# Last deglacial environmental change in the Tropical South Pacific from Tahiti Corals

Oliver Knebel, Thomas Felis, Ryuji Asami, Pierre Deschamps, Martin Kölling, Denis Scholz

► To cite this version:

Oliver Knebel, Thomas Felis, Ryuji Asami, Pierre Deschamps, Martin Kölling, et al. Last deglacial environmental change in the Tropical South Pacific from Tahiti Corals. Paleoceanography and Paleoclimatology, 2024, 39 (2), e2022PA004585 [21 p.]. 10.1029/2022PA004585. hal-04731044

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# Paleoceanography and
Paleoclimatology®

## RESEARCH ARTICLE

10.1029/2022PA004585

**Key Points:**

* Last deglacial changes in Tahiti coral Sr/Ca suggest an influence of Northern-Hemisphere climate events on the tropical South Pacific

* Variations in last deglacial Tahitian climate were modulated by meridional migrations of the South Pacific Convergence Zone

* Last deglacial *Porites* corals revealed reduced linear extension rates (1.0 ± 0.6 cm year⁻¹) compared to today (1.6 ± 0.3 cm year⁻¹)

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

O. Knebel,
oknebel@marum.de

**Citation:**

Knebel, O., Felis, T., Asami, R., Deschamps, P., Kölling, M., & Scholz, D. (2024). Last deglacial environmental change in the tropical South Pacific from Tahiti corals. *Paleoceanography and Paleoclimatology, 39*, e2022PA004585. https://doi.org/10.1029/2022PA004585

Received 23 NOV 2022
Accepted 18 JAN 2024

**Author Contributions:**

**Conceptualization:** Thomas Felis

**Data curation:** Oliver Knebel

**Formal analysis:** Oliver Knebel

**Funding acquisition:** Thomas Felis

**Investigation:** Oliver Knebel, Thomas Felis, Ryuji Asami, Pierre Deschamps, Martin Kölling, Denis Scholz

**Project administration:** Thomas Felis

**Supervision:** Thomas Felis

**Visualization:** Oliver Knebel

**Writing – original draft:** Oliver Knebel

## Last Deglacial Environmental Change in the Tropical South Pacific From Tahiti Corals

<figure><img src="image_3.png" /></figure>

Oliver Knebel¹ ID, Thomas Felis¹ ID, Ryuji Asami² ID, Pierre Deschamps³ ID, Martin Kölling¹ ID, and Denis Scholz⁴ ID

¹MARUM - Center for Marine Environmental Sciences, University of Bremen, Bremen, Germany, ²Graduate School of Science, Institute of Geology and Paleontology, Tohoku University, 6-3 Aramaki-aza-Aoba, Sendai, Japan, ³Aix-Marseille Universités, CNRS, IRD, INRAE, CEREGE, Aix-en-Provence, France, ⁴Institute for Geosciences, Johannes Gutenberg-University Mainz, Mainz, Germany

**Abstract** On glacial-interglacial time scales, changes in the Earth's orbital configuration control climate seasonality and mean conditions. Tropical coral skeletons can be sampled at a sufficient resolution to reconstruct past seasonality. Here, last deglacial *Porites* skeletons from Integrated Ocean Drilling Program Expedition 310 to Tahiti are investigated and, supported by a modern calibration, monthly resolved time series in geochemical proxies (Sr/Ca, δ¹⁸O, δ¹³C) are constructed. For most of the deglaciation, Sr/Ca seasonality was similar to modern (0.139 ± 0.010 mmol mol⁻¹; 2.8 ± 0.2°C) reflecting the small change in insolation seasonality. However, during the Younger Dryas, high values in Sr/Ca seasonality (0.171 ± 0.017 mmol mol⁻¹; 3.4 ± 0.3°C) suggest a reduced mixed layer depth and enhanced influence of the South Pacific Subtropical Gyre due to South Pacific Convergence Zone (SPCZ) inactivity. Furthermore, high amplitudes in Younger Dryas skeletal δ¹⁸O (0.40 ± 0.22 ‰) and δ¹³C (0.86 ± 0.22 ‰) seasonality compared to modern (δ¹⁸O = 0.29 ± 0.08 ‰; δ¹³C = 0.27 ± 0.08 ‰) point to elevated winter-summer discrepancies in rainfall and runoff. Mean coral Sr/Ca variability suggests an influence of Northern Hemisphere climate events, such as the Younger Dryas cooling (+0.134 ± 0.012 mmol mol⁻¹; -2.6 ± 0.2°C), or the Bølling-Allerød warming (+0.032 ± 0.040 mmol mol⁻¹; -0.6 ± 0.4°C). Deglacial mean coral Δδ¹⁸O (δ¹⁸O₃ seawater contribution to skeletal δ¹⁸O), corrected for the ice volume effect, was elevated pointing to more saline, thus dryer conditions, likely due to a northward migration of the SPCZ. Seasonal cycles in coral δ¹³C were likely caused by variations in linear extension rates that were reduced during the last deglaciation (1.00 ± 0.6 cm year⁻¹) compared to today (1.6 ± 0.3 cm year⁻¹).

**Plain Language Summary** Changes in the Earth's movement around the sun are the primary cause of changes in the annual cycle and annual mean of environmental conditions during the last glacial-interglacial periods. Distinct from most climate archives, the annual bands of coral skeletons can be sampled at a monthly resolution that allows for reconstructions of past seasonal cycles. In this study, the chemical composition of coral skeletons from the genus *Porites* that grew during the last deglaciation at Tahiti (~15,000–9,000 years ago) was analyzed to infer seasonal variations and mean conditions in past sea surface temperature and salinity. Prior to application, the methods were successfully calibrated using modern coral skeletons from Tahiti. Results showed that sea surface temperature seasonality was similar to today during most of the deglaciation. However, for the Younger Dryas, a particular cold interval, a marked increase in sea surface temperature seasonality suggests that Tahiti was at that time located outside the South Pacific Convergence Zone, a rain band extending across the tropical Southwest Pacific. Furthermore, coral reconstructions revealed that, although distant, North Atlantic last deglacial climate events influenced mean climate conditions at Tahiti. In addition, we show that last-deglacial coral growth rates were reduced compared to today.

