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\begin{document}

\title{Tropical Sea Surface Temperatures following the Middle Miocene Climate
Transition from Laser-Ablation ICP-MS analysis of glassy foraminifera}



\author[1,1]{Michael Grahame Nairn}%
\author[1,1]{Caroline H. Lear}%
\author[1,1]{Sindia Sosdian}%
\author[2,2]{Trevor Bailey}%
\author[3,3]{Simon Beavington-Penney}%
\affil[1]{Cardiff University}%
\affil[2]{National Museums and Galleries of Wales}%
\affil[3]{University of Manchester}%


\vspace{-1em}



  
  \date{May 20, 2024}


\begingroup
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\selectlanguage{english}
\begin{abstract}
The mid-to-late Miocene is proposed as a key interval in the transition
of the Earth's climate state towards that of the modern-day. However, it
remains a poorly understood interval in the evolution of Cenozoic
climate, and the sparse proxy-based climate reconstructions are
associated with large uncertainties. In particular, tropical sea surface
temperature (SST) estimates largely rely on the unsaturated alkenone
Uk37 proxy, which fails to record temperatures higher than 29@\selectlanguage{english}C, the
TEX86 proxy which has challenges around its calibration, and Mg/Ca
ratios of poorly preserved foraminifera. We reconstruct robust,
absolute, SSTs between 13.5 Ma and 9.5 Ma from the South West Indian
Ocean (paleolatitude \textasciitilde{}5.5@\selectlanguage{english}S) using Laser-Ablation (LA-)
ICP-MS microanalysis of glassy planktic foraminiferal Mg/Ca. Employing
this microanalytical technique, and stringent screening criteria,
permits the reconstruction of paleotemperatures using foraminifera which
although glassy, are contaminated by authigenic coatings. Our absolute
estimates of 24-310\selectlanguage{english}C suggest that SST in the tropical Indian Ocean was
relatively constant between 13.5 and 9.5 Ma, similar to those
reconstructed from the tropics using the Uk37 alkenone proxy. This
finding suggests an interval of enhanced polar amplification between 10
and 7.5 Ma, immediately prior to the global late Miocene Cooling.%
\end{abstract}%



\sloppy


Michael G. Nairn\textsuperscript{1}, Caroline H.
Lear\textsuperscript{1}, Sindia M. Sosdian\textsuperscript{1}, Trevor R
Bailey\textsuperscript{2}, and Simon
Beavington-Penney\textsuperscript{3}

\textsuperscript{1}School of Earth and Environmental Sciences, Cardiff
University, Cardiff, UK,

\textsuperscript{2}Department of Natural Sciences, Amgueddfa Cymru,
National Museum Wales, Cathays Park, Cardiff,

\textsuperscript{3}BG Group, 100 Thames Valley Park Drive, Reading, RG6
1PT, UK*

Corresponding author: Michael G. Nairn (NairnMG@cardiff.ac.uk)

*Current address: Department of Earth and Environmental Sciences, The
University of Manchester, Williamson Building, Oxford Road, Manchester,
M13 9PL, UK.

Key Points:

\begin{itemize}
\tightlist
\item
  Laser-Ablation ICP-MS facilitates absolute sea surface temperature
  reconstructions using foraminifera with diagenetic coatings.
\item
  Tropical sea surface temperatures remained relatively stable at
  24-31\selectlanguage{ngerman}°C following the Miocene Climate Transition.
\item
  Development of an increased latitudinal temperature gradient began
  prior to the Late Miocene Cooling.
\end{itemize}

Abstract

The mid-to-late Miocene is proposed as a key interval in the transition
of the Earth's climate state towards that of the modern-day. However, it
remains a poorly understood interval in the evolution of Cenozoic
climate, and the sparse proxy-based climate reconstructions are
associated with large uncertainties. In particular, tropical sea surface
temperature (SST) estimates largely rely on the unsaturated alkenone
U\textsuperscript{k}\textsubscript{37} proxy, which fails to record
temperatures higher than 29\selectlanguage{english}@C, the TEX\textsubscript{86} proxy which has
challenges around its calibration, and Mg/Ca ratios of poorly preserved
foraminifera. We reconstruct robust, absolute, SSTs between 13.5 Ma and
9.5 Ma from the South West Indian Ocean (paleolatitude
\textasciitilde{}5.5@\selectlanguage{english}S) using Laser-Ablation (LA-) ICP-MS microanalysis
of glassy planktic foraminiferal Mg/Ca. Employing this microanalytical
technique, and stringent screening criteria, permits the reconstruction
of paleotemperatures using foraminifera which although glassy, are
contaminated by authigenic coatings. Our absolute estimates of 24-310\selectlanguage{english}C
suggest that SST in the tropical Indian Ocean was relatively constant
between 13.5 and 9.5 Ma, similar to those reconstructed from the tropics
using the U\textsuperscript{k}\textsubscript{37} alkenone proxy. This
finding suggests an interval of enhanced polar amplification between 10
and 7.5 Ma, immediately prior to the global late Miocene Cooling.

1 Introduction

The mid-late Miocene is an important interval in the evolution of global
climate through the Cenozoic, representing a key period in the
transition out of the warm, dynamic climate state of the Miocene
Climatic Optimum (MCO) into a more stable unipolar icehouse world
(\emph{Badger et al.} , 2013; \emph{Foster et al.} , 2012; \emph{Greenop
et al.} , 2014; \emph{Sosdian et al.} , 2018). Despite being
characterized by similar to modern day atmospheric CO\textsubscript{2}
concentrations (\emph{Foster et al.} , 2012; \emph{Sosdian et al.} ,
2018; \emph{Super et al.} , 2018), middle Miocene mean global
temperatures were likely significantly warmer than the modern day
(\emph{Pound et al.} , 2011;\emph{Rousselle et al.} , 2013). This has
been used to suggest a decoupling of global temperature and atmospheric
CO\textsubscript{2}forcing (\emph{LaRiviere et al.} , 2012; \emph{Pagani
et al.} , 1999), a characteristic which general circulation models
struggle to simulate (\emph{Knorr et al.} , 2011; \emph{von der Heydt
and Dijkstra} , 2006). It has also been suggested that the late Miocene
was an additional important key step in the transition to our modern
climate state, as high latitudes cooled more than low latitudes, leading
to a marked steepening of latitudinal temperature gradients
(\emph{Herbert et al.} , 2016).

The late Miocene Cooling (LMC) between \textasciitilde{} 7.5 and 5.5 Ma
was a global phenomenon (\emph{Herbert et al.} , 2016) perhaps
associated with decreasing atmospheric pCO\textsubscript{2} (\emph{Stoll
et al.} , 2019). The increase in the equator to pole surface temperature
gradients was not associated with an increase in the benthic
foraminiferal oxygen isotope record, implying that it occurred in the
absence of a large increase in continental ice volume (\emph{Herbert et
al.} , 2016). Polar amplification in the LMC is consistent with
estimates for other time intervals (e.g., \emph{Cramwinckel et al.}
(2018)). However, the LMC was also preceded by a significant cooling of
mid to high southern and northern latitudes, a heterogenous cooling at
high northern latitudes, and a muted, limited cooling in the tropics
(\emph{Herbert et al.} , 2016). This heterogenous cooling perhaps
suggests an unusually high polar amplification factor for the interval
immediately preceding the LMC. Potential changes in the Earth System
that could impact the magnitude of polar amplification include sea ice
extent, vegetation induced changes in albedo, cloud cover, or
ocean-atmosphere heat transport. Constraining the magnitude and timing
of the steepening of latitudinal temperature gradients is therefore
important for understanding the factors driving the late Miocene surface
cooling specifically, and Earth System feedbacks more generally.
Ideally, this would be achieved through a combined data-modelling
approach using multi-proxy temperature reconstructions spanning a range
of latitudes to increase confidence in calculated changes in temperature
gradients.

