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).