## 1. Introduction

Global climate variability on a glacial-interglacial time scale is primarily controlled by variations in the earth's orbital parameters. According to the Milankovitch hypothesis, the transition from glacial to interglacial climate states is initiated by an increase in the Northern Hemisphere summer insolation, while total annual insolation remains nearly constant (Berger, 1980). Although insolation seasonality is a fundamental driver of past climate variability, there is only little paleoclimatological proxy data with sufficient resolution to resolve climate variations on the sub-annual scale and reconstruct the amplitude of annual cycles of the past (Brummer et al., 2020).
For instance, sediment cores covering the last deglacial transition in the central Pacific are sparse and exhibit very low sedimentation rates of ~1 cm ka⁻¹ (Seo et al., 2016). In contrast, tropical corals calcify at rates of ~1 cm year⁻¹ and record ambient environmental conditions in the geochemical composition of their aragonitic annual bands (Goreau, 1977; Knutson et al., 1972). The Sr/Ca of coral carbonate is an established proxy for sea surface temperature (SST; Le Cornec & Correge, 1997; Smith et al., 1979) and can be sampled at a sub-seasonal resolution that allows for reconstructions in paleo-SST seasonality (Asami et al., 2020; Brocas et al., 2016; Felis et al., 2015; Giry et al., 2012). Additionally, the oxygen isotope values (δ¹⁸O) and carbon isotope values (δ¹³C) are geochemical proxies routinely measured in coral skeletons. Coral δ¹⁸O is a function of both SST and seawater δ¹⁸O which relates to sea surface salinity (SSS; Hendy et al., 2002; Quinn et al., 1993; Weber & Woodhead, 1972). The contribution of seawater δ¹⁸O to the coral skeletal δ¹⁸O signal (Δδ¹⁸O) can be derived from paired δ¹⁸O-Sr/Ca analysis (Cahyarini et al., 2008). Coral δ¹³C is controlled by metabolic processes related to photosynthesis and respiration, kinetic effects related to calcification rate, and the state of the ambient seawater carbonate system (Felis et al., 2022; Linsley et al., 2019).

While the precessional cycle is the component of the orbital forcing mainly driving changes in climate seasonality, changes in the obliquity cycle govern variations in climate mean state (Timmermann, Lorenz, et al., 2007). Such orbital forcing initiated the early last deglacial warming but a range of other processes contributed to the continuing warming trend until the Early Holocene, including changes in ice sheet extent (Clark et al., 2009), sea level (De Nezio et al., 2016), and greenhouse gas forcing (Shakun et al., 2012). The deglacial warming trend was interrupted by meltwater-forced slow-downs or even a temporal shutdown of the Atlantic Meridional Overturning Circulation (AMOC) that led to deglacial cooling events, such as Heinrich Stadial 1 and the Younger Dryas (McManus et al., 2004; Ng et al., 2018; Praetorius et al., 2020). To which extent and pathways these North Atlantic deglacial cooling events affected other basins, such as the tropical South Pacific, is a matter of debate. It is likely that oceanic and, more importantly, atmospheric teleconnections, through a reorganization of the low latitude large-scale atmospheric circulations, influenced Pacific climate conditions (Okumura et al., 2009; Timmermann, Okumura, et al., 2007).

In this study, *Porites* coral skeletons recovered during the Integrated Ocean Drilling Program (IODP) Expedition 310 off Tahiti are analyzed for Sr/Ca, δ¹⁸O, and δ¹³C aiming to reconstruct central tropical South Pacific seasonality and mean environmental conditions, especially SST, for key periods of the last deglacial transition (Heinrich Stadial 1, Bølling-Allerød, Melt-Water Pulse 1A, Younger Dryas, Early Holocene). Samples have been dated using the ²³⁰Th-U method (Andersen et al., 2008; Edwards et al., 1986) and revealed ages of 15.15–9.52 ka BP (Deschamps et al., 2012) with most being deposited around Meltwater Pulse 1A (MWP1A; 14.65–14.34 ka BP). Furthermore, five modern corals collected from the vicinity of IODP Expedition 310 drilling sites are geochemically analyzed and serve as a benchmark for present-day conditions.

The results of this study provide rare proxy data of tropical Pacific environmental conditions on a sub-seasonal timescale for the last deglacial transition. These data are highly relevant for comparison with climate model simulations and improve our understanding of the controls in environmental seasonality, as well as the response in climate variability and global teleconnections to abrupt perturbations. Also today, under ongoing climate change, the interrelation between changes in mean conditions, seasonality, and other compartments of the climate system, such as the El Niño–Southern Oscillation (ENSO), are complex and not fully understood, especially in the tropics. Consequently, the results of this study of the last deglacial warming in the tropical South Pacific also contribute to improving projections for future tropical Pacific climate.

## 2. Materials and Methods

### 2.1. Study Area

Tahiti (French Polynesia) is a volcanic island with a pronounced topography of up to 2241 masl (meters above sea level), located in the central tropical South Pacific (S17°41′ W149°27′; Figure 1; Pouteau et al., 2011). Total irradiance largely follows the seasonal insolation cycle (Figure 2b) and is highest in austral summer (February) with 206.4 ± 25.0 W m⁻² (SD, standard deviation; if not mentioned otherwise, uncertainty in this study is indicated at 1σ level) and lowest in winter (June) with 130.5 ± 11.3 W m⁻² (1960–1998; Cayan, 1992). The seasonal SST cycle follows the total irradiance with a lag of 1–4 months (Figures 2a and 2b). In situ SST measurements from Tahiti (IRD, 1979–1992; Figure 1) and the neighboring island Moorea (CRIOBE, 1992–2018; Figure 1) have shown that maximal mean monthly SST was 28.7 ± 0.5°C and occurred 55.0% of these years
<figure><img src="image_4.png" /><figcaption>**Figure 1.** Satellite imagery of Tahiti and Moorea islands (French Polynesia) showing drilling locations of IODP Expedition 310 (Camoin et al., 2007), modern corals, as well as locations of in situ measurements of seawater temperature. Measurements at Papeete and Tiahura were performed by IRD and CRIOBE respectively.</figcaption></figure>

