5.1. Mg-Temperature Record
 Bottom water temperatures were estimated from the C. mundulus Mg/Ca data (Figure 3) using the C. mundulus-specific Mg-temperature calibration of Lear et al. :
Assuming a minimal change in seawater Mg/Ca from the Miocene to present [Wilkinson and Algeo, 1989; Rowley, 2002], BWT estimates should be accurate to within ±1°C (2σ; Lear et al. ).
 Unlike the Mg/Ca of planktonic foraminifera, the uptake of Mg/Ca into benthic foraminifera remains poorly understood. As a result, there is an ongoing debate in the paleoceanographic community as to whether the relationship between temperature and Mg/Ca in benthic foraminifer is exponential or linear (Lear et al. [2002, 2003], Lea , Marchitto et al. , and others). We chose the Lear et al.  equation to convert Site 1171 Mg/Ca to BWT because our record is exclusively derived from C. mundulus and because of the definite offset between C. mundulus and C. wullerstorfi observed in multispecies records [Lear et al., 2003]. Lear et al.  revised their original core-top Cibicidoides calibration [Lear et al., 2002] to obtain a species specific C. mundulus equation by assuming that the two species have a similar temperature sensitivity but a different preexponential constant.
 More recently, a core-top calibration for C. pachyderma (or C. mundulus) from the Little Bahamas Bank was generated by Marchitto et al. , which suggests a strong linear relationship between Mg/Ca and BWT, with a sensitivity of 0.12 mmol/mol Mg/Ca per °C:
The standard error of the equation is 2.4°C. Marchitto et al.  favor a linear fit to the core top data because an exponential curve exaggerates the temperatures at the cold extreme of the calibration.
 In order to illustrate the effect of using different calibrations on our Site 1171 middle Miocene C. mundulus Mg/Ca record, we have plotted the results of three different equations in Figure 2: The solid line with open circles uses the exponential calibration Lear et al.  and the dashed line employs the linear calibration of Marchitto et al. . We have also plotted the data using the exponential equation of Lear et al.  with the temperature sensitivity of the Marchitto et al.  equation (heavy black line; standard error of ±1.7°C); the Marchitto et al.  sensitivity was substituted into the Lear et al.  equation to reflect the likelihood that the temperature sensitivity of C. mundulus (C. pachyderma) is species specific. Furthermore, we argue that at the cold end of the calibration, the slopes of the linear and exponential equations are essentially identical:
Temperatures derived from the Lear et al.  equation are the warmest of the three equations. The temperatures estimated from hybrid Lear-Marchitto equation are slightly lower than the Lear et al.  temperatures, but within the error of the Lear et al.  equation. The Marchitto et al.  equation yields significantly cooler temperatures, which may reflect the lack of data at the cool (<5°C) end of the calibration [Lear et al., 2003; Marchitto et al., 2007]. Furthermore, because the standard error of the Lear et al.  equation is 1.0°C and the offset between the temperatures estimated from equation (1) and equation (3) are within error of the Lear et al.  calibration, we have chosen to employ this calibration.
 Using the published equation of Lear et al. , we estimate that the Middle Miocene BWTs at Hole 1171C oscillated around 5.3 ± 1.0°C (Figure 3), which is within the range of existing low-resolution Cenozoic benthic Mg/Ca records (4–8°C [Lear et al., 2000; Billups and Schrag, 2002]). However, Site 1171 BWTs are routinely cooler than Mg/Ca records generated at lower latitude sites, a finding consistent with the South Tasman Rise's geography and oceanographic setting [Lear et al., 2000; Billups and Schrag, 2002]. Three intervals of relatively warm BWTs exist in the Site 1171 record, including those at 16.4–16.2 Ma (6.3 ± 0.3°C), 14.5–14.1 Ma (5.6 ± 1.0°C), and 13.5–13.3 Ma (5.7 ± 0.7°C) (Figure 3). Two of these warm intervals occurred during the Miocene Climatic Optimum (17–14 Ma [Flower and Kennett, 1994, and references therein]) and the third at ∼13.5 Ma following the middle Miocene δ18O increase; warming of regional surface water temperatures [Shevenell et al., 2004] and Southwest Pacific (26°S) benthic foraminifer d18O records [Flower and Kennett, 1994] has also been inferred at this time (Figure 3). Cooler BWTs observed at ∼16 Ma (4.7 ± 1.0°C) and 13.6 Ma (4.6 ± 0.7°C) correspond with intervals of more positive δ18O (glacial events Mi-2 and Mi-3 of Miller et al. [1991a]) (Figure 3). Across the middle Miocene d18O increase (14.2–13.8 Ma), BWTs at Site 1171 cooled ∼2°C ± 1.5°C (∼6° to 4°C) (Figure 3), which is similar to cooling estimated using indirect methods [Miller et al., 1991a; Wright et al., 1992; Flower and Kennett, 1994; John et al., 2004].