Despite the significance of this climate interval, the evolution of
global sea surface temperatures (SST) and hence temperature gradients
during the mid-late Miocene is relatively poorly constrained due to a
paucity of complete well-preserved sedimentary successions (\emph{Lunt
et al.} , 2008). The widespread carbonate dissolution, which
dramatically reduced the sediment carbonate content and preservation
quality in deep marine sediments, is termed the middle-late Miocene
carbonate crash (\emph{Farrell et al.} , 1995; \emph{Jiang et al.} ,
2007; \emph{Keller and Barron} , 1987; \selectlanguage{ngerman}\emph{Lübbers et al.} , 2019;
\emph{Lyle et al.} , 1995). In addition to these dissolution issues, the
majority of foraminifera-bearing Miocene sections are comprised of
carbonate rich sediments which have undergone some degree of
recrystallisation. The oxygen isotopic composition of planktic
foraminifera that have undergone recrystallisation in seafloor sediments
has been shown to be biased to colder temperatures (\emph{Pearson et
al.} , 2001). While planktic foraminiferal Mg/Ca appears to be less
affected than \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O, the impact of recrystallisation on
reconstructed Mg/Ca sea surface temperatures remains an additional
source of uncertainty (\emph{Sexton et al.} , 2006). As a consequence,
many mid-late Miocene absolute sea surface temperature reconstructions
are restricted to estimates based on the unsaturated alkenone proxy and
the TEX\textsubscript{86} proxy (\emph{Herbert et al.} , 2016;
\emph{Huang et al.} , 2007; \emph{LaRiviere et al.} , 2012;
\emph{Rousselle et al.} , 2013; \emph{Seki et al.} , 2012; \emph{Zhang
et al.} , 2014). These records show a cooling in the late Miocene which
begins around 10 Ma at high northern and southern latitudes. However,
significant cooling in the tropics is not apparent in the alkenone
records until \textasciitilde{}7.5 Ma, while atmospheric
pCO\textsubscript{2}reconstructions also suggest a significant decline
from this time (\emph{Sosdian et al.} , 2018; \emph{Stoll et al.} ,
2019). At face value therefore, these records imply an interval of
enhanced polar amplification between 10 Ma and 7.5 Ma in the absence of
significant drawdown of CO\textsubscript{2} or increase in ice volume
(\emph{Herbert et al.} , 2016; \emph{Sosdian et al.} , 2018). One
significant caveat to this interpretation is that the
Uk\textsubscript{37} alkenone proxy becomes saturated above 280\selectlanguage{english}C
(\selectlanguage{ngerman}\emph{Müller et al.} , 1998) and the late Miocene tropical SSTs prior
to 7.5 Ma are at this limit (\emph{Herbert et al.} , 2016). Therefore,
an alternative interpretation of the data would be that the high
latitudes and the tropics cooled synchronously from \textasciitilde{}10
Ma, but the initial cooling in the tropics was not able to be recorded
by the Uk\textsubscript{37} alkenone proxy. Corroboration of the
absolute Uk\textsubscript{37} alkenone temperatures by an independent
proxy would therefore confirm the timing of the global late Miocene
Cooling and the possible interval of enhanced polar amplification
between 10 Ma and 7.5 Ma.

Here we present a new planktic foraminiferal Mg/Ca record from the
Sunbird-1 industry well cored offshore Kenya by BG Group. Critically,
middle to late Miocene sediments in Sunbird-1 are hemipelagic clays,
which has resulted in glassy preservation of the foraminifera. However,
the foraminifera are coated with metal-rich authigenic coatings, which
are not removed by standard cleaning techniques. Planktic foraminifera
were therefore analyzed by laser ablation ICP-MS to obtain Mg/Ca from
the primary foraminiferal test and hence enable estimation of absolute
SSTs.

2 Materials and Methods

2.1 Site location, stratigraphy, and age control

This study utilizes 91 cuttings, spanning 273 meters at burial depths
ranging from 630 m to 903 m, recovered by BG Group from the Sunbird-1
well offshore Kenya (04° 18' 13.268'' S, 39° 58' 29.936'' E; 723.3 m
water depth) (Figure 1, Supplementary Table S1). Sedimentation at
Sunbird-1 through the studied interval (9.5-13.5 Ma) is dominated by
clays; the fraction of the sediment \textgreater{}63\selectlanguage{greek}µ\selectlanguage{english}m averages 11.5\%
(Supplementary Table S1), much lower than typical carbonate-rich
deep-water sites. The impermeable nature and chemical composition of
clay-rich sediment reduces diagenetic alteration of primary
foraminiferal calcite, making them ideal targets for geochemical
analysis (\emph{Pearson et al.} , 2001; \emph{Sexton et al.} , 2006).
Tests displaying the desired exceptional preservation appear glassy and
translucent under reflected light, and SEM imaging shows retention of
the foraminiferal original microstructure (\emph{Pearson and Burgess} ,
2008). This style of preferential glassy preservation, as displayed in
the Sunbird-1 well, is rare to absent in published records from Miocene
foraminifera.\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image1/image1}

\end{center}
\end{figure}

\textbf{Figure 1:} Location of the Sunbird-1 study site (black square).
Other sites for which there are mid to late Miocene sea surface
temperature reconstructions from Mg/Ca (red circles),
ẟ\textsuperscript{18}O (blue circles), unsaturated alkenones (green
circles) and TEX\textsubscript{86} (pink diamonds) are shown. Figure
produced using Ocean Data Viewer (\emph{Schlitzer, R.,} 2018) using
modern-day mean annual sea surface temperature data from the World Ocean
Database.

Micropaleontological and calcareous nannoplankton assemblages for
Sunbird-1 were analyzed by Haydon Bailey and Liam Gallagher of Network
Stratigraphic Consulting. Biostratigraphic datums, correlated with the
astronomical timescale of \emph{Raffi et al.} (2020), are based on the
planktic foraminifera zonations of \emph{Wade et al.} (2011) and
calcareous nanofossil zonations of \emph{Backman et al.} (2012). An age
model was constructed by linear interpolation between these
biostratigraphic datums (Supplementary Figure S1). Sedimentation rates
were \textasciitilde{}3 cm/kyr immediately following the middle Miocene
Climate Transition (MMCT), and subsequently increased to
\textasciitilde{}17 cm/kyr between 11.8 and 11.5 Ma, before decreasing
to \textasciitilde{}8 cm/kyr until 9.5 Ma.

2.2 Foraminiferal stable isotope analysis

Up to 12 individual tests of the planktic
foraminifer\emph{Globigerinoides obliquus} showing glassy preservation
were used.\emph{G. obliquus} is an extinct, symbiont-bearing species
with a tropical to subtropical paleogeographical distribution, and is
interpreted as a surface mixed-layer dweller (\emph{Aze et al.} ,
2011;\emph{Keller} , 1985). The assertion that \emph{G. obliquus}
inhabits and calcifies in the surface mixed layer (\emph{Aze et al.} ,
2011;\emph{Keller} , 1985) is supported by multispecies analyses from a
10.0 Ma sediment sample from the Indian Ocean offshore Tanzania
showing\emph{G. obliquus} to have the most negative
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O (-2.5%0) of all species (Paul Pearson, personal
communication, 2019). Tests were crushed between two glass plates
ensuring all chambers were opened. Any visible infill was removed using
a fine paintbrush under a binocular microscope. Fine clays and other
detrital material on the outer surface of the test were removed by
rinsing three times in 18.2 M DI water, ultrasonicating for 5-10
seconds in analytical grade methanol, and finally rinsing a further time
in 18.2 M DI water. Samples were analyzed at Cardiff University on a
ThermoFinnigan MAT253 with online sample preparation using an automated
Kiel IV carbonate device. Results are reported relative to Vienna Pee
Dee Belemnite, and long-term uncertainty based on repeat analysis of
NBS-19 is \selectlanguage{ngerman}±0.08 \selectlanguage{english}%0 (n=469, 2 standard deviations) and on repeat analysis
of BCT63 is +-0.07 \selectlanguage{english}%0 (n=310, 2 standard deviations). Data is available in
Supplementary Table S2.

2.3 Solution ICP-MS trace metal analysis

Between 10 and 15 individuals of the planktic
foraminifer\emph{Dentoglobigerina altispira} from the 250 -- 355 \selectlanguage{greek}µ\selectlanguage{english}m size
fraction were picked and weighed on a six-decimal-place balance to
determine average test weight. Individual tests were then crushed
between two glass plates ensuring all chambers were opened. Due to the
low foraminiferal abundance it was not possible to analyze the same
species for stable isotope and trace metal composition. Any visible
infill was removed using a fine paintbrush under a binocular microscope.
Fragments were cleaned to remove clays and organic matter following the
standard protocol (\emph{Barker et al.} , 2003; \emph{Boyle and Keigwin}
, 1985). Due to the clay-rich nature of the sediment the clay removal
procedure was conducted twice. To test for the possible presence of
metal oxides half of the samples were reductively cleaned between the
clay removal and oxidative cleaning steps. Samples were dissolved in
trace metal pure 0.065 M HNO\textsubscript{3} and diluted with trace
metal pure 0.5M HNO\textsubscript{3} to a final volume of 350 \selectlanguage{greek}µ\selectlanguage{english}l.
Samples were analyzed at Cardiff University on a Thermo Element XR
ICP-MS using standards with matched calcium concentrations to reduce
matrix effects (\emph{Lear et al.} , 2010; \emph{Lear et al.} , 2002).
Together with Mg/Ca, several other ratios (Al/Ca, Mn/Ca, and U/Ca) were
analyzed to screen for potential contaminant phases. Data are available
in Supplementary Table S3. Long-term analytical precision for Mg/Ca
throughout the study is better than 2\%.