in March, 27.5% in April, 15.0% in February, and 2.5% in May (Figure 3e). Similarly, minimal mean monthly SST is 26.5 ± 0.5°C and occurred in 47.5% of the years in August, 32.5% in September, 10.0% in July, and 5.0% in January and October (Figure 3f). The mean annual SST at Tahiti is 27.6 ± 0.4°C. The regional climate is dominated by the trade winds that blow from the Northeast during summer and from the Southeast during winter (Bessemoulin et al., 2004). Consequently, rainfall occurs predominantly on the eastern side of the island with an annual average of 1876 ± 236 mm year⁻¹, (1979–1991; Gerard, 1992; SD). Tahiti is located within the SPCZ

<figure><img src="image_5.png" /><figcaption>**Figure 2.** Mean seasonal cycles for coral colonies. The seasonality of the OISST v2.1 (Huang et al., 2020) in (a) is denoted in red, the seasonality (weighted mean) of modern records in (b) is denoted in black, and the inter- and intra-colony variability of the seasonality in (a) and (c) is denoted in gray. In (a) record TIAREI-1 (black line), TIAREI 3B (gray), TIAREI-4 (purple), MARAA-7 (green), and MARAA-8 (orange) are shown. In (b) insolation at the latitude of Tahiti (17.67° S) from Berger (1978) and the irradiance at Tahiti from Cayan (1992) are shown. In (c) the seasonal cycles from annual bands of similar age from fossil record 310-20A-23R-1 (0–78 cm) are shown to illustrate the intra-colony variability.</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>**Figure 3.** (a) Sr/Ca time series from modern corals with (b) applied offsets. (c) Weighted least squares regression between composite Sr/Ca time series and OISST v2.1 SST. (d) Time series of in situ measurements of SST at Tahiti and Moorea (gray and black), OISST v2.1 SST data from the same period (red), and SST reconstructed from composite Sr/Ca (blue) using the slope from (c). Furthermore, histograms show the frequency of the warmest (e) and (f) coldest months of each year from the in situ data, as well as (g) rainfall and SSS from in situ measurements and data products. The SSS seasonality from SODA is 0.39 PSU (Carton & Giese, 2008), from IRD is 0.42 PSU, and from Delcroix et al. (2011) is 0.28 PSU.</figcaption></figure>

(South Pacific Convergence Zone) that extends south-eastwards from the equatorial West Pacific Warm Pool toward the tropical South Pacific (Brown et al., 2020). Environmental conditions within the SPCZ are characterized by high rainfall rates and elevated SST. The mixed layer depth is enhanced in the SPCZ compared to the South Pacific Subtropical Gyre (de Boyer Montégut et al., 2004) and, at Tahiti, varies seasonally between 34 and 98 m (monthly mean; Sauzède et al., 2020). Rainfall at Tahiti is highest in summer (December) with a monthly average of 384 ± 186 mm, when SPCZ activity reaches the maximum, and lowest in winter (June) with 47 ± 26 mm (Figure 3g). On a seasonal scale, SSS follows the pattern in rainfall and varies between 35.8 ± 0.3 PSU in April and 36.0 ± 0.2 PSU in August (1985–2009; Delcroix et al., 2011; Figure 3g). Previous studies have shown that the annual variability in SSS at Tahiti as observed today is too small to be resolved from paired coral δ¹⁸O-Sr/Ca analysis (Cahyarini et al., 2008). Interannual climate variability in the tropical South Pacific is largely determined by ENSO and the associated meridional migration of SPCZ (Gouriou & Delcroix, 2002). However, at Tahiti, the influence of ENSO on interannual SST is only small, because Tahiti is located near the nodal line of ENSO-induced SST anomalies, resulting in anomalies close to zero during El Niño and La Niña events (Felis et al., 2012).

## 2.2. Coral Samples

Samples analyzed in this study (Table 1) were collected during the IODP Expedition 310 off Tahiti in 2005 (Camoin et al., 2007). Deglacial sequences containing *Porites* skeletons were recovered in three areas (Tiarei, Faaa, Maraa) by drilling into reef terraces at approximately 40–150 mbsl (meters below sea level). Previous studies have utilized some of these samples for sea level (Deschamps et al., 2012; Thomas et al., 2012) and
<table><caption>**Table 1**

