The evolution of pCO2, ice volume and climate during the middle Miocene
Highlights
► Boron isotope based pCO2 reconstruction of the middle Miocene. ► Coupled changes in climate (ice-volume) and pCO2 through the middle Miocene. ► High pCO2 (350–400 ppm) during the middle Miocene Climatic Optimum. ► pCO2 starts to decrease at ∼14.7 Ma to ∼200 ppm by ∼13 Ma.
Introduction
Over at least the last 550 thousand years Earth's climate, global sea level, and the CO2 content of the atmosphere have been tightly coupled (Rohling et al., 2009). However, the nature of this relationship in the future in response to anthropogenic climate change is hard to predict (Rahmstorf et al., 2007) due to questions regarding the behaviour of the large continental scale ice sheets of Greenland and Antarctica in a rapidly warming climate (Alley et al., 2005). Observations of continental ice sheet stability in the past, when the Earth's climate was significantly warmer than today, may not represent the immediate future due to hysteresis in ice sheet behaviour (Pollard and DeConto, 2005), but do provide a valuable test-bed of our understanding of ice sheet growth and decay (Pollard and DeConto, 2005).
The Cenozoic era saw a fundamental transition in the Earth's climate state from the greenhouse climate of the Cretaceous to our modern icehouse, characterised by continental ice on both poles (Zachos et al., 2001). Large, continental scale ice sheets first formed on Antarctica at the Eocene–Oligocene boundary (33.7 Ma; Coxall et al., 2005), while the northern hemisphere is thought to have remained largely ice free until the Late Pliocene (∼3 Ma; DeConto et al., 2008, Zachos et al., 2001, Mudelsee and Raymo, 2005). The modern Antarctic Ice Sheet (AIS) is made up of two parts, the large, sluggish, land-based Eastern AIS and the smaller, more dynamic, largely marine-based Western AIS. Ice sheet modelling indicates that, once formed, the land-based AIS is very difficult to melt (Pollard and DeConto, 2005). This hysteresis effect arises because the bright surface of the ice cap maintains a cold, elevated interior that resists melting during a global warming event. Such modelling also indicates that CO2 levels must rise to significantly higher values in order to melt the EAIS than are required to form it in the first place (∼1000 ppm vs. 750–840 ppm; Pollard and DeConto, 2005). In contrast, no significant hysteresis is thought to characterise the marine-based AIS (Pollard and DeConto, 2009) or is documented in records of the northern hemisphere ice sheets response to climate forcing (Rohling et al., 2009).
Throughout the Oligocene and Miocene there is a growing body of evidence for orbitally paced and relatively large (±>20 m) changes in sea level and ice volume, often attributed to growth/decay of the land-based Antarctic Ice Sheet (e.g. Pekar and deConto, 2006, Holbourn et al., 2007, Shevenell et al., 2008, Kominz et al., 2008, Liebrand et al., 2011, Passchier et al., 2011). The middle Miocene Climatic Optimum (MCO; 17–15 Ma) is one such time period where the Antarctic Ice Sheet may have reduced in size, with some studies suggesting a decrease to around 10–25% of its modern volume (De Boer et al., 2010). This period is however characterised by rather limited global warmth (+2–4 °C compared to preindustrial; You et al., 2009) with reconstructions of CO2 showing either relatively stable levels with concentrations similar to the pre-industrial (200–300 ppm; Pagani et al., 1999) or, if elevated at all (350–450 ppm; Kürschner et al., 2008), well below the modelled threshold values required to melt the land-based AIS (e.g. ∼1000 ppm; Pollard and DeConto, 2005). Therefore, a puzzling aspect of the climate evolution of the MCO is why the observed rapid and orbitally paced variations in sea level and ice volume (e.g. Holbourn et al., 2007) take place without appearing to show significant hysteresis. This discrepancy between geological observations and ice sheet modelling implies that either our understanding of how the continental ice sheets grow and decay is at fault or our reconstructions of the middle Miocene CO2 and climate are in error. Here we present a new multi-site boron isotope-based reconstruction of surface water palaeo-pH and use this to quantify the evolution of atmospheric CO2 during this time period (e.g. Hönisch and Hemming, 2005, Foster, 2008). This new pCO2 reconstruction allows, for the first time, a direct investigation of the relationship between pCO2 and the cyrosphere of the middle Miocene.
Section snippets
Sample locations and site details
We examine Miocene aged sediments from two open ocean Sites: ODP 761 from the Wombat Plateau (16°44.23′S, 115°32.10′E and water depth of 2179 m) and ODP 926 from the Ceara Rise (3°43.148′N, 42°54.507′W and water depth of 3598 m). Both Sites are currently located in regions where surface water is close to equilibrium with the atmosphere with respect to CO2 (Fig. 1; Takahashi et al., 2009).
The foraminifera sampled from ODP Hole 761B were taken from between 35 and 50 m below sea floor (mbsf), which
δ11B of G. sacculifer from the middle Miocene
The boron isotope data for G. sacculifer from our two Sites are in good agreement and show lower δ11B values (15.0±0.4‰ to 15.8±0.3‰) during the MCO, with higher values (16.0±0.3‰ to 17.1±0.2‰) either side of this warm period, largely sympathetic with the benthic foraminiferal δ18O record for this time period (Fig. 5). The δ11B of foraminifera such as G. sacculifer (δ11Bsac) has a positive relationship with pH (Fig. 2), and thus a negative relationship with [CO2]aq. Therefore, the δ11B of G.
Conclusions
Through a combination of planktic δ11B and benthic δ18O we have shown that there is a close coupling between pCO2 and ice-volume (hence climate) through the middle Miocene. The nature of this coupling indicates the presence of a dynamic ice sheet(s) during the MCO and MMCT that, within our sampling resolution (<300 kyr), exhibited little or no apparent hysteresis and a linear relationship with climate forcing by CO2. With anthropogenic emissions of CO2 continuing to rise (and potentially
Acknowledgements
This work was supported by NERC grants (NE/D00876X/2, NE/I006176/1 and NE/D008654/1) to G.L.F. and C.H.L. and a NERC studentship to J.W.B.R. We thank the (I)ODP for supplying samples that were essential to this study. Kat Allen and two anonymous reviewers are thanked for their comments that greatly improved this manuscript. Chris Coath and Huw Boulton are acknowledged for their assistance in the laboratory and Paul Wilson, Martin Palmer, Damon Teagle and Eelco Rohling are thanked for their
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