Figure 6: Sunbird-1 LA-ICP-MS Mg/Ca derived SST using the approach of Anand et al. (2003) without a pH correction compared to and SST estimates at contemporaneous sites from (a) Uk37, and (b) foraminiferal geochemistry. Estimates applying Uk37are from ODP Site 722 (Huang et al., 2007) in the Arabian Sea, ODP & IODP Sites 846 (Herbert et al., 2016), 850 (Zhang et al., 2014), 1241 (Seki et al., 2012), and U1338 (Rousselle et al., 2013) in the Eastern Equatorial Pacific, terrestrial outcrops in Malta (Badger et al., 2013). Estimates applying the foraminiferal Mg/Ca proxy are from ODP Sites 761 (Sosdian and Lear, 2020) and terrestrial outcrops in Malta (Badger et al.,2013). ODP Site 761 data is displayed on an alternative axis as SST anomalies relative to the baseline average from 16.0 – 15.5 Ma. Two temperature estimates using the δ18O of exceptionally preserved foraminifera from Tanzania are also shown (Stewart et al., 2004). The upper limit for the Uk37 proxy (29⁰C) is marked by the thick dashed black line. All previously published records used for comparison are kept on their original age models. Supplementary Figure S9 provides LA-ICP-MS Mg/Ca sea surface temperatures using the alternative approach of Evans et al. , (2016).
Although not a true tropical location, and consisting of only three data points, the Badger et al. (2013) Mg/Ca record from the Mediterranean estimates SST of ~27.5⁰C between 13.5 and 13 Ma, within the Sunbird-1 SST uncertainty envelope (Figure 6b). Mg/Ca-SST records based on less well-preserved planktic foraminifera also suggest stable tropical SST between 13.8 and 11.4 Ma (Sosdian and Lear, 2020) (Figure 6b). Furthermore, well preserved planktic foraminifera from clay-rich sediments of coastal Tanzania yield Indian Ocean sea surface temperatures of 27⁰C at 12.2 Ma and 29⁰C at 11.55 Ma using the δ18O paleo-thermometer (Stewart et al. , 2004), again in agreement with the Sunbird-1 temperature estimates (Figure 6b). It is worth noting that this study, as well as the tropical SST records of Herbert et al. (2016) and references therein, do not sample the warm pool of the Western Pacific. Sea surface temperature estimates for the western equatorial Pacific using the TEX86 paleothermometer suggest a slight, ~1°C, SST decrease between 12 Ma and 9 Ma, whilst those for the eastern equatorial Pacific are more or less constant across the same interval (Zhang et al., 2014).
Although the estimates provided by the Sunbird-1 record suggest absolute tropical sea surface temperatures remained relatively stable through the mid-late Miocene, some temporal variability does persist. Between 11.8 Ma and 11.7 Ma SST drops sharply by ~3⁰C. Excluding one value of 28.6⁰C at 11.62 Ma, this decrease in SST to ~24-25⁰C persists for ~300 kyr before recovering to pre excursion values by 11.5 Ma. However, no transient decrease in sea surface temperature is recorded from contemporaneous alkenone based estimates of tropical SST utilizing the Uk37 proxy from the Arabian Sea (Huang et al. , 2007), and the Eastern Equatorial Pacific (Herbert et al. , 2016; Rousselle et al. , 2013; Seki et al. , 2012; Zhang et al. , 2014) (Figure 6a). We therefore suggest that the observed transient ~3⁰C SST decrease is not the result of a global driver, and supports a mechanism causing local ocean cooling of the surface waters at Sunbird-1. An alternative hypothesis is that an unaccounted increase in local salinity and/or pH, lowering foraminiferal Mg/Ca ratios, caused a bias to cooler temperatures between ~11.8 and 11.5 Ma. Assuming constant SST, the observed ~0.7 mmol/mol decrease in Mg/Ca would require a salinity increase on the order of 5.0 PSU (Hönisch et al. , 2013; Gray et al. , 2018). This salinity increase equates to a 0.8 ‰ change in δ18O using the Indian Ocean δ18Osw-salinity relationship of LeGrande and Schmidt (2006) (Equation 5). As well as being an extremely large change in salinity, the planktic foraminiferal δ18O record does not support such a significant change in sea surface salinity between ~11.8 and 11.5 Ma (Figure 5b). However, we do acknowledge that a contribution from increased salinity control cannot be discounted. Despite incorporating varying pH from a globally distributed set of open ocean sites (Sosdian et al. , 2018), a localized increase in pH at Sunbird-1 cannot be ruled out. This possibility may be particularly relevant considering the land-proximal, tectonically active nature of the study site. A further possibility is that selective dissolution of foraminiferal chambers precipitated during warmer seasons occurred during post-burial diagenetic alteration, causing an apparent ~3°C lowering of SST between 11.8 Ma and 11.5 Ma. However, mean D. altispira test weights suggest that there was no increased dissolution of the foraminiferal tests through this interval of lower LA-ICP-MS Mg/Ca derived SST (Supplementary Table S11 and Supplementary Figure S10).
Therefore, our preferred interpretation is for a local cooling between ~11.8 and 11.5 Ma. The lack of a marked increase in the planktic δ18O record at this time implies that the cooling was associated with a freshening of surface waters (Figure 5c). Interestingly, this interval corresponds to a period of very high sedimentation rates (Supplementary Figure S1), which might be consistent with enhanced precipitation and runoff, lowering regional surface salinity.
4.3 Implications for the global climate state during the mid-late Miocene
Previous studies utilizing the Uk37proxy suggest a substantial cooling of sea surface temperature at mid-to-high latitudes in both hemispheres between 10 and 5.5 Ma, whilst tropical sea surface temperatures show limited cooling in the late Miocene prior to ~7 Ma (Herbert et al. , 2016;LaRiviere et al. , 2012). The absolute tropical SST record reported in this study supports the finding that the latitudinal temperature gradient steepened from ~10 Ma, as the climate system transitioned towards its modern-day state. Furthermore, support for the absolute temperatures reconstructed by the alkenone proxy suggests that the interval between 10 and 7.5 Ma was associated with enhanced polar amplification, significantly greater than that calculated for the greenhouse climate of the Eocene (Cramwinckel et al. , 2018). There is little evidence for a significant change in pCO2 in this interval (Sosdian et al. , 2018;Stoll et al. , 2019) (Figure 7). We speculate that the marked regional cooling between 10 and 7.5 Ma perhaps reflects processes internal to the climate system, involving for example ocean-atmospheric heat transport, sea ice extent, or changes in regional cloud cover. A combined data-modelling approach would help constrain possible factors and explore potential relationships between this highly heterogenous cooling and the CO2 drawdown that was associated with the subsequent global late Miocene Cooling starting ~7.5 Ma (Figure 7).