Overview of Modern and IODP Expedition 310 Fossil Tahiti Coral Samples Used in This Study</caption><thead><tr><th>Sample ID</th><th>Location</th><th>Annual cycles (n)</th><th>Sampling resolution (samples year<sup>-1</sup>)</th><th>Annual linear extension ± 1SD (cm year<sup>-1</sup>)</th><th>Sr/Ca seasonality ± 1SE (mmol mol<sup>-1</sup>)</th><th>Mean Sr/Ca ± 1SE (mmol mol<sup>-1</sup>)</th><th>Age ± 2SD</th></tr></thead><tbody><tr><td colspan="8"><strong>Modern</strong></td></tr><tr><td>TIAREI-A1</td><td>Tiarei</td><td>7.9</td><td>17.6</td><td>1.58 ± 0.31</td><td>0.139 ± 0.010</td><td>9.052 ± 0.023</td><td></td></tr><tr><td>TIAREI-3B</td><td>Tiarei</td><td>7.5</td><td>18.5</td><td>1.82 ± 0.36</td><td>0.123 ± 0.016</td><td>9.074 ± 0.007</td><td>1999–2008</td></tr><tr><td>TIAREI-4</td><td>Tiarei</td><td>6.5</td><td>11.3</td><td>1.62 ± 0.46</td><td>0.127 ± 0.027</td><td>9.035 ± 0.007</td><td>2000–2007</td></tr><tr><td>MARAA-7</td><td>Maraa</td><td>7.9</td><td>17.2</td><td>1.19 ± 0.25</td><td>0.127 ± 0.013</td><td>8.982 ± 0.013</td><td>2001–2007</td></tr><tr><td>MARAA-8</td><td>Maraa</td><td>2.5</td><td>12.8</td><td>1.78 ± 0.75</td><td>0.171 ± 0.029</td><td>9.112 ± 0.009</td><td>1999–2008</td></tr><tr><td>MARAA-8</td><td>Maraa</td><td>2.5</td><td>12.8</td><td>1.02 ± 0.10</td><td>0.164 ± 0.043</td><td>9.031 ± 0.009</td><td>2002–2005</td></tr><tr><td colspan="8"><strong>Early Holocene</strong></td></tr><tr><td>310-5A-12R-1 (25–60 cm)</td><td>Maraa</td><td>15.0</td><td>12.4</td><td>0.95 ± 0.43</td><td>0.074 ± 0.011</td><td>9.093 ± 0.007 (9.079)</td><td>11.03 ± 0.02 ka BP</td></tr><tr><td>310-7A-13R-1 (23–37 cm)</td><td>Maraa</td><td>4.6</td><td>10.8</td><td>0.59 ± 0.15</td><td>0.159 ± 0.023</td><td>9.171 ± 0.014 (9.165)</td><td>9.06 ± 0.01 ka BP</td></tr><tr><td>310-7A-15R-1 (23–32 cm)</td><td>Maraa</td><td>4.0</td><td>13.6</td><td>0.83 ± 0.27</td><td>0.177 ± 0.041</td><td>9.259 ± 0.006 (9.252)</td><td>9.37 ± 0.02 ka BP</td></tr><tr><td>310-7A-18R-1 (77–109 cm)</td><td>Maraa</td><td>6.0</td><td>36.3</td><td>1.76 ± 0.61</td><td>0.159 ± 0.030</td><td>9.374 ± 0.006 (9.364)</td><td>10.03 ± 0.02 ka BP</td></tr><tr><td>310-7B-11R-2 (0–14 cm)**</td><td>Maraa</td><td>8.7</td><td>–</td><td>1.13 ± 0.06</td><td>0.140 ± 0.031</td><td>9.293 ± 0.006 (9.285)</td><td>9.52 ± 0.03 ka BP</td></tr><tr><td>310-7B-21R-1 (0–33 cm)****</td><td>Maraa</td><td>4.4</td><td>9.5</td><td>0.79 ± 0.23</td><td>0.204 ± 0.021</td><td>9.122 ± 0.018 (9.107)</td><td>11.06 ± 0.04 ka BP</td></tr><tr><td colspan="8"><strong>Younger Dryas</strong></td></tr><tr><td>310-23A-5R-1 (73–78 cm)</td><td>Tiarei</td><td>1.1</td><td>14.0</td><td>0.60</td><td>0.071</td><td>9.170 (9.150)</td><td>12.33 ± 0.02 ka BP</td></tr><tr><td>310-23A-6R-1 (3–20 cm)*</td><td>Tiarei</td><td>10.2</td><td>–</td><td>0.77 ± 0.12</td><td>0.185 ± 0.017</td><td>9.205 ± 0.013 (9.184)</td><td>12.40 ± 0.05 ka BP</td></tr><tr><td colspan="8"><strong>Bølling–Allerød</strong></td></tr><tr><td>310-9A-6R-1 (11–26 cm)</td><td>Tiarei</td><td>3.0</td><td>12.7</td><td>0.84 ± 0.20</td><td>0.090 ± 0.028</td><td>9.194 ± 0.013 (9.165)</td><td>14.25 ± 0.02 ka BP</td></tr><tr><td>310-9D-7R-1 (15–50 cm)*</td><td>Tiarei</td><td>16.2</td><td>–</td><td>0.86 ± 0.40</td><td>0.159 ± 0.010</td><td>9.134 ± 0.006 (9.105)</td><td>14.22 ± 0.06 ka BP</td></tr><tr><td>310-18A-18R-1 (55–71 cm)</td><td>Maraa</td><td>8.0</td><td>10.9</td><td>0.96 ± 0.36</td><td>0.102 ± 0.024</td><td>9.088 ± 0.005 (9.059)</td><td>14.10 ± 0.03 ka BP</td></tr><tr><td>310-18A-19R-1 (29–65 cm)</td><td>Maraa</td><td>10.0</td><td>9.5</td><td>0.62 ± 0.04</td><td>0.097 ± 0.014</td><td>9.073 ± 0.013 (9.044)</td><td>14.14 ± 0.03 ka BP</td></tr><tr><td colspan="8"><strong>Meltwater Pulse 1A</strong></td></tr><tr><td>310-9E-9R-1 (32–36 cm)</td><td>Tiarei</td><td>2.4</td><td>9.0</td><td>0.45 ± 0.05</td><td>0.138 ± 0.049</td><td>9.197 ± 0.009 (9.168)</td><td>14.36 ± 0.04 ka BP</td></tr><tr><td>310-20A-22R-2 (123–150 cm)</td><td>Faaa</td><td>7.0</td><td>7.4</td><td>0.44 ± 0.08</td><td>0.125 ± 0.034</td><td>9.142 ± 0.008 (9.112)</td><td>14.37 ± 0.04 ka BP</td></tr><tr><td>310-20A-23R-1 (0–78 cm)</td><td>Faaa</td><td>32.9 (51.8)</td><td>9.5</td><td>0.60 ± 0.26</td><td>0.100 ± 0.030</td><td>9.087 ± 0.017 (9.057)</td><td>14.46 ± 0.06 ka BP</td></tr><tr><td>310-20A-23R-1 (78–132 cm)</td><td>Faaa</td><td>4.6</td><td>10.6</td><td>0.73 ± 0.12</td><td>0.097 ± 0.015</td><td>9.315 ± 0.010 (9.285)</td><td>14.35 ± 0.03 ka BP</td></tr><tr><td colspan="8"><strong>Heinrich Stadial</strong></td></tr><tr><td>310-9D-9R-1 (80–90 cm)</td><td>Tiarei</td><td>3.6</td><td>30.5</td><td>1.49 ± 0.31</td><td>0.199 ± 0.032</td><td>9.358 ± 0.004 (9.327)</td><td>14.66 ± 0.02 ka BP</td></tr><tr><td>310-18A-22R-1 (6–17 cm)</td><td>Maraa</td><td>3.0</td><td>5.4</td><td>0.43 ± 0.04</td><td>0.084 ± 0.036</td><td>9.135 ± 0.005 (9.104)</td><td>14.74 ± 0.09 ka BP</td></tr><tr><td>310-20A-24R-1 (30–41 cm)</td><td>Faaa</td><td>7.4</td><td>12.4</td><td>1.05 ± 0.33</td><td>0.113 ± 0.019</td><td>9.163 ± 0.011 (9.132)</td><td>14.74 ± 0.06 ka BP</td></tr><tr><td>310-24A-11R-2 (1–62 cm)***</td><td>Tiarei</td><td>22.2</td><td>–</td><td>1.99 ± 0.52</td><td>0.156 ± 0.013</td><td>9.203 ± 0.007 (9.171)</td><td>15.01 ± 0.07 ka BP</td></tr><tr><td colspan="8"></td></tr><tr><td>310-24A-13R-1 (27–45 cm)</td><td>Tiarei</td><td>9.0</td><td>14.9</td><td>0.88 ± 0.44</td><td>0.123 ± 0.031</td><td>9.225 ± 0.011 (9.193)</td><td>15.15 ± 0.03 ka BP</td></tr><tr><td>310-25B-9R-1 (35–45 cm)****</td><td>Tiarei</td><td>6.0</td><td>7.3</td><td>1.42 ± 0.37</td><td>0.083 ± 0.016</td><td>9.197 ± 0.007 (9.166)</td><td>14.74 ± 0.06 ka BP</td></tr></tbody></table>