2.4 Laser ablation-ICP-MS analysis

Direct sampling of solid phase material via laser ablation (LA-) allows
for geochemical analyses through individual foraminiferal tests at the
sub-micron scale when coupled to an inductively-coupled-plasma mass
spectrometer (ICP-MS) (\emph{Detlef et al.} , 2019; \emph{Eggins et al.}
, 2004; \emph{Evans et al.} , 2015a; \emph{Fehrenbacher et al.} ,
2015;\emph{Hines et al.} , 2017; \emph{Petersen et al.} , 2018;
\emph{Reichart et al.} , 2003). A key advantage of analyzing the trace
element composition of foraminifera using LA-ICP-MS over the more
traditional solution-based ICP-MS is the ability to recognize the
diagenetically altered portions of the tests, allowing identification of
the primary calcite (\emph{Creech et al.} , 2010; \emph{Hasenfratz et
al.} , 2016;\emph{Pena et al.} , 2005). The elemental composition of
this primary calcite can provide important information about
palaeotemperature (\emph{Nooijer et al.} , 2017; \emph{Eggins et al.} ,
2003; \emph{Pena et al.} , 2005) and other paleo-environmental
conditions such as pH (\emph{Mayk et al.} , 2020; \emph{Thil et al.} ,
2016) and oxygenation (\emph{Koho et al.} , 2015; \emph{Petersen et al.}
, 2018).

Up to six specimens of \emph{D. altispira} per sample were selected from
44 depth intervals through the Sunbird-1 core for LA-ICP-MS analysis.
Foraminiferal sample preparation included the removal of fine clays and
other detrital material on the outer surface of the test using DI water
and methanol, but the more aggressive oxidative and reductive steps
(\emph{Barker et al.} , 2003; \emph{Boyle and Keigwin} , 1985), were not
required for laser ablation analysis (\emph{Vetter et al.} , 2013). The
cleaned tests were mounted onto glass slides using double sided carbon
tape and were allowed to dry before being mounted into the sample cell
(\emph{Evans et al.} , 2015b; \emph{Fehrenbacher et al.} ,
2015;\emph{Hines et al.} , 2017).

Analyses were performed using an ArF excimer (193nm) LA- system with
dual-volume laser-ablation cell (RESOlution S-155, Australian Scientific
Instruments) coupled to a Thermo Element XR ICP-MS. Optimized ablation
parameters and analytical settings determined for analyzing foraminifera
in the Cardiff University CELTIC laboratory (Supplementary Table S4;
(\emph{Detlef et al.} , 2019; \emph{Nairn} , 2018)) were used for this
study. Three cleaning pulses to remove any contaminant on the outer
\textasciitilde{}0.5 \selectlanguage{greek}µ\selectlanguage{english}m of the test surface were included prior to
analysis. We analyzed \textsuperscript{25}Mg,
\textsuperscript{27}Al,\textsuperscript{43}Ca, \textsuperscript{55}Mn
and\textsuperscript{88}Sr, each isotope having a constant 50 ms dwell
time. Typically, intervals with elevated Mn and Al in concert with
elevated Mg are interpreted as being contaminant phases (e.g., Fe-Mn
oxides-hydroxides or clays), and are commonly found on the inner and
outer test surface (\emph{Barker et al.} , 2003; \emph{de Nooijer et
al.} , 2014; \emph{Hasenfratz et al.} , 2016; \emph{Koho et al.} ,
2015;\emph{Pena et al.} , 2005).

Where possible, three laser spot depth profiles were collected on each
of the penultimate (f-1) and previous (f-2) chambers by ablating with
100 consecutive laser pulses in one position on the test. Assuming that
each laser pulse only ablates a \textasciitilde{}0.1 \selectlanguage{greek}µ\selectlanguage{english}m layer of calcite
(\emph{Eggins et al.} , 2003), we estimate the profile to represent a
transect through the test wall approximately 15 \selectlanguage{greek}µ\selectlanguage{english}m long. However, in
some cases older chambers were required to ensure six laser profiles per
specimen were analyzed (\emph{Nairn} , 2018). NIST SRM 610 glass
standard was measured between every six laser profiles, and NIST SRM 612
at the beginning and end of analyses from each sample depth. The
reference values for elemental concentrations in both silicate glass
standards are taken from the GEOREM website
(http://georem.mpch-mainz.gwdg.de/sample\_query\_pref.asp), updated from
\emph{Jochum et al.} (2011a). NIST SRM 612 was used to determine long
term external reproducibility using NIST SRM 610. For Mg/Ca, NIST 612
(n=90) had an accuracy of 12.0\% and a precision of 3.7\% relative to
the reported value. A similar \textasciitilde{}12\% negative offset
relative to the reported value of NIST 610-calibrated NIST 612 has been
observed over a much longer period of data collection (\selectlanguage{ngerman}\emph{Evans and
Müller} , 2018). To supplement this assessment, we also conducted
accuracy tests using the GOR-132 and KL-2 MPI-DING glasses (\emph{Jochum
et al.} , 2011b). For this, GOR-132 and KL2 were treated as unknowns,
with both NIST 610 and NIST 612 as calibration standards. For Mg/Ca,
GOR-132 (n=25) had an accuracy of 1.1\% and a precision of 3.2\%
relative to the reported value, and KL-2 (n=25) had an accuracy of 0.6\%
and a precision of 2.6\% relative to the reported value when calibrated
using NIST 610. These values increased to 10.9\% and 9.4\% for GOR-132,
and 8.2\% and 5.4\% for KL-2 when calibrated using NIST 612. The NIST
610-calibrated data presented here supports the determination
of\emph{Evans et al.,} (2015a) that the Mg values for NIST 612 requires
reassessment.

An important issue related to accuracy is that because a
well-characterized, homogenous calcite reference material is not
currently available for laser ablation use, the glass standards we used
have a different matrix to the calcite foraminifera tests (\emph{Evans
et al.} , 2015a; \emph{Evans and Müller} , 2018; \emph{Fehrenbacher et
al.} , 2015). Therefore, while we have high confidence in the accuracy
of the intra-and inter-specimen geochemical variability described in
Section 2.6, we must consider the possibility of an analytical bias in
the absolute geochemical composition of foraminiferal tests determined
by laser ablation ICP-MS. One way to assess the magnitude of such
potential bias is to analyse foraminiferal samples by both solution and
laser ablation ICP-MS. However, it is important to note that the
corrosive cleaning protocol for solution analysis tends to slightly
lower primary test calcite Mg/Ca, an issue that is routinely
circumvented by employing the same cleaning on calibration samples for
paleotemperature reconstructions (\emph{Barker et al.,} 2003). For the
purpose of this study, it is therefore important that our LA-ICP-MS
technique gives values that are consistent with our samples analysed by
solution ICP-MS. We are able to make a direct comparison of our youngest
samples in this way, because these do not have significant authigenic
coatings biasing the solution analyses. For these samples, our solution
and laser ablation results are in excellent agreement, which gives us
confidence in the LA-ICP-MS values for the older samples, where we know
the solution ICP-MS results are compromised by authigenic coatings
(Supplementary Figure S2). Furthermore, we note that if future work
indicates a consistent offset between laser ablation Mg/Ca analyses of
carbonates and silicate glasses, owing to their differing matrices, our
standard values reported above will allow our data to be corrected to
obtain an accurate composition of the uncleaned foraminiferal calcite.

2.5 LA-ICP-MS data processing and screening

Each individual laser ablation profile was carefully inspected and
processed using the SILLS data reduction software package
(\emph{Guillong et al.} , 2008) following the established protocol
outlined in \emph{Longerich et al.} (1996). Profiles generally followed
one of two patterns: (i) a rise from background values to a transient
peak, followed by a somewhat lower plateau, or (ii) a rise from
background values to a general plateau (Figure 2). There are two likely
explanations for the initial transient peak in some isotope profiles:
ablation of authigenic coatings enriched in some trace metals, or laser
ablation induced isotope fractionation (so-called ``pit effects''). We
favor the first explanation because we used the same operating
parameters on every profile, and would therefore expect any ``pit
effects'' to be consistent among the profiles. Furthermore, profiles
containing the transient peaks were more prevalent in the older part of
the record, where our solution Mn/Ca analyses demonstrate the presence
of authigenic coatings. Therefore, we assume the transient peaks
represent contaminated portions of the test and exclude those regions.
The integration interval for the profile was selected based upon the
following three criteria: (i) stable \textsuperscript{43}Ca counts,
indicating ablation of calcite, (ii) stable Mg/Ca signal, indicating a
consistent primary calcite phase, (iii) flat Mn/Ca and Al/Ca signals,
avoiding any peaks indicating intervals of contamination (Figure 2).\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image2/image2}

\end{center}
\end{figure}

\textbf{Figure 2:} Representative LA-ICP-MS Mg, Al, and Mn profiles
demonstrating the selection of background (grey panel) and sample (blue
panel) signals for a profile with an authigenic coating (A, B) and for a
profile without an authigenic coating (C, D). Both examples are shown in
raw isotopic counts (A, C), and ratios mode (B, D) where the isotopes of
interest are relative to \textsuperscript{43}Ca, the internal standard.
In both examples the x axis is analysis time (seconds), and the y axis
is the raw intensity of the isotopes or ratios on a log scale. The
sample interval is selected to avoid the elevated Mg/Ca, Mn/Ca, and
Al/Ca at the outer surface of the test.