5.1.1. Accounting for Temporal Seawater Mg/Ca Variability
 The largest uncertainty in estimating Cenozoic paleotemperatures using Mg/Ca relates to temporal variations in seawater Mg/Ca [Lear et al., 2000; Billups and Schrag, 2002]. Changes in CaCO3 sedimentation, dolomite cycling, hydrothermal activity, and/or the hydrologic cycle could alter seawater Mg/Ca [Wilkinson and Alego, 1989; Stanley and Hardie, 1998; Lear et al., 2000; Billups and Schrag, 2002]. However, the long residence times of oceanic Mg2+ (13 Ma) and Ca2+ (1 Ma) [Broecker and Peng, 1982] indicate that, while the absolute values of BWTs may be affected by changing seawater Mg/Ca, the magnitude of temperature change across rapid (<1 Ma) climate transitions should remain unchanged. Furthermore, if the seawater Mg/Ca ratio were driven purely by changes in oceanic Ca cycling, then the ratio of Sr/Ca in seawater should exhibit similar changes to that of Mg/Ca. The Site 1171 benthic foraminifer Sr/Ca record shows no similarity to that of the Mg/Ca [Lear et al., 2003] record and no systematic long-term trend. Thus, we assume that the long-term trends observed in the record are not related to changes in oceanic Ca cycling.
 To assess the uncertainty in the absolute values of our BWT record related to temporal seawater Mg/Ca variations [Wilkinson and Alego, 1989; Stanley and Hardie, 1998], we modified the Lear et al.  calibration equation following Lear et al. :
Where Mg/CaM refers to measured C. mundulus Mg/Ca (this study), Mg/CaSWM to the modeled Mg/Ca of seawater at ∼14 Ma (4.2 mmol/mol [Wilkinson and Algeo, 1989]), and Mg/CaSWP to the Mg/Ca of present day seawater (5.1 mmol/mol [Broecker and Peng, 1982; Stanley and Hardie, 1998]). BWT estimates derived from equation (4) are 1.8°C warmer than those from equation (1) [Lear et al., 2003]. Taking into account uncertainty related to temporal seawater Mg/Ca variations and the calibration, we estimate the absolute uncertainty in Miocene BWT estimates at ±1–3°C. Because the uncertainty related to the changing ratio of Mg to Ca in seawater is within the error of the Lear et al.  calibration, our discussions will focus on the temperature record derived from that equation, not equation (4). Owing to the long residence times of Mg/Ca in the ocean, our data will not be affected on timescales of <1 Ma. Furthermore, on longer timescales it is unlikely that the shape of the middle Miocene curve will change with a shift in the seawater Mg/Ca, but rather the absolute values will be influenced by ±1–3°C.
5.1.2. Middle Miocene Benthic Foraminifer Mg Temperatures (∼14 Ma)
 The middle Miocene δ18O increase at ∼14 Ma is one of the three major δ18O increases of the Cenozoic. At Site 1171, a δ18O increase of 1.2‰ occurs between 14.1 and 13.7 Ma. Across the same time interval, C. mundulus Mg-temperatures cool ∼2 ± 2°C from ∼6 to 4°C, suggesting that roughly 0.75‰ (∼65%) of the 1.2‰ δ18O increase across the middle Miocene climate transition relates to global ice volume [O'Neil et al., 1969]. Thus the Site 1171 benthic Mg/Ca record suggests no significant or permanent cooling following the middle Miocene climate step. Interestingly, the cooling immediately follows the relatively warm temperatures of the Miocene Climatic Optimum (∼15–14 Ma) and is similar in magnitude and range to bottom water temperatures observed in the lower resolution benthic foraminifer Mg/Ca records of both Lear et al.  and Billups and Schrag , despite different Mg-temperature relationships used in each of the studies.