Note. The number of annual cycles refers to the number of annual cycles used for reconstructions in (paleo-) seasonality. For record 310-20A-23R-1, the total number of annual cycles constructed including synchronous annual bands used for quantification of intra-colony variability is indicated in brackets. Mean Sr/Ca values shown are corrected for inter-laboratory offsets using the JCP-1 standard. The mean JCP-1 Sr/Ca value representative for the mean Sr/Ca in this study is 8.901 ± 0.020 mmol mol⁻¹. The mean Sr/Ca values that are corrected for last deglacial seawater [Sr] are shown in brackets assuming an elevated [Sr] during the Last Glacial Maximum of 0.5 %. Average annual linear extension rates, Sr/Ca-seasonality, and Sr/Ca mean values are also shown for each age interval. Sr/Ca records that have been published previously are indicated: Asami et al., 2009*; DeLong et al., 2010**; Felis et al., 2012***; Hathorne et al., 2011****. All deglacial coral ages are given relative to 1950 (i.e., BP).
climate (Asami et al., 2009; DeLong et al., 2010; Felis et al., 2012; Hathorne et al., 2011; Kubota et al., 2014) reconstructions.

In addition, five samples from modern *Porites* colonies were collected from the fore-reef slope at Tiarei and the lagoon at Maraa at 5–15 mbsl in the proximity of the drilling sites of IODP Expedition 310 (Figure 1; Table 1) in July 2008. Coordinates of sampling sites can be found in Table S1 in Supporting Information S1.

## 2.3. Sample Preparation and Diagenesis Screening

Samples analyzed in this study have been prepared and screened for diagenetic alteration at the MARUM—Center for Marine Environmental Science and the Faculty of Geosciences, University of Bremen, Germany (modern corals, 5A-12R-1 23–37 cm, 18A-18R-1 55–71 cm, 18A-19R-1 29–65 cm, 18A-22R-1 6–17 cm, 20A-22R-2 123–150 cm, 20A-23R-1B 78–132 cm, 20A-24R-1 30–41 cm, 24A-13R-127–45 cm), as well as at the Department of Earth Science, Tohoku University, Japan (7A-13R-1 23–37 cm, 7A-15R-1 23–32 cm, 7A-18R-1 77–109 cm, 9A-6R-1 48–67 cm, 9D-9R-1 80–90 cm, 9E-9R-1 32–36 cm, 23A-5R-1 73–78 cm). Cores were cut into 0.5–1 cm thick slices, ultrasonically cleaned, and subsequently X-radiographed to reveal the coral's annual banding pattern. Modern and last-deglacial corals may become subject to diagenetic alteration (Felis et al., 2012; Hathorne et al., 2011; Hendy et al., 2007). Therefore, areas in coral skeletons diagenetically altered or affected by microbial micritic products were identified through visual inspection of X-radiographs and excluded from sampling. Continuous sampling transects were defined parallel to the coral's growth axes. However, sampling transects had to be sectioned in some areas to ensure subsampling of pristine material only. Furthermore, powder X-ray diffraction (XRD) analysis was carried out to assess the preservation, especially of deglacial coral skeletons. Along sampling transects, samples for powder XRD analysis were collected to determine the presence of calcite. In order to detect areas affected by secondary aragonite precipitation, thin sections were produced from subsamples adjacent to the sampling transects, and subsamples cut from the coral slices were inspected for dissolution using a scanning electron microscope (SEM). Mg/Ca is the trace element ratio most sensitive to diagenetic alteration (Felis et al., 2012; Hathorne et al., 2011; Hendy et al., 2007) and, consequently, samples showing anomalies in Mg/Ca were removed from the analysis. Furthermore, the alignment of the corallites along the sampling path can impact the resultant geochemical time series (DeLong et al., 2007, 2013). Therefore, an assessment of corallite alignment was performed following Reed et al. (2021). Results of the diagenesis screening and assessment of sampling transects can be found in Table S2 in Supporting Information S1. X-radiographs with sampling tracks, as well as the sampling locations for diagenesis screening and ²³⁰Th-U dating, can be found in the Supplement, Figures S1–S24 in Supporting Information S1.