Individual depth profiles were corrected by first subtracting the mean
background signal (determined from \textasciitilde{}15 seconds of data
acquired when the laser was turned off prior to ablation). The repeated
analysis of the NIST 610 standard reference material was used to
linearly correct for any instrumental drift. Typically, this is small,
\textless{}2\%, because of the good counting statistics and stable data
acquisition during ablation. The ablation profiles were normalized
to\textsuperscript{43}Ca as the internal standard and elemental
concentrations (TM/Ca) were calculated, assuming 40 wt \% for
CaCO\textsubscript{3}.

Following data processing, rigorous screening of the Mg/Ca ratios for
the influence of intratest contamination was conducted. It is important
to recognise that Mn/Mg and Al/Mg of contaminant phases vary greatly,
such that there is no single universal threshold for these elements that
can be applied in every situation (\emph{Lear et al} ., 2015). For the
Sunbird-1 samples we examined co-variation of Mg/Ca and Mn/Ca and chose
to exclude all samples above a Mn/Ca threshold of 200 \selectlanguage{greek}µ\selectlanguage{english}mol/mol
(Supplementary Figure S3). Consideration of Al/Ca was more complex, as
some samples with extremely high Al/Ca (\textgreater{}1000 \selectlanguage{greek}µ\selectlanguage{english}mol/mol) was
not associated with markedly elevated Mg/Ca. This result demonstrates
that aluminum is sporadically present in foraminiferal tests in variable
phases (with differing Al/Mg). We therefore used a dual-pronged
approach, considering both Al/Ca and intra-sample heterogeneity. We
excluded profiles where two conditions were met: (i) Al/Ca was
\textgreater{}100 \selectlanguage{greek}µ\selectlanguage{english}mol/mol, and (ii) the associated Mg/Ca was
substantially elevated relative to the other depth profiles from the
same sample.

2.6 Determination of mean foraminiferal test Mg/Ca by laser ablation

Geochemical heterogeneity exists both within an individual foraminiferal
test and between foraminiferal tests from the same sample (\emph{Eggins
et al.} , 2004; \emph{Fehrenbacher and Martin} , 2014; \emph{Sadekov et
al.} , 2008; \emph{Sadekov et al.} , 2005). Therefore, several laser
ablation profiles are required to produce a consistent Mg/Ca ratio for
temperature reconstructions. Here we analyzed ten depth profiles through
each of ten individual \emph{D. altispira} tests from the 1551-1554 m
(11.74 Ma) sample to determine representative inter-specimen variability
for these samples (Figure 3). Approximately one third (n=28) of the 100
depth profiles were excluded during screening for elevated Al/Ca and
Mn/Ca indicative of diagenetic contamination. The Mg/Ca value of
individual depth profiles in \emph{D. altispira} from the 1551-1554 m
sample ranges from 2.67 mmol/mol to 5.23 mmol/mol, with a mean of 3.63 \selectlanguage{ngerman}±
0.14 mmol/mol (n=72) (Figure 3a; Supplementary Table S5). The mean Mg/Ca
value from four specimens, a total of 28 profiles, is 3.41 ± 0.18
mmol/mol (Figure 3a). Averaging profiles from ten individual tests did
therefore not produce significantly better accuracy or precision than
averaging profiles from four individual tests (Figure 3b). Therefore,
for a Mg/Ca ratio to be considered representative it must represent an
average of at least 28 laser ablation profiles, from at least four
specimens, with the analytical uncertainty (2 SE) indicating the intra-
and inter-specimen variability this incorporates. To account for depth
profiles excluded due to contamination, where possible the number of
measurements per sample was increased to 36, six depth profiles per
specimen and six specimens per sample. This result is in line with other
LA-ICP-MS studies (\emph{Rathmann et al.} , 2004; \emph{Sadekov et al.}
, 2008). Future studies are advised to conduct similar testing to
determine the number of measurements required for a mean sample Mg/Ca to
be representative, as this will likely be site dependent.\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image3/image3}

\end{center}
\end{figure}

\textbf{Figure 3:} Distribution of \emph{D. altispira} Mg/Ca values from
LA-ICP-MS profiles of the 1551-1554 m sample. (A) A summary of all Mg/Ca
values, where open circles denote individual measurements, and filled
circles denote mean Mg/Ca values for each specimen. The horizontal black
line is the mean of all depth profiles from the sample, and the gray bar
the \selectlanguage{ngerman}±2 SE sample uncertainty. (B) The evolution of the sample 2 SE with
increasing specimens. Only profiles that passed data screening are
included (n=72). Data is provided in Supplementary Table S5.

2.7 Mg/Ca paleo-sea surface temperature calculations

The influence of calcification temperature (T) on the Mg/Ca ratio of
foraminiferal calcite can be explained by an exponential curve of
general form Mg/Ca = Bexp\textsuperscript{AT} where the pre-exponential
constant (B) and exponential constant (A) are species specific
(\emph{Anand et al.} , 2003; \emph{Lear et al.} , 2002; \emph{Nürnberg
et al.} , 1996; \emph{Rosenthal et al.} , 1997). To convert raw Mg/Ca
ratios to absolute temperatures, several secondary controls on Mg/Ca
must be considered, and accounted for (\emph{Gray et al.} , 2018;
\emph{Hollis et al.} , 2019; \emph{Holland et al.,} 2020).

In this study we use Mg/Ca values from \emph{D. altispira} , a near
surface dweller present from the Oligocene to the Pliocene. Since this
is not an extant species, we consider two approaches to calculating SST:
(i) using the multi-species calibration equation from \emph{Anand et
al.} (2003) and (ii) using the \emph{Globigerinoides ruber} Mg/Ca-SST
equation and pH correction from \emph{Evans et al.} 2016.

In scenario (i), we apply a compilation of nine modern planktic
foraminifera (\emph{Anand et al.,} 2003). This calibration is commonly
applied to extinct planktic foraminiferal species such as \emph{D.
altispira} and applies a power law relationship, where H is a constant
that describes the sensitivity of Mg/Ca\textsubscript{CALCITE} to
seawater Mg/Ca (Mg/Ca\textsubscript{SW}) (\emph{Hasiuk and
Lohmann,}2010; \emph{Cramer et al.,} 2011; \emph{Evans and Müller} ,
2012) (Equation 1).

\begin{equation}
\frac{\mathbf{Equation\ 1:\ \ \ Mg}}{\mathbf{\text{Ca}}}\mathbf{=}\frac{\mathbf{B}}{\frac{\mathbf{\text{Mg}}}{\mathbf{\text{Ca}}}_{\mathbf{\text{SW}}}^{\mathbf{t=0}^{\mathbf{H}}}}\mathbf{\text{\ x\ }}\frac{\mathbf{\text{Mg}}}{\mathbf{\text{Ca}}}_{\mathbf{\text{SW}}}^{\mathbf{t=t}^{\mathbf{H}}}\mathbf{\text{ex}}\mathbf{p}^{\mathbf{\text{AT}}}\nonumber \\
 \end{equation}

Fluxes of Mg\textsuperscript{2+} and Ca\textsuperscript{2+} into and out
of the oceans leads to secular variation in Mg/Ca\textsubscript{sw}.
This variability must be accounted for when determining absolute sea
surface temperatures on Cenozoic timescales (\emph{Hollis et al.} ,
2019). Reconstructions of Mg/Ca\textsubscript{sw} based on large benthic
foraminifera (\emph{Evans et al.} , 2018), calcite veins (\emph{Coggon
et al.} , 2010), fluid inclusions (\emph{Horita et al.} , 2002), and
echinoderms (\emph{Dickson} , 2002) have constrained this variability
through the Cenozoic (Supplementary Figure S4). The Eocene-Oligocene
demonstrates relatively stable values of 2.0-2.5 mol/mol (\emph{Coggon
et al.} , 2010; \emph{Evans et al.} , 2018). However, only one data
point exists from the Miocene, through which Mg/Ca\textsubscript{sw}
more than doubles from \textasciitilde{}2.2 mol/mol in the late
Oligocene (\emph{Coggon et al.} , 2010) to the well constrained value of
5.2 mol/mol in the modern ocean (\emph{Broecker et al.} ,
1982;\emph{Dickson} , 2002; \emph{Horita et al.} , 2002; \selectlanguage{polish}\emph{Kısakürek
et al.}\selectlanguage{ngerman} , 2008). Therefore, the method of \emph{Lear et al.} (2015) is
followed by fitting the fourth-order polynomial curve fit through the
compiled Mg/Ca\textsubscript{sw} proxy records (Supplementary Figure
S4). We use a ±0.5 mol/mol uncertainty window in the following
temperature calculations, this error envelope incorporating the majority
of the spread in the proxy data.

