boundary, 818O values were relatively stable and of low-amplitude variability (<0.5 m; Paul et al., 2000). In contrast, the Mi1 event represents a dramatic ~1 m increase in d18O, over a 250 ky period immediately prior to the Oligocene-Miocene boundary peaking coincident with the boundary (Billups et al., 2004; Fig. 9.3). Peak values persisted for only ~20ky before returning to similarly low amplitude but slightly increased late Oligocene mean 818O values over the first ~ 120 ky of the early Miocene (Paul et al., 2000). The covariance of the isotope signal in both benthic and planktic species in the late Oligocene at Equatorial Atlantic ODP Site 929 led Paul et al. (2000) to conclude that the variability was primarily ice volume driven. However, the Mi1 event, itself, is of relatively lower amplitude in planktic records which led Paul et al. (2000) to suggest that only 0.5 m is likely due to ice-volume effects, which, using the late Pleistocene calibration, represents growth of an ice sheet in Antarctica of similar proportion to the present-day East Antarctic Ice Sheet (EAIS). The remaining 0.5 m, they concluded was due to a 2°C cooling of bottom waters at ODP Site 929 in the western Equatorial Atlantic.

Another approach to deciphering temperature versus ice-volume components of the deep-sea d18O signal was employed by Lear et al. (2004) who determined palaeotemperature independently from Mg/Ca ratios in for-aminifera tests across the Oligocene-Miocene boundary at ODP Site 1218 in the eastern Equatorial Pacific (Figs. 9.1 and 9.3). The 2-3°C of cooling predicted in the equatorial Atlantic by Paul et al. (2000) immediately prior to Oligocene-Miocene boundary was confirmed in the Equatorial Pacific by Lear et al. (2004). However, peak cooling preceded peak 818O values at ODP Site 1218 and a slight (1°C) warming in bottom waters of the Equatorial Pacific is recorded immediately prior to and during peak Mi1 818O values followed by another 2°C of warming post-Mi1. A similar temperature variation was observed ~ 800 ky prior to the Oligocene-Miocene boundary. Lear et al. (2004) concluded that the Mi1 event did indeed represent a significant increase in continental ice volume and, given the warming influence predicted by Mg/Ca ratios, perhaps a larger ice-volume growth than that predicted by Paul et al. (2000). Billups and Schrag (2002) also suggested that the d18O record from ODP Site 747 represented an ice-volume signal because paired Mg/Ca measurements suggested little change in ocean temperature through the early Miocene. Recent work, however, suggests caution when interpreting stable intervals in Mg/Ca ratios from deep-water sites due to potential saturation of carbonate, which might affect the partitioning of Magnesium into benthic foraminifera (Elderfield et al., 2006; Lear et al., 2008).

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