## 2.4. ²³⁰Th-U Dating

Samples of 1–2 g were cut from pristine coral material, ultrasonically cleaned, and dried. Subsequently, ²³⁰Th-U dating was performed at CEREGE in Aix-en-Provence, France, and the Institute for Geosciences, Johannes Gutenberg-University, Mainz (JGU), Germany. At CEREGE subsamples were analyzed for ²³⁰Th-U using a VG-54 thermo-ionization mass spectrometer equipped with a 30-cm electrostatic analyzer and a pulse-counting Daly detector, following methods published in Deschamps et al. (2012). At JGU, subsamples of ca. 300 mg were prepared by column chemistry and analyzed using a Neptune Plus multi-collector inductively coupled plasma mass spectrometer (MC-ICPMS). The weighed samples were dissolved in 7N HNO₃ and a mixed ²²⁹Th–²³³U–²³⁶U spike was added (see Gibert et al., 2016 for details on spike calibration). Potential organic material was removed by adding a mixture of concentrated HNO₃, HCl and H₂O₂. The dried samples were dissolved in 6N HCl, and U and Th were separated using ion exchange columns (Yang et al., 2015). Technical details about the MC-ICPMS procedures are described in Obert et al. (2016). ²³⁰Th-U ages were corrected for detrital contamination assuming a ²³²Th/²³⁸U weight ratio of 3.8 for the detritus and ²³⁰Th, ²³⁴U and ²³⁸U in secular equilibrium.

## 2.5. Micro-Sampling and Geochemical Analysis

Sampling of pristine coral skeletal material along defined sampling transects at MARUM, University of Bremen, Germany, was conducted by manual low-speed drilling using a dentist tool with a drill bit of 0.7 mm diameter. Subsamples removed were adjacent to each other with minimal overlap approximately every 0.75–1 mm apart
(i.e., continuous spot sampling; Felis et al., 2004). Samples were collected by drilling maximal 2 mm into the coral skeleton. Sampling at Tohoku University, Japan, employed a low-speed diamond dental disk where a sample was removed approximately every 0.5 mm.

An aliquot of 0.05 mg was removed from each subsample for δ¹⁸O and δ¹³C analysis. Oxygen and carbon isotope analyses at the Stable Isotope Laboratory of MARUM, University of Bremen, Germany, employed a Finnigan MAT 251 mass spectrometer equipped with an automated Kiel carbonate preparation device and a dual inlet system. Oxygen and carbon isotope analyses at the Japan Oil, Gas, and Metals National Corporation (JOGMEC) in Tokyo, Japan, and the Center for Advanced Marine Core Research in Kōchi, Japan, employed a Kiel III automated sample preparation system coupled to a Finnigan MAT 252 mass spectrometer, MAT253 mass spectrometer respectively. Isotopic values are reported in ‰ relative to the Vienna PeeDee Belemnite (VPDB) reference standard. For all isotope measurements performed in this study, instrumental precision was better than 0.05 ‰ for δ¹⁸O and better than 0.03 ‰ for δ¹³C.

Trace element analysis for all samples was performed at the Sediment Geochemistry Laboratory of MARUM, University of Bremen, Germany. For Sr/Ca and Mg/Ca analysis, aliquots of 0.25 mg were dissolved in 7 ml of 2% suprapure HNO₃ containing 1 ppm Scandium as internal standard. The Ca concentration of dissolved samples varied between 7 and 16 ppm. Trace element analysis employed an Agilent 720 (axial) inductively coupled plasma optical emission spectrometer (ICP-OES; modern corals), a Perkin-Elmer Optima 3300R (simultaneous radial) ICP-OES, respectively (last-deglacial corals), and followed methods detailed in Felis et al. (2012). Instrumental drift was offline corrected by a sample-standard bracketing approach and instrumental precision for Sr/Ca determinations was better than 0.2%. Mean Sr/Ca for the JCp-1 standard (Hathorne et al., 2013; Okai et al., 2002) measured during this study was 8.901 ± 0.020 mmol mol⁻¹ (n = 20) for measurements carried out on the Agilent 720, and 8.909 ± 0.026 mmol mol⁻¹ (n = 57) for the Perkin-Elmer Optima 3300R. For comparison between Sr/Ca records, Sr/Ca measurements from different measurement sessions were corrected for the offset in the mean Sr/Ca value of the JCp-1 standard relative to the mean measured with the modern *Porites* (8.901 ± 0.020 mmol mol⁻¹). This was also done for the mean value of the JCp-1 standard reported in previous studies: 8.836 ± 0.018 mmol mol⁻¹ (Asami et al., 2009), 8.850 ± 0.020 mmol mol⁻¹ (Hathorne et al., 2011), 8.924 ± 0.014 mmol mol⁻¹ (Felis et al., 2012), and 8.841 ± 0.029 mmol mol⁻¹ (DeLong et al., 2010). Consequently, the mean JCp-1 Sr/Ca value representative for the mean Sr/Ca values shown in this study is 8.901 ± 0.020 mmol mol⁻¹.