The power law function negates the assumption that the temperature
sensitivity remains constant, independent of changing
Mg/Ca\textsubscript{SW} through the Cenozoic era. We apply a power law
constant of H=0.41, similar to the value applied for \emph{T. trilobus}
, a symbiont-bearing, mixed layer dweller (\emph{Delaney et al.,}
1985;\emph{Evans and Müller} , 2012). Adapting Equation 1 to include our
H value, a modern-day Mg/Ca\textsubscript{sw} value of 5.2 mol/mol, and
the calibration constants of \emph{Anand et al.,} (2003) derives
Equation (2).

\begin{equation}
\mathbf{Equation\ 2:\ }\nonumber \\
 \end{equation}\begin{equation}
\frac{\mathbf{\text{Mg}}}{\mathbf{\text{Ca}}}\mathbf{=}\frac{\mathbf{0.38\pm 0.02}}{\mathbf{5.2}^{\mathbf{0.41}}}\mathbf{\text{\ x\ }}{\mathbf{Mg/Ca}_{\mathbf{\text{sw}}}}^{\mathbf{0.41}}\mathbf{\text{\ ex}}\mathbf{p}^{\mathbf{(0.090\pm 0.003\ x\ SST)}}\nonumber \\
 \end{equation}

This first calibration approach assumes that foraminiferal Mg/Ca is not
influenced by changes in the carbonate system. However, studies have
shown that planktic foraminiferal Mg/Ca is influenced by changes in the
carbonate system, the ratio increasing with decreased pH and/or
\selectlanguage{english}\selectlanguage{greek}Δ{[}CO\textsubscript{3}\textsuperscript{2-}{]} (\emph{Evans et al.} ,
2016; \emph{Gray and Evans} , 2019; \emph{Gray et al.} ,
2018;\emph{Russell et al.} , 2004; \emph{Yu and Elderfield} , 2008).
However, the ultimate driver of this effect is not certain and some
species are insensitive to changes in the carbonate system. Further, it
has been shown that for \emph{Orbulina universa} dissolved inorganic
carbon (DIC) plays a role in test Mg/Ca variability (\emph{Holland et
al.} , 2020). We follow recent results which interpret pH, as opposed to
\selectlanguage{english}\selectlanguage{greek}Δ{[}CO\textsubscript{3}\textsuperscript{2-}{]} or DIC, as the parameter
which controls the carbonate system's influence on Mg/Ca (\emph{Evans et
al.} , 2016; \emph{Gray et al.} , 2018). Furthermore, unlike with either
DIC or \selectlanguage{english}\selectlanguage{greek}Δ{[}CO\textsubscript{3}\textsuperscript{2-}{]}, it is possible to
reconstruct pH through the Neogene using boron isotopes in foraminifera
(\emph{Foster and Rae} , 2015; \emph{Greenop et al.} , 2014;
\emph{Henehan et al.} , 2013; \emph{Sosdian et al.} , 2018). For these
reasons we use the recent Neogene boron isotope compilation of
\emph{Sosdian et al.}(2018), which provides well constrained estimates
of pH across this time interval (Supplementary Figure S5; Supplementary
Table S9). Linear interpolation between these pH values allows us to
estimate a mean pH value, and associated uncertainty envelope, for each
Sunbird-1 sample, where the uncertainty envelope is maximum and minimum
pH at the 17\% and 83\% confidence interval (\textasciitilde{}\selectlanguage{ngerman}± 0.06 pH
units).

Therefore, in addition to scenario (i) we also consider the approach
from \emph{Evans et al} . (2016) which corrects for pH changes using the
interpolated Neogene pH record of \emph{Sosdian et al.} (2018)
(Supplementary Figure S5). Measured planktic foraminiferal Mg/Ca values
are corrected for this influence of pH using the equation of \emph{Evans
et al.} (2016) (Equation 3).

\begin{equation}
\frac{\mathbf{Equation\ 3:\ \ Mg}}{\mathbf{\text{Ca}}_{\mathbf{\text{CORRECTED}}}\mathbf{\ =\ }\frac{\frac{\mathbf{\text{Mg}}}{\mathbf{\text{Ca}}_{\mathbf{\text{MEASURED}}}}}{\frac{\mathbf{0.66}}{\mathbf{1+exp(6.9}\left(\mathbf{pH-8.0}\right)\mathbf{)}}\mathbf{+0.76}}}\nonumber \\
 \end{equation}

The preferred equation of \emph{Evans et al.} (2016) is used to account
for the influence of changing Mg/Ca\textsubscript{sw} when estimating
SST. These authors determined that the best fit to culture-derived
calibration lines is when both the pre-exponential (B) and exponential
(A) coefficients vary quadratically with
Mg/Ca\textsubscript{sw}(Equation 4 and 5).

Equation 4: B = (0.019 x Mg/Ca\textsubscript{sw}\textsuperscript{2}) --
(0.16 x Mg/Ca\textsubscript{sw}) + 0.804

Equation 5: A = (-0.0029 x Mg/Ca\textsubscript{sw}\textsuperscript{2}) +
(0.032 x Mg/Ca\textsubscript{sw})

We substitute these equations into the general exponential calibration,
Mg/Ca = Bexp\textsuperscript{AT}, to account for changing
Mg/Ca\textsubscript{sw}. Although the \emph{Evans} \emph{et al.} (2016)
equation is specific to \emph{G. ruber} , this species inhabits a
shallow water depth of 0-50m (\emph{Schiebel and Hemleben} , 2017)
similar to the inferred mixed-layer habitat depth \emph{D. altispira}
(\emph{Aze et al.} , 2011). Furthermore, as with \emph{G. ruber} ,
\emph{D. altispira}was a tropical/subtropical species, with symbionts
(\emph{Aze et al.} , 2011).

Salinity can exert a secondary effect on foraminiferal Mg/Ca,
sensitivity measurements from culture and core-top studies show this to
be \textasciitilde{}3-5\% per practical salinity unit (psu) (\emph{Gray
et al.} , 2018; \emph{Hollis et al.} , 2019; \selectlanguage{ngerman}\emph{Hönisch et al.} ,
2013;\selectlanguage{polish}\emph{Kısakürek et al.}\selectlanguage{ngerman} , 2008). In the absence of a robust,
independent salinity proxy (although we do note the promise of Na/Ca
(\emph{Bertlich et al.} , 2018; \emph{Geerken et al.} , 2018)) and the
relatively minor effect of salinity on foraminiferal Mg/Ca, this
potential secondary control is not empirically accounted for. Sunbird-1
was located in a coastal setting and likely experienced a highly
variable hydrological cycle due to changes in the position of the ITCZ
making it susceptible to changes in salinity. Therefore, an error of ±
0.5\selectlanguage{english}0C is incorporated into the final sea surface temperature estimates,
equivalent to an assumed salinity variability of \textasciitilde{}\selectlanguage{ngerman}± 1
PSU.

Mg/Ca-derived sea surface temperature estimates calculated using both
approaches (i) and (ii) yield extremely similar trends (Supplementary
Figure S6). Across the time interval of the Sunbird-1 dataset
(\textasciitilde{}13.5 Ma -- 9.5 Ma) pH changes by a small amount and
thus the choice of approach has little influence on the Sunbird-1
absolute SST record. In our discussion below, we adopt approach (i); the
multi-species calibration equation from \emph{Anand et al.} (2003)
without a pH correction. This approach avoids any potential
species-specific effects from applying the \emph{Evans et al.} (2016)
calibration specific to \emph{G. ruber} to the extinct \emph{D.
altispira} used in this study. Furthermore, \emph{D. altispira} has been
considered to be symbiont bearing, so may demonstrate a muted response
to changes in pH and insensitivity to pH changes, similar
to\emph{Trilobatus trilobus} (\emph{Gray and Evans} , 2019).