 A ∼2°C warming trend (13.7 to 13.5 Ma) immediately follows the middle Miocene cooling and δ18O increase. Despite the uncertainties presently associated with the benthic Mg-temperature relationship, there is some evidence to suggest that the observed pattern of temperature change across the Middle Miocene δ18O increase is robust. A similar pattern of warming is observed following the major Cenozoic δ18O excursions across the Eocene/Oligocene and Oligocene/Miocene boundaries [Lear et al., 2004]. These authors propose a scenario in which warming after the major Eocene/Oligocene (Oi-1) and early Miocene (Mi-1) Antarctic glaciations is thought to reflect a negative feedback of the climate system related to the reduction of global chemical weathering rates due to extensive continental glaciation. This reduction in weathering is thought to result in an increase in atmospheric pCO2 and a partial melting of Antarctic ice sheets, referred to by Lear et al.  as the “missing sink” mechanism, which is consistent with the evidence from both the middle Miocene global benthic foraminifer δ13C record and proxy atmospheric pCO2 records [Pagani et al., 1999]. Carbon isotope records from Site 1171, which exhibit the global benthic foraminifer δ13C signal, indicate a rise in benthic foraminifer δ13C immediately following the major δ18O increase that is associated with the final orbitally paced global carbon maximum events (CM6) of the Monterey δ13C excursion [Vincent and Berger, 1985; Woodruff and Savin, 1991] and an increase in atmospheric pCO2 as inferred from the δ13C of alkenones preserved in southwest Pacific sediments [Pagani et al., 1999]. Support for reduced Antarctic ice volume at this time comes from both our calculated δ18Osw record (Figure 4; see below for discussion) as well as from the dating of relict surfaces in the Antarctic Dry Valley region (see below for further discussion [Sugden and Denton, 2004; Lewis et al., 2006]).
5.1.3. Ice Volume Estimates from δ18Osw
 To calculate the δ18O of regional seawater (δ18Osw) from paired BWT (1 s.d.: ±1–3°C) and benthic foraminifer δ18O (1 s.d.: ±0.1‰ [Shevenell and Kennett, 2004]) records, we used the C. pachyderma (C. mundulus) specific δ18O paleotemperature equation of Lynch-Stieglitz et al.  combined with the BWTs derived from the Lear et al.  equation:
Results reveal that 0.82 ± 0.43‰ (∼70%) of the 1.2‰ Cibicidoides mundulusδ18O increase between 14.1 and 13.7 Ma relates to Antarctic ice growth and ∼30% relates to cooling (Figure 3). Comparison of the Site 1171 δ18Osw curve with lower resolution records of Lear et al.  and Billups and Schrag [2002, 2003] suggests similarities between the three records (δ18Osw: -0.5–1‰) and confirms at high-resolution previous indirect or lower resolution ice volume estimates [Shackleton and Kennett, 1975; Wright et al., 1992; Lear et al., 2000; Billups and Schrag, 2002, 2003; John et al., 2004]. Thus δ18Osw results from Site 1171 likely reflect global ice volume and not just regional changes in δ18Osw.
 The calculated δ18Osw record from Site 1171 provides novel high-resolution insight into the phasing of ice growth and Southern Ocean temperatures in the Southwest Pacific during the middle Miocene climate transition not available from the previous lower resolution Mg/Ca records of Lear et al.  and Billups and Schrag . A general trend toward more positive δ18Osw values between 15 and 13.8 Ma (Figure 3) at Site 1171 suggests that Antarctic cryosphere expansion began at the height of the warm Miocene Climatic Optimum (∼15 Ma [Shevenell et al., 2004]) and this expansion progressed in a stepwise fashion until 13.8 Ma (Figures 2 and 3). Three intervals of more positive and variable δ18Osw values are superimposed on the general trend and interpreted as intervals of glacial advance (midpoints: 14.9, 14.4, and 13.8 Ma). These glacial advances are generally associated with warmer surface and bottom water temperatures in the South Tasman Rise region (Figure 4). Three intervals of more negative δ18Osw values occur at 15.0, 14.7, and 14.0 Ma and are interpreted as times of glacial retreat. These interglacials are associated with times of cooler surface and bottom water temperatures at Site 1171 (Figure 4). Interestingly, the final interglacial prior to the middle Miocene δ18O increase is associated with a stepwise cooling of South Tasman Rise surface waters interpreted to reflect an increase in the strength and/or northward progression of the Antarctic Circumpolar Current in the Southwest Pacific (Figure 4 [Shevenell et al., 2004]), as well as with the 2 ± 2°C cooling of regional bottom waters (Figure 4).
