Climate model sensitivity to changes in Miocene paleotopography

Herold et al. (2009) consider factors underlying the midMiocene climate optimum (MMCO), comparing atmospheric paleo-CO2 levels [Pagani et al. 1999 (pre-10 Ma: 5300 ppm CO2); Royer et al. 2001 (18–14 Ma: *300– 370 ppm CO2)] and paleotemperatures (Flower & Kennett 1994: þ68C) with 20–21st century climate parameters. On this basis, the authors suggest a ‘. deficiency in our understanding of the global climate system’ (Herold et al. 2009 p. 1049), and proceed to test the effects of ocean–atmosphere heat exchange, plate tectonics, subduction, orographic elevation, uplift and erosion on the MMCO, using general circulation models. Revised estimates by You et al. (2009) refer to MMCO CO2 level of 460–580 ppm, narrowed from earlier estimates of 300–600 ppm. These values are consistent with foraminiferal B/Ca proxies that indicate MMCO CO2 levels of 430 ppm and temperatures *3–68C warmer than the present (Tripati et al. 2009). Cenozoic climates (Pearson & Palmer 2000; Zachos et al. 2001, 2008; Berner 2004; Royer et al. 2004, 2007) underwent overall cooling during the Pleistocene, in particular during the last 800 ka for which ice-core studies indicate atmospheric CO2 levels between *180 and 280 ppm, closely coupled with temperature and sealevel (Hansen et al. 2007). However, the controls of earlier warm periods, such as the MMCO and the midPliocene (Robinson et al. 2008), when ice sheets were smaller relative to the Pleistocene, are less well constrained. CO2 estimates referred to by Herold et al. (2009) are based on dC paleo-CO2 proxies measured on diunsaturated alkenones and shallow-dwelling planktonic foraminifera (Pagani et al. 1999) and proxy stomatal index of Ginkgo biloba and Metasequoia glyptostroboides (Royer et al. 2001). The more recent stomatal-based paleo-CO2 estimates for the period 25–12 Ma (Late Oligocene–Middle Miocene) use fossil leaf remains (Laurophyllum pseudoprinceps, Laurus nobilis, Ocotea foetens, Ginkgo biloba), suggesting CO2 levels in the approximate range of 300–600 ppm (Kürschner et al. 2008; Royer 2008) (Figure 1). Since stomatal indices reflect CO2 levels during the original life-span of the fossil plant, rather than time-averaged values, they can be compared with contemporaneous CO2 levels, as contrasted with CO2-e levels (CO2-e: CO2þ equivalent CH4 forcing). This allows estimates of climate sensitivity (CS: mean global temperature rise per doubling of CO2: Charney 1979; Hansen et al. 2007) as follows. Mid-Miocene: CO2 rise from minimum Early and Middle Miocene values of *280 ppm to the MMCO maximum value of 5520 ppm in relation to a temperature rise of 2.48C (Kürschner et al. 2008 figure 3) yields a CS value of about 2.88C, similar to Charney’s (1979) CS value of 3.0+ 1.58C. Mid-Pliocene: An estimate based on a rise of CO2 level to about 400 ppm and temperatures by 38C higher than at present (Lunt et al. 2009), assuming background CO2 levels of about 250–300 ppm, would yield a CS value of about 5–98C. Last Glacial termination: Estimates based on the last glacial termination (180–280 ppm CO2; *4.58C: Hansen et al. 2007) yield a CS value of *88C, probably reflecting melting of the extensive Laurentian and Fennoscandian ice sheets and the ice melt feedback loop, i.e. loss of albedo and gain of infrared absorption by opened water, which precedes CO2 rise (Hansen et al. 2007). 18th century to present: CS estimates based on a CO2 rise from 280 to 388 ppm (IPCC 2007) and a rise of mean global temperature by about þ0. 88C since 1750 (IPCC 2007), plus about þ0. 58C currently masked by sulfur aerosols (IPCC 2007), yields CS value of *3.48C, not accounting for the effects of current albedo loss due to melting cryosphere. The more than doubling of the CS index during Pleistocene glacial terminations probably reflects the greater extent of ice sheets during the glacial periods and therefore the more important role of their melting upon glacial terminations. Whereas geotectonic transitions, including orographic elevation, continent–ocean patterns, plate subduction, ocean current systems, ocean–atmosphere heat exchange, uplift and erosion rates (Ruddiman & Prell 1997) affect long-term CO2 levels, the relatively shorter mid-Miocene and midPliocene (Robinson et al. 2008; Lunt et al. 2009) warm Australian Journal of Earth Sciences (2010) 57, (377–379)