The uncertainties (± 2SE) associated with the conversion from Mg/Ca to
absolute SST estimates incorporate the uncertainty on the
Mg/Ca\textsubscript{sw} record, and the potential uncertainty due to
varying salinity. Additionally, scenario (i) incorporates the
uncertainty in the calibration of \emph{Anand et al.} (2003) (Equation
2), and scenario (ii) using the approach of \emph{Evans et al. (} 2016)
incorporates the uncertainty in the pH correction. These combined are
termed the calibration uncertainty and are considerably greater than the
independent analytical uncertainty, which only incorporates the intra-
and inter- specimen variability (±2 SE). Absolute sea surface
temperature estimates, and associated uncertainties, calculated using
approach (i) and (ii) are available in Table 1 and Supplementary Table
S9 respectively\selectlanguage{english}
\begin{longtable}[]{@{}llllllll@{}}
\toprule
Age (Ma) & Minimum Age (Ma) & Maximum Age (Ma) & Temperature (\selectlanguage{ngerman}°C) &
Maximum Temperature (\selectlanguage{ngerman}°C) & Minimum Temperature (\selectlanguage{ngerman}°C) & Analytical Error
Only Maximum Temperature (\selectlanguage{ngerman}°C) & Analytical Error Only Minimum
Temperature (\selectlanguage{ngerman}°C)\tabularnewline
\midrule
\endhead
9.53 & 9.43 & 9.62 & 27.73 & 31.08 & 24.64 & 28.30 &
27.12\tabularnewline
9.86 & 9.86 & 9.86 & 28.15 & 31.86 & 24.68 & 29.06 &
27.17\tabularnewline
10.19 & 10.05 & 10.33 & 29.54 & 33.49 & 25.81 & 30.62 &
28.34\tabularnewline
10.43 & 10.43 & 10.43 & 26.82 & 30.55 & 23.32 & 27.78 &
25.77\tabularnewline
10.48 & 10.48 & 10.48 & 28.13 & 31.58 & 24.94 & 28.77 &
27.44\tabularnewline
10.57 & 10.57 & 10.57 & 27.81 & 31.47 & 24.41 & 28.66 &
26.90\tabularnewline
10.62 & 10.62 & 10.62 & 24.88 & 28.12 & 21.90 & 25.43 &
24.30\tabularnewline
10.78 & 10.73 & 10.89 & 29.48 & 33.14 & 26.09 & 30.27 &
28.64\tabularnewline
10.92 & 10.92 & 10.92 & 28.95 & 32.78 & 25.36 & 29.92 &
27.89\tabularnewline
11.13 & 10.98 & 11.28 & 28.42 & 32.47 & 24.58 & 29.61 &
27.09\tabularnewline
11.40 & 11.40 & 11.40 & 28.18 & 31.50 & 25.15 & 28.67 &
27.68\tabularnewline
11.50 & 11.46 & 11.55 & 29.36 & 32.85 & 26.15 & 29.98 &
28.71\tabularnewline
11.61 & 11.10 & 11.61 & 26.78 & 30.39 & 23.44 & 27.59 &
25.91\tabularnewline
11.63 & 11.63 & 11.63 & 26.65 & 30.34 & 23.21 & 27.55 &
25.68\tabularnewline
11.64 & 11.64 & 11.64 & 29.57 & 33.29 & 26.11 & 30.40 &
28.68\tabularnewline
11.67 & 11.67 & 11.67 & 25.44 & 29.02 & 22.12 & 26.26 &
24.55\tabularnewline
11.72 & 11.69 & 11.74 & 28.19 & 31.84 & 24.81 & 28.99 &
27.33\tabularnewline
11.77 & 11.77 & 11.77 & 26.89 & 30.09 & 23.96 & 27.30 &
26.46\tabularnewline
11.82 & 11.82 & 11.82 & 26.39 & 29.69 & 23.37 & 26.92 &
25.85\tabularnewline
11.87 & 11.87 & 11.87 & 28.84 & 32.47 & 25.47 & 29.60 &
28.02\tabularnewline
12.03 & 12.03 & 12.03 & 28.10 & 31.66 & 24.82 & 28.81 &
27.35\tabularnewline
12.71 & 12.57 & 12.85 & 29.14 & 32.80 & 25.77 & 29.90 &
28.34\tabularnewline
13.23 & 13.13 & 13.33 & 28.85 & 32.54 & 25.43 & 29.64 &
28.00\tabularnewline
\bottomrule
\end{longtable}

\textbf{Table 1.} Sunbird-1 LA-ICP-MS Mg/Ca derived SST using the
approach of \emph{Anand et al.} (2003) without a pH correction. Minimum
and maximum age refer to the age range of the pooled samples
(Supplementary Table S7). Maximum and Minimum temperatures refer to the
full range of absolute temperatures derived incorporating the analytical
and calibration uncertainty, whereas Analytical Error Only Maximum and
Minimum temperatures refer to the range of temperatures derived from the
analytical uncertainty only.

2.8 \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O paleo-sea surface temperature calculations

Due to the limited sampling resolution of the trace metal data, SST is
also calculated using foraminiferal \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O.
Foraminiferal \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O
(\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{calcite}) is converted to
temperature (T) using the palaeotemperature equation of \emph{Bemis et
al.} (1998) (Equation 4), changes in global ice volume being corrected
using the \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{sw} value from the
nearest 0.1 Myr time interval in the compilation of \emph{Cramer et
al.}(2011).

\begin{equation}
\mathbf{Equation\ 4:\ \ }\left(\mathbf{\delta}^{\mathbf{18}}\mathbf{O}_{\mathbf{\text{calcite}}}\mathbf{\ \ }\mathbf{\delta}^{\mathbf{18}}\mathbf{O}_{\mathbf{\text{sw}}}\mathbf{\ +0.27}\right)\mathbf{=\ -0.21\ \pm 0.003\ T\ +\ 3.10\ \pm 0.07}\nonumber \\
 \end{equation}

The absence of a robust, independent salinity proxy makes any
quantitative attribution of its influence on foraminiferal
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O challenging. Therefore, we incorporate potential
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O variability due to salinity into any temperature
estimate uncertainty. Salinity of the upper water column in a 0.75\selectlanguage{ngerman}° x
0.75° grid square around the modern-day study site varies between 34.9
and 35.4 PSU (\emph{Boyer et al.} , 2013). Using the Indian Ocean
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{sw}-salinity relationship
of\emph{LeGrande and Schmidt} (2006) (Equation 5) this equates to a
maximum \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{sw} uncertainty of \selectlanguage{ngerman}±0.091\selectlanguage{english}%0.
Using Equation 4 this equates to a 0.4 \selectlanguage{english}0C uncertainty in the calculated
surface temperature.

\begin{equation}
\mathbf{Equation\ 5:\ \delta}^{\mathbf{18}}\mathbf{O}_{\mathbf{\text{sw}}}\left(\mathbf{\text{SMOW}}\right)\mathbf{=}\left(\mathbf{0.16\pm 0.004\ x\ Salinity}\right)\mathbf{-5.31\pm}\mathbf{0.135}\nonumber \\
 \end{equation}

We acknowledge the likelihood of variability in sea surface salinity in
this downcore record. We use the paleolatitude calculator of \emph{van
Hinsbergen et al.} (2015) to calculate a paleolatitude for Sunbird-1 at
10 Ma of approximately 5.5 \selectlanguage{ngerman}°S. The latitudinal correction of\emph{Zachos
et al.} (1994) gives a \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{sw} of 0.1%0.
The absence of a significant offset from SMOW (0%0) suggests that this
will have a negligible influence on the isotopic SST reconstructions.

3 Results

3.1 Solution ICP-MS trace element chemistry

\emph{D. altispira} Mg/Ca measured by solution ICP-MS ranges from 3.15 \selectlanguage{ngerman}±
0.1 to 40.2 ± 0.2 mmol/mol (Figure 4a), translating to unrealistically
high reconstructed sea surface temperatures. The high Mg/Ca ratios
strongly suggest the addition of magnesium from a secondary,
post-depositional source, prior to 11.75 Ma. The elevated Mg/Ca ratios
are associated with correspondingly high Mn/Ca, Al/Ca, and U/Ca (Figure
4b-d). Six of the sixteen Mn/Ca ratios are in excess of the proposed 200
\selectlanguage{greek}µ\selectlanguage{english}mol/mol threshold, from our LA-ICP-MS analysis, above which Mg/Ca
ratios are excluded due to contamination (Supplementary Figure S3).
Furthermore, every foraminiferal U/Ca ratio is considerably higher than
typical U/Ca ratios of primary foraminiferal calcite, which range from
\textasciitilde{}3-23 nmol/mol (\emph{Chen et al.} , 2017;
\emph{Raitzsch et al.} , 2011; \emph{Russell et al.} , 2004). In
addition, foraminiferal Al/Ca exceeds the commonly applied 100 \selectlanguage{greek}µ\selectlanguage{english}mol/mol
threshold in all but the four youngest samples.\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image4/image4}

\end{center}
\end{figure}

\textbf{Figure 4:} Downcore solution ICP-MS (a) Mg/Ca, (b) Mn/Ca, (C)
U/Ca, and (d) Al/Ca records for \emph{D. altispira} in the Sunbird-1
core, distinguishing between sample that were reductively cleaned (red
circles) and those that were not (blue squares).

The presence of elevated foraminiferal Mn/Ca, Al/Ca, and U/Ca ratios
does not necessarily mean that the Mg/Ca ratios are contaminated.
However, the downcore, point to point correlation (Figure 4) and
covariance (Supplementary Figure S7) between Mg/Ca and contaminant
indicators suggest a strong association. This downcore association
between Mg/Ca and contaminant indicators, despite a rigorous chemical
cleaning protocol, suggests one of two things; (i) the chemical cleaning
protocol is not fully effective at removing contaminant coatings, and/or
(ii) an Mg-rich contaminant phase is pervasive throughout the calcite
test.

Including the reductive cleaning step lowers Mg/Ca, Mn/Ca, and U/Ca
ratios in the post 11.8 Ma portion of the record, but has a negligible
effect on Al/Ca. Neither cleaning protocol is effective at removing the
authigenic coatings on the Sunbird-1 foraminifera in the pre 11.8 Ma
portion of the record (Figure 4). For this reason, we also analyzed
Sunbird-1 planktic foraminifera by laser ablation ICP-MS.