## 2.6. Age Chronology

For Sr/Ca, δ¹⁸O, and δ¹³C time series, data points were considered outliers if they fell outside the 99% (3σ) confidence interval, and were subsequently removed from the analysis. The sampling resolution among coral samples varied due to variations in linear extension (Table 1). To reduce the effect of variable sampling resolution on reconstructions in SST seasonality, for each time series, data points within approximately the same month were averaged (Piecewise Aggregate Approximation) so that each annual cycle consisted of no more than 12 data points (i.e., normalization for sampling resolution). Furthermore, following Lawman et al. (2020) a 2-month low-pass filter was applied to reduce high-frequency noise. An age chronology was established by identifying annual cycles in Sr/Ca where the lowest Sr/Ca was set to austral summer and the highest Sr/Ca was set to austral winter. The warmest or coldest months do not occur at the same time every year (Figures 3e and 3f). Consequently, for last-deglacial coral records, the lowest and highest Sr/Ca values were randomly set to one of the warmest and coldest months according to their respective relative occurrences. For modern corals, the exact year each annual band was precipitated could be determined. Thus, for these records, the lowest and highest Sr/Ca values were set to the precise month with the highest and lowest SSTs of the respective year as inferred from the NOAA Optimum Interpolation Sea Surface Temperature Version 2.1 (OISST v2.1; Huang et al., 2020) data set (Figure 3d). The age of data points between the tie points was calculated linearly based on their relative depth. The age-depth model construction was further supported by the alignment of the Sr/Ca time series with the annual banding pattern in the coral skeletons visible in the X-radiographs (Figures S1–S24 in Supporting Information S1) and annual cycles of the other geochemical proxies. The same age-depth model based on seasonal cycles in Sr/Ca was applied to all other proxies. Both modern and last-deglacial corals precipitated the high-density band during austral summer. Annual linear extension rates were calculated by measuring the distance along the growth axis for each annual cycle. Finally, time series were linearly interpolated to monthly resolution for comparison with instrumental and climate model data.
## 2.7. Statistics

Statistical analyses were performed using the statistical computing software R (R Core Team, 2021). For the calibration of modern coral Sr/Ca against SST, the OISST v2.1 data set for the grid cell 17°–18° South, 149°–150° West was used that covers entirely Tahiti and the neighboring island of Moorea (Huang et al., 2020; Figures 1 and 2a). Global gridded SST data sets provide spatially and temporally smoothed data that are often, especially in remote areas, based on only a few observations (Kennedy et al., 2011a, 2011b). In contrast, coral geochemical data represents very local seawater conditions. Thus, to improve our calibration we follow the approach from Sayani et al. (2019) and use a composite Sr/Ca record from the five modern corals that grew synchronously at different locations at Tiarei and Maraa on Tahiti (Figure 1; Figure 3a). Therefore, the Sr/Ca time series were adjusted by subtracting the offsets between the mean Sr/Ca of the longest modern record TIAREI-A1 and the means of the other records (Figures 3a and 3b). Subsequently, a composite record was calculated from the 5 modern records. This composite record aims to better represent average conditions at Tahiti and is, therefore, more suitable for calibration against the OISST v2.1 data set. To account for the uncertainties in regressor (SST) and response (Sr/Ca) variables a weighted least squares (WLS) regression was applied following Thirumalai et al. (2010) using the standard error of the mean (SE) of the composite Sr/Ca record and the OISST v2.1 SST uncertainty to assign weights (Figure 3c; York et al., 2004; Xu et al., 2015).

The seasonality and mean values in geochemical proxies were calculated for each coral colony. Therefore, a mean seasonal cycle was calculated by averaging the proxy values of the same months (Figures 2 and 7). The SE was calculated for each month's average which takes into account how many monthly proxy values were included in each month's average. Subsequently, the seasonality for each coral colony was calculated by subtracting the highest from the lowest monthly proxy value in the mean annual cycle. The SE of the seasonality was derived through Gaussian principals of error propagation, that is, the root of the sum of the squares of the SE of the month with the highest and lowest proxy value in each coral colony's mean annual cycle. Furthermore, a mean was calculated from annual proxy values for each coral colony. For each of these coral colony mean values, a SE was determined that takes into account how many annual cycles were included.

Additionally, for last-deglacial corals, a combined error (CE) was calculated for each coral colony's seasonality and mean following Giry et al. (2012), Felis et al. (2015), and Brocas et al. (2018) that accounts for the inter-colony variability in proxy seasonality and mean value. The CE was calculated as the root of the sum of the squares of the SD around the mean of the five modern corals' seasonality (Figure 2a) and the SE of the respective colony's seasonality. The CE for the mean proxy values for each colony was calculated accordingly. Subsequently, the SE and CE of seasonality and mean values are used to determine if the differences between values are significant.

Furthermore, colonies were grouped based on their age (Table 1, Figure 4). Following Abram et al. (2009) a weighted mean and uncertainty for each proxy seasonality and mean proxy value was calculated for each age group. Weights were assigned to each coral colony's seasonality or mean value based on the number of seasonal cycles in the respective coral colony. Estimates of mean $\Delta\delta^{18}$O and associated uncertainty (SE and CE) for each age interval were calculated following Cahyarini et al. (2008) using the $\delta^{18}$O-SST and Sr/Ca-SST slopes from Gagan et al. (1998) and this study, respectively.

Absolute reconstructions in SST from mean coral Sr/Ca and $\delta^{18}$O on a glacial-interglacial time scale are complicated by changes in seawater chemistry. Seawater $\delta^{18}$O varies with global ice volume and thus fluctuates following sea level (Shackleton, 1967). Seawater [Sr] is thought to be modulated by increased weathering of continental shelves that are exposed during sea level low stands (Stoll & Schrag, 1998). Furthermore, Sr incorporation into the coral skeletons has been suggested to be affected by seawater pH (Cole et al., 2021) which has likely fluctuated on a glacial-interglacial time scale (Pelejero et al., 2010). Therefore, the mean $\delta^{18}$O of coral records was corrected for seawater $\delta^{18}$O at a given age following Waelbroeck et al. (2002). For each record, mean Sr/Ca was corrected for seawater [Sr] by linear interpolation of seawater [Sr] dependent on the sea level at a given age (Waelbroeck et al., 2002) assuming seawater Sr/Ca was increased by 0.5% (Asami et al., 2009; Felis et al., 2012; Stoll & Schrag, 1998) relative to today (8.539 mmol mol$^{-1}$; De Villiers, 1999) at minimum sea level during the Last Glacial Maximum.
<figure><img src="image_7.png" /></figure>