 The δ18Osw record from Site 1171 indicates that the Antarctic cryosphere underwent a phase of rapid, seemingly orbitally paced ice growth between 15 and ∼14 Ma. The inferred glacial-interglacial cycles in the δ18Osw record (Figure 4) recur every ∼400-ka, suggesting eccentricity pacing of the Middle Miocene Antarctic cryosphere expansion. Similar pacing is observed in global middle Miocene benthic foraminifer δ13C records of the Monterey interval (17–13.5 Ma [Vincent and Berger, 1985; Woodruff and Savin, 1989; Flower and Kennett, 1994; Holbourn et al., 2007]). Such a strong long-period eccentricity signal suggests a central role for internal climate system feedbacks (e.g., ice/albedo, global carbon cycling, ocean circulation changes) in this major Cenozoic climate transition [Shackleton, 2000]. The apparent increase in the intensity of Antarctic glaciations approaching ∼14 Ma provides further evidence for such internal climate system feedbacks.
 Our interpretation of the δ18Osw record from Site 1171 as a proxy record for changes in ice volume is consistent with terrestrial and marine geologic records from Antarctica and its continental margins, which indicate that expansion of Antarctic ice sheets began at ∼15 Ma. Radiogenic isotope records from the circum-Antarctic indicate a shift towards more physical weathering of Antarctica at ∼15 Ma [Vlastelic et al., 2005] while exposure dating of relict surfaces in the Antarctic Dry Valley region suggest an expansion of the ice sheet into the region at 14.8 Ma and a retreat by 13.6 Ma [Sugden and Denton, 2004]. Furthermore, sequence stratigraphic records of eustasy in the middle Miocene [Haq et al., 1987; Miller et al., 1991b; Wright et al., 1992; John et al., 2004] indicate broad similarities to the calculated Site 1171 δ18Osw curve, within current dating resolution of the records. However, none of these records are presently of sufficient orbital-scale resolution to be definitively compared to the calculated δ18Osw record from Site 1171. Ongoing sequence stratigraphic studies of the middle Miocene interval similar to those conducted across the Eocene/Oligocene boundary [Kominz and Pekar, 2001] will likely yield accurate orbital scale records of changes in global eustasy that may be useful for comparison with our record [Miller et al., 2005; K. G. Miller et al., personal communication, 2006].
5.1.4. Importance of Moisture Supply in Middle Miocene Antarctic Ice Sheet Expansion
 Southern Ocean temperatures and seawater δ18O estimates reveal substantial Antarctic ice growth began during the warm Miocene Climatic Optimum (∼15 Ma) (Figure 4 [Shevenell et al., 2004]) when Southwest Pacific bottom water and sea surface temperatures were relatively warm. The Site 1171 benthic Mg/Ca record confirms, at higher resolution, the Mg/Ca findings of Lear et al.  and Billups and Schrag  and other more indirect estimates (Shackleton and Kennett , Wright et al. , John et al. , and others) which suggest that the majority of the middle Miocene δ18O increase at ∼14 Ma was related to an increase in global ice volume. Our records are especially significant in that they suggest that ice growth began during the warmest period of the Neogene (the Miocene Climatic Optimum at ∼15 Ma), ∼1 Ma prior to the globally recognized climate step during a time of inferred low atmospheric pCO2 (Vincent and Berger , Pagani et al. , and others), and progressed in a stepwise orbitally paced fashion between ∼15 and 14 Ma. This pattern of glaciation is further supported by lower latitude benthic foraminifer δ18O records from the subtropical Pacific, which show orbitally paced changes inferred as glacial-interglacial cycles between 14.7 and 13.8 Ma [Holbourn et al., 2005]. On orbital timescales, our records suggest that Antarctic ice growth appears to coincide with times when Southern Ocean temperatures were particularly warm (Figure 4). The observed relationship (both the long-term trend and on orbital scales) between Southwest Pacific temperature and inferred ice growth challenges the notion that meridional heat flux limited Antarctic ice growth [Woodruff and Savin, 1989] and instead provides support for hypotheses positing that poleward heat/moisture supply was essential for Antarctic cryosphere expansion [Schnitker, 1980; Prentice and Matthews, 1991]. It remains unclear as to the origin of this heat/moisture (e.g. Warm Saline Deep Water from the Tethys, proto-North Atlantic Deep Water, or surface/atmospheric sources).