3.2 Downcore Laser Ablation ICP-MS Mg/Ca

Our laser ablation profiles clearly demonstrate that the metal-rich
contaminant is present as an authigenic surface coating on the glassy
foraminifera (e.g., Figure 2a-b). Because the alteration is not
pervasive throughout the calcite test, laser ablation ICP-MS is an ideal
approach to determine primary test Mg/Ca on these coated samples
(section 2.6). \emph{D. altispira} Mg/Ca determined by laser-ablation
ICP-MS ranges from 3.03 to 5.07 mmol/mol, with an average value of 4.18
\selectlanguage{ngerman}± 0.40 mmol/mol, and errors (±2SE) range from 0.10 to 1.04 mmol/mol
(Supplementary Table S6). However, due to elevated Al/Ca and Mn/Ca
ratios, only 14 of the 44 samples are represented by at least 28 laser
profiles. To alleviate this problem, adjacent samples have been combined
into longer time slices to ensure that the absolute mean Mg/Ca
measurements are robust (Supplementary Table S7). Samples comprising the
mean of at least 28 laser profiles are termed ``un-pooled samples''.
Samples pooled to achieve a minimum of 28 laser profiles are termed
``pooled samples''. It is acknowledged that combining adjacent samples,
which span up to 420 kyr, could incorporate orbital scale climatic
variability into these pooled samples. However, we do not infer climatic
variability on orbital timescales because the coarse sampling resolution
could incorporate aliasing of any precessional or obliquital periodicity
into longer term eccentricity cycles (\emph{Pisias and Mix} , 1988).
Combining adjacent samples to generate a representative mean Mg/Ca for a
longer time-slice could smooth orbital scale variability, could reduce
uncertainty and assist the interpretation of longer-term climatic
trends.

The mean Mg/Ca of representative samples after incorporating the nine
pooled Mg/Ca samples with the 14 un-pooled samples ranges from 3.08 to
4.70 mmol/mol, with an average value of 4.04 ± 0.29 mmol/mol, and errors
(±2SE) range from 0.14 to 0.48 mmol/mol (Supplementary Table S8). These
values are in good agreement with the reductively cleaned solution
ICP-MS data for the post-11.8 Ma portion of the record (Supplementary
Figure S2), coinciding with the interval when contaminant indicators
(Mn/Ca, Al/Ca, and U/Ca) are substantially lower (Figure 4b-d). This
agreement between Mg/Ca values obtained by LA-ICP-MS and solution ICP-MS
following effective reductive-cleaning supports the suitability of the
LA-ICP-MS analyses. Because we can be more confident that the laser
ablation data are not biased by authigenic coatings, the laser-ablation
approach has the advantage that we can also determine original test
Mg/Ca in the older part of the record.

There is no obvious long-term trend in Mg/Ca through the interval
(Figure 5a). Between 11.8 Ma and 11.7 Ma there is a 0.7-0.8 mmol/mol
decrease in Mg/Ca followed by a recovery to approximately previous
values at 11.5-11.4 Ma. There is a Mg/Ca decrease of similar magnitude
from between 10.7 Ma and 10.36 Ma, recovering by 9.85 Ma. We acknowledge
that the coarse sampling frequency, and the combining of samples could
be obscuring similar variability through the rest of the record.

3.3 \emph{G. obliquus} \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O

\emph{G. obliquus} \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O ranges from -3.63%0 to -2.34%0
with a mean value of -2.92%0. The \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O record shows
very little variability, values remaining stable at -3.4%0 prior to a
positive 0.6%0 shift at \textasciitilde{}12.5 Ma, and -2.7%0 after (Figure
5b). The low variability translates to a stable \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O
SST record, temperatures ranging between 27\selectlanguage{ngerman}°C and 31°C with the only
distinctive trend being a \textasciitilde{}3°C decrease between
\textasciitilde{}12.7 Ma and 12.0 Ma. The coeval positive 0.3\selectlanguage{english}%0 shift in
seawater \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O (\emph{Cramer et al.} , 2011) dampens
the influence on the SST estimate of the positive 0.6%0 shift in \emph{G.
obliquus} \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O at \textasciitilde{}12.5 Ma.\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image5/image5}

\end{center}
\end{figure}\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image6/image6}

\end{center}
\end{figure}

\textbf{Figure 5:} (a) Mean \emph{D. altispira} LA-ICP-MS Mg/Ca ratios
(mmol/mol) for unpooled (black squares) and pooled (grey squares)
samples from Sunbird-1. Error bars denote the age range for pooled
samples, and the \selectlanguage{ngerman}± 2SE of Mg/Ca from all depth profiles in the sample.
(b) \emph{G. obliquus} \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O from Sunbird-1. Solid line
is a five-point moving average. (c) Sea surface temperature records at
Sunbird-1 from planktic foraminiferal \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O, and
LA-ICP-MS Mg/Ca using our preferred approach that applies the
calibration of \emph{Anand et al.} (2003) without a pH correction.
Symbols are the same as in (a) and (b). Error bars on the
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O record denote the analytical uncertainty (\selectlanguage{ngerman}± 2SD),
and error bars on the Mg/Ca record denote the sample uncertainty (±
2SE). As in (a), pooled Mg/Ca samples also have horizontal error bars
denoting the sample age range. Dashed blue and black lines denote the
full uncertainty on the temperature estimates, including that derived
from the calibration uncertainty, for \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O and
LA-ICP-MS Mg/Ca respectively. Supplementary Figure S8 provides LA-ICP-MS
Mg/Ca sea surface temperatures using the alternative approach
of\emph{Evans et al.} , (2016).

4 Discussion

4.1 Reconstructing sea surface temperature from diagenetically altered
foraminifera using laser ablation ICP-MS

Robust paleotemperature reconstructions using foraminiferal Mg/Ca ratios
are reliant upon the Mg/Ca ratio recording a primary environmental
signal, unaltered by diagenetic alteration. Despite employing a thorough
cleaning protocol (\emph{Barker et al.} , 2003; \emph{Boyle and Keigwin}
, 1985), our Mg/Ca ratios from solution-based ICP-MS analysis in the
\textgreater{}11.8 Ma portion of the record are clearly influenced by a
diagenetic contaminant phase containing elevated magnesium (Figure 4).
This finding demonstrates that foraminifera with a glassy appearance
under the light microscope are not necessarily free from the influence
of all modes of diagenetic alteration. We therefore emphasize the
importance of complementary trace metal ratios indicative of
contamination (i.e. Al/Ca, Mn/Ca, U/Ca) to assess the reliability of
foraminiferal Mg/Ca ratios (Figure 4). The application of LA-ICP-MS to
collect high resolution elemental profiles through the foraminiferal
tests, excluding regions displaying diagenetic contamination, has
facilitated the identification of what we interpret to be primary
paleotemperatures from diagenetically altered foraminifera (\emph{Hines
et al.} , 2017; \emph{Hollis et al.} , 2015).

The Sunbird-1 \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{PF} SST record
from\emph{G. obliquus} reconstructs very similar absolute temperatures
to the planktic foraminiferal Mg/Ca SST record (Figure 5c). Mean SST
from the Sunbird-1 \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{PF} record
(290\selectlanguage{english}C) is 20\selectlanguage{english}C higher than mean SST from the Mg/Ca record (270\selectlanguage{english}C),
although with the exception of the two transient decreases in Mg/Ca
reconstructed SST initiating at 11.8 Ma and 10.7 Ma the records are
within error. The similarity of the absolute SSTs reconstructed by the
two proxies strengthens the case for the LA-ICP-MS Mg/Ca SST record
recording a primary temperature signal, and that these absolute sea
surface temperatures at Sunbird-1 should be considered primary.

The majority of the uncertainty in the absolute temperature estimates is
derived from the uncertainties incorporated from the relevant
calibrations, in particular the seawater Mg/Ca and seawater
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O records (Figure 5c). This is true for both
LA-ICP-MS Mg/Ca (Table 1 and Supplementary Table S9) and
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O (Supplementary Table S10). Therefore, despite
being appreciable, the uncertainty resulting from the geochemical
heterogeneity both within an individual foraminiferal test and between
foraminiferal tests from the same sample (Figure 3) is not the primary
contributor to the final absolute temperature uncertainty.

4.2 Mid-late Miocene sea surface temperatures in the equatorial Indian
Ocean

The results from Sunbird-1 indicate that SST in the equatorial Indian
Ocean remained stable at \textasciitilde{}270\selectlanguage{english}C-290\selectlanguage{english}C through the 13.3 Ma
to 9.5 Ma interval (Figure 5c). This finding suggests that tropical
climate was relatively stable following the global cooling associated
with the expansion of the East Antarctic Ice Sheet across the MMCT.
These records from Sunbird-1 supports the robustness of contemporaneous
alkenone based studies which exhibit similar absolute tropical SST
estimates (\emph{Herbert et al.} , 2016; \emph{Huang et al.} ,
2007;\emph{Rousselle et al.} , 2013; \emph{Seki et al.} , 2012;
\emph{Zhang et al.} , 2014) (Figure 6a). The
U\textsuperscript{k}\textsubscript{37} SST calibration fails to
reconstruct SST\textgreater{}290\selectlanguage{english}C (\selectlanguage{ngerman}\emph{Müller et al.} , 1998) but
these results using Mg/Ca paleo-thermometry suggest that outside the
western Pacific warm pool this restriction does not apply to this time
interval, unlike the preceding Miocene Climatic Optimum during which
Mg/Ca temperature estimates are higher than those estimated with the
U\textsuperscript{k}\textsubscript{37} proxy (\emph{Badger et al.} ,
2013).\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image7/image7}

\end{center}
\end{figure}

\textbf{Figure 6:} Sunbird-1 LA-ICP-MS Mg/Ca derived SST using the
approach of \emph{Anand et al.} (2003) without a pH correction compared
to and SST estimates at contemporaneous sites from (a)
U\textsuperscript{k}\textsubscript{37}, and (b) foraminiferal
geochemistry. Estimates applying
U\textsuperscript{k}\textsubscript{37}are from ODP Site 722 (\emph{Huang
et al.,} 2007) in the Arabian Sea, ODP \& IODP Sites 846 (\emph{Herbert
et al.,} 2016), 850 (\emph{Zhang et al.,} 2014), 1241 (\emph{Seki et
al.,} 2012), and U1338 (\emph{Rousselle et} \emph{al.,} 2013) in the
Eastern Equatorial Pacific, terrestrial outcrops in Malta (\emph{Badger
et al.,} 2013). Estimates applying the foraminiferal Mg/Ca proxy are
from ODP Sites 761 (\emph{Sosdian and Lear,} 2020) and terrestrial
outcrops in Malta (\emph{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 \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O of exceptionally preserved foraminifera from
Tanzania are also shown (\emph{Stewart et al.,} 2004). The upper limit
for the U\textsuperscript{k}\textsubscript{37} proxy (290\selectlanguage{english}C) 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 \emph{Evans et al.} , (2016).

Although not a true tropical location, and consisting of only three data
points, the \emph{Badger et al.} (2013) Mg/Ca record from the
Mediterranean estimates SST of \textasciitilde{}27.50\selectlanguage{english}C 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 (\emph{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 270\selectlanguage{english}C at 12.2 Ma and 290\selectlanguage{english}C at 11.55 Ma
using the \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O paleo-thermometer (\emph{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 \emph{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
TEX\textsubscript{86} paleothermometer suggest a slight,
\textasciitilde{}1\selectlanguage{ngerman}°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 (\emph{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 \textasciitilde{}3\selectlanguage{english}0C. Excluding one
value of 28.60\selectlanguage{english}C at 11.62 Ma, this decrease in SST to
\textasciitilde{}24-250\selectlanguage{english}C persists for \textasciitilde{}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
U\textsuperscript{k}\textsubscript{37} proxy from the Arabian Sea
(\emph{Huang et al.} , 2007), and the Eastern Equatorial Pacific
(\emph{Herbert et al.} , 2016; \emph{Rousselle et al.} , 2013;
\emph{Seki et al.} , 2012; \emph{Zhang et al.} , 2014) (Figure 6a). We
therefore suggest that the observed transient \textasciitilde{}30\selectlanguage{english}C 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 \textasciitilde{}11.8 and 11.5 Ma. Assuming
constant SST, the observed \textasciitilde{}0.7 mmol/mol decrease in
Mg/Ca would require a salinity increase on the order of 5.0 PSU
(\selectlanguage{ngerman}\emph{Hönisch et al.} , 2013; \emph{Gray et al.} , 2018). This salinity
increase equates to a 0.8 \selectlanguage{english}%0 change in \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O using the
Indian Ocean \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O\textsubscript{sw}-salinity
relationship of \emph{LeGrande and Schmidt} (2006) (Equation 5). As well
as being an extremely large change in salinity, the planktic
foraminiferal \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O record does not support such a
significant change in sea surface salinity between \textasciitilde{}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 (\emph{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
\textasciitilde{}3\selectlanguage{ngerman}°C lowering of SST between 11.8 Ma and 11.5 Ma.
However, mean \emph{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
\textasciitilde{}11.8 and 11.5 Ma. The lack of a marked increase in the
planktic \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O 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
U\textsuperscript{k}\textsubscript{37}proxy 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
\textasciitilde{}7 Ma (\emph{Herbert et al.} , 2016;\emph{LaRiviere et
al.} , 2012). The absolute tropical SST record reported in this study
supports the finding that the latitudinal temperature gradient steepened
from \textasciitilde{}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 (\emph{Cramwinckel et al.} , 2018). There is little evidence
for a significant change in pCO\textsubscript{2} in this interval
(\emph{Sosdian et al.} , 2018;\emph{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 CO\textsubscript{2} drawdown that was
associated with the subsequent global late Miocene Cooling starting
\textasciitilde{}7.5 Ma (Figure 7).\selectlanguage{english}
\begin{figure}[H]
\begin{center}
\includegraphics[width=0.70\columnwidth]{figures/image8/image8}

\end{center}
\end{figure}

\textbf{Figure 7} : Summary of global climate through the mid-to-late
Miocene. (a) Sea surface temperature estimates from Sunbird-1, fellow
low latitude ODP sites 850 (\emph{Zhang et al.} , 2014) and 761
(\emph{Sosdian and Lear,} 2020), mid latitude Northern Hemisphere ODP
site 1021 (\emph{LaRiviere et al.,} 2012), and mid-latitude Southern
Hemisphere site 1125 (\emph{Herbert et al.,} 2016), and high-latitude
Northern Hemisphere ODP Site 982 (\emph{Herbert et al.} , 2016). ODP
Site 761 data is displayed on an alternative axis as SST anomalies
relative to the baseline average from 16.0 -- 15.5 Ma. (b)
pCO\textsubscript{2}reconstructions, with Y axis on a log scale, of
\emph{Sosdian et al.}(2018) applying the CCD reconstruction of
\selectlanguage{ngerman}\emph{Pälike et al.} (2012) and the
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{11}B\textsubscript{SW} scenario of\emph{Greenop et
al.} (2017), and \emph{Stoll et al.} (2019) applying temperature
estimates from \emph{Bolton et al.} (2016) and \emph{Zhang et al.}
(2013). Confidence intervals (95\%) are displayed as dashed lines and
error bars respectively. (c) Composite benthic \selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O
record showing data that have been smoothed by a locally weighted
function over 20 kyr (blue curve) and 1 Myr (red curve)
(\emph{Westerhold et al.} , 2020). Blue, yellow, and gray panels
indicate intervals of ice sheet expansion across the Mid Miocene Climate
Transition (MMCT) associated with CO\textsubscript{2} decline, the
steepening of latitudinal temperature gradeints in the absence of a
CO\textsubscript{2} trend, and the Late Miocene Cooling (LMC).

5 Conclusions

Our Sunbird-1 sea surface temperature estimates from LA-ICP-MS Mg/Ca
analyses are in good agreement with those using the
\selectlanguage{greek}δ\selectlanguage{english}\textsuperscript{18}O paleo-thermometer on glassy foraminifera,
supporting the use of LA-ICP-MS micro-analysis across multiple specimens
for reconstructing paleotemperatures. This analytical technique has
allowed the reconstruction of reliable Mg/Ca derived paleotemperatures
using foraminifera whose bulk trace element ratios demonstrate
diagenetic contamination by authigenic coatings. This finding opens the
potential for Mg/Ca paleothermometry on other challenging time
intervals, and locations, where contaminant coatings have previously
inhibited the geochemical analysis of primary foraminiferal calcite. We
present new sea surface temperature records from planktic foraminiferal
Mg/Ca for the south west Indian Ocean between 13.5 Ma and 9.5 Ma.
Absolute estimates of 24-310\selectlanguage{english}C suggest that sea surface temperature was
relatively constant through the interval, although our record also
suggests two intervals of regional cooling and freshening of surface
waters at 11.8 and 10.7 Ma. The late Miocene represented a key interval
in the transition of Earth's climate to its modern state, including the
development of stronger latitudinal temperature gradients. Our new
temperature record suggests that different mechanisms may have been
responsible for this cooling. The initial cooling from
\textasciitilde{}10 Ma at mid to high latitudes in both hemispheres was
not associated with significant cooling at low latitudes. On the other
hand, the late Miocene cooling between \textasciitilde{}7.5 and 5.5 Ma
was global in nature and associated with a drawdown in
pCO\textsubscript{2}. Further work should therefore explore the
mechanisms responsible for the enhanced polar amplification between 10
and 7.5 Ma, and the possibility of carbon cycle feedbacks contributing
to the subsequent late Miocene Cooling.

Acknowledgments, Samples, and Data

This study uses samples from the Sunbird-1 core provided by BG-Group.
All data from this study can be found in Table 1 and Supplementary
Tables S1 to S11,and aredeposited in the Zenodo online data repository
http://doi.org/10.5281/zenodo.4472994. We thank Alexandra Nederbragt and
Anabel Morte-Rode\selectlanguage{ngerman}ñas for laboratory assistance. We thank the reviewers
and editor for their insightful comments that improved the manuscript.
This research was supported by NERC iCASE studentship BW/22003105
(M.G.N.), and NE/L009633/1 grant to C.H.L.

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