 The importance of moisture availability to the development and maintenance of Antarctic ice sheets is highlighted in terrestrial records from the Antarctic Dry Valleys (Ross Sea sector) (Sugden and Denton  and others). Presently, the majority of Antarctica's precipitation is concentrated in the coastal regions. However, several coastal locations, including the Dry Valleys, appear to have remained predominantly ice-free since ∼13.6 Ma due to limited regional moisture availability [Sugden and Denton, 2004]. This inferred long-term environmental stability suggests that middle Miocene Antarctic cryosphere expansion altered regional heat/moisture supply [Sugden and Denton, 2004] and/or the Antarctica's sensitivity to mid- to low-latitude derived heat/moisture [DeConto and Pollard, 2003]. Geomorphology of the Dry Valleys indicates that relatively humid conditions prevailed at ∼15 Ma and that the East Antarctic Ice Sheet inundated the region by 14.8 Ma and retreated by ∼13.6 Ma [Marchant et al., 1993; Sugden and Denton, 2004]. The timing of this inferred glacial expansion and retreat corresponds with our marine geochemical evidence (Figure 4), providing independent support for our interpretation of δ18Osw as a record of Antarctic ice growth and for changes in moisture flux to the region at the end of the Miocene Climatic Optimum.
 Intensified oceanic and atmospheric circulation capable of altering the flux of low-latitude-derived heat/moisture to the Southern Ocean is inferred during the middle Miocene [Kennett et al., 1985; Vincent and Berger, 1985; Woodruff and Savin, 1989; Flower and Kennett, 1994; Shevenell et al., 2004]. A progressive orbitally paced increase in Antarctic Circumpolar Current strength between 14.2 and 13.8 Ma, inferred from the stepwise cooling of regional surface waters [Shevenell et al., 2004], could have isolated Antarctica from low-latitude heat/moisture sources and acted as a negative feedback towards further ice growth. Alternatively, the progressive reduction of warm low-latitude-derived deep-water associated with the tectonic restriction of the eastern Tethys may have removed an oceanic heat source from the Southern Ocean [Hsu and Bernoulli, 1978; Woodruff and Savin, 1989; Flower and Kennett, 1994]. Although the origin and transport mode (atmosphere or ocean) of Southern Ocean heat remain unconstrained, evidence suggests that the continued isolation of Antarctica (via tectonics and circulation) was critical to the progressive Cenozoic development of the Antarctic cryosphere.
5.1.5. Is There a Role for pCO2 in Middle Miocene Antarctic Cryosphere Expansion?
 Antarctica may have been poised to respond sensitively to poleward heat/moisture transport in the middle Miocene due to relatively low inferred atmospheric pCO2 (220–250 ppmv; Vincent and Berger , Pagani et al. , Pearson and Palmer , DeConto and Pollard , and others). Model results suggest that under declining pCO2 and other greenhouse gases (e.g., CH4), Antarctic snowline elevations would have dropped, increasing the area of Antarctica available for glaciation [DeConto and Pollard, 2003]. Our high-resolution geochemical records from the Southwest Pacific suggest that Antarctic cryosphere expansion began in earnest at ∼15 Ma (Figures 2 and 3), when atmospheric pCO2 levels reached the lowest inferred levels of the Miocene [Pagani et al., 1999; Pearson and Palmer, 2000] and regional Southern Ocean temperatures were relatively warm (Figure 4 [Shevenell et al., 2004]). Thus Site 1171 geochemical records provide support for models suggesting that the Antarctic cryosphere may be especially sensitive to poleward heat/moisture flux under low pCO2 boundary conditions [DeConto and Pollard, 2003]. In the early stages of middle Miocene ice expansion (∼15 Ma), warm Southern Ocean waters may have supplied heat/moisture to the Antarctic continent (Figure 4). As glaciation progressed, internal climate feedbacks (e.g., invigorated circumpolar circulation, ice/albedo) likely further isolated Antarctica from lower latitude derived heat/moisture. The stepwise character of the δ18Osw (ice volume) record and the Mg-derived SST records (Figure 4 [Shevenell et al., 2004]) support this interpretation. By ∼13.5 Ma, Antarctic cryosphere expansion had slowed/ceased, Southern Ocean temperatures cooled, and higher threshold pCO2 levels are inferred (Figures 2 and 3 [Pagani et al., 1999; Shevenell et al., 2004]).
 Not only did Antarctica's sensitivity to poleward heat/moisture flux appear to decrease with rising atmospheric pCO2 [DeConto and Pollard, 2003; Sugden and Denton, 2004], but that the expansion of Antarctic ice sheets likely exerted influence on the global carbon cycle in a similar fashion to that observed following both the Oi-1 and Mi-1 glaciations [Lear et al., 2004]. We speculate that this control resulted in both the final and largest marine δ13C increase (CM6) of the “Monterey interval” at 13.7 Ma (Flower and Kennett  and others) as well as the decline in marine carbonate δ13C at 13.5 Ma that marked the end “Monterey” interval (16.5–13.5 Ma [Vincent and Berger, 1985]) and a shift in global carbon cycle dynamics.