**Figure 4.** Modern and last deglacial temperature for (a) the Northern Hemisphere reconstructed from NGRIP ice core $\delta^{18}$O (North Greenland Ice Core Project members, 2004) and (c) the Southern Hemisphere reconstructed from EDML ice core $\delta^{18}$O (EPICA Community Members, 2006). (b) Mean Sr/Ca from modern and IODP Expedition 310 fossil Tahiti corals. Symbols represent each record mean, and thick horizontal bars represent the weighted mean of colony means within a given period with uncertainty (CE) as light blue area. The light red column indicates the interval of Meltwater Pulse 1A (Deschamps et al., 2012). Reconstructed SST anomaly using the slope of this study (–0.050 mmol mol$^{-1}\circ$C$^{-1}$)*, Gagan et al., 2012 (–0.084 mmol mol$^{-1}\circ$C$^{-1}$)**, Felis et al., 2009 (–0.140 mmol mol$^{-1}\circ$C$^{-1}$) ***. SST reconstructions are not adjusted for changes in seawater Sr/Ca and thus provide an upper estimate of the magnitude of cooling. Uncertainty for record means is 2$\sigma$ for the $^{230}$Th-U age and 1$\sigma$ of the mean for the reconstructed SST. Also shown is the mean JCp-1 Sr/Ca measured along with the modern *Porites* corals from this study (8.901 ± 0.020 mmol mol$^{-1}$) that is used for corrections of inter-laboratory offsets. Mean coral Sr/Ca from previous studies are also shown (Asami et al., 2009; DeLong et al., 2010; Felis et al., 2012; Hathorne et al., 2011).

# 3. Results

## 3.1. Geochemistry

Powder XRD and SEM analyses, as well as the inspection of thin sections, allowed the detection of areas likely affected by the deposition of secondary calcium carbonate and dissolution. Respective skeletal material was not subsampled for geochemical analysis. Furthermore, the alignment between sampling transects and corallites determines the quality of the constructed proxy time series (DeLong et al., 2013; Reed et al., 2021). Full results of sampling path quality assessment, XRD analyses, SEM imagery, and thin section inspection can be found in Table S2 (sampling path quality assessment), Table S3 (XRD), Figures S26–S42 (thin section imagery), and Figures S43–S73 (SEM imagery) in Supporting Information S1.

Results of $^{230}$Th-U dating revealed that coral samples span a time interval between the early Holocene (9.06 ± 0.01 ka BP) to Heinrich Stadial 1 (15.15 ± 0.03 ka BP) with major clusters in the Bølling–Allerød warm period and around Meltwater Pulse 1A (MWP1A; Figure 4b; Table 1). Full results of $^{230}$Th-U ages presented in this study for the first time can be found in Tables S4 and S5 in Supporting Information S1. All previously published $^{230}$Th-U dates used in this study (Asami et al., 2009; DeLong et al., 2010; Deschamps et al., 2012; Felis et al., 2012; Hathorne et al., 2011) are listed in Table S6 in Supporting Information S1.

Coral $^{230}$Th-U samples were largely pristine and did not contain large amounts of detrital Th. Nevertheless, for samples 310-9A-6R-1 (11–26 cm) and 310-7A-15R-1 (23–32 cm), the correction for detrital contamination is between 30 and 40 years and, thus, significant within error. This is accounted for by slightly enlarged age
uncertainties. Samples showed initial ($^{234}$U/$^{238}$U) activity ratios that are in agreement with the value for modern seawater (Chutcharavan et al., 2018) providing further evidence that the corals have not been affected by diagenetic alteration.

For seasonality reconstructions, 153.9 monthly resolved annual cycles in Sr/Ca, $\delta^{18}$O, and $\delta^{13}$C were obtained from 20 *Porites* colonies and are presented in this study for the first time (Figure 5; Figures S75 and S76 in Supporting Information S1). These include 32.3 annual cycles from five modern and 121.6 annual cycles from 16 fossil colonies. Additionally, previously published 67.7 monthly resolved annual cycles from six corals were included in this study (Asami et al., 2009; DeLong et al., 2010; Felis et al., 2012; Hathorne et al., 2011). Furthermore, for record 310-20A-23-1 (0–78 cm), 21.7 annual cycles in geochemical proxies were constructed from annual bands that grew synchronously and thus have similar ages (Figure 5u). Of these 21.7 annual cycles 14.8 annual cycles were used for the assessment of intra-colony variability. Hence, in total, 32.3 monthly resolved annual cycles from five modern corals and 189.3 monthly resolved annual cycles from 22 last-deglacial corals are presented in this study.

Furthermore, samples were binned according to their age (Figures 4b and Table 1) with 32.3 annual cycles from five corals representing modern Tahiti climate (1986–2008), 42.7 annual cycles from six corals representing the Early Holocene (9.06–11.06 ka BP), 11.3 annual cycles from two corals representing the Younger Dryas (12.33–12.40 ka BP), 37.2 annual cycles from four corals representing the Bølling–Allerød (14.10–14.25 ka BP), 46.9 annual cycles from four corals representing MWP1A (14.35–14.46 ka BP), and 51.2 annual cycles from six corals representing Heinrich Stadial 1 (14.66–15.15 ka BP).

The correlation between annual linear extension rates and corresponding annually averaged Sr/Ca values from all modern and fossil corals was very weak ($r^2 = 0.03, p = 0.01$) indicating that skeletal Sr/Ca is primarily SST controlled with minimal influence of coral growth and calcification rate effects (Figure S25a in Supporting Information S1). This is in line with results from culturing experiments using *Porites* corals (Inoue et al., 2007).

## 3.2. Inter- and Intra-Colony Variability

Environmental conditions including SST and SSS on Tahiti reefs may vary spatially due to, for instance, local differences in upwelling and internal waves (Wolanski & Delesallet, 1995; Wyatt et al., 2020) or orographic effects on rainfall (Bessemoulin et al., 2004). Therefore, synchronously growing corals across Tahiti may record different environmental signals. Furthermore, the geochemical composition of *Porites* skeletons exhibits substantial inter- and intra-colony variability as a result of physiological differences between colonies and kinetic effects induced by, for instance, variations in linear extension rates (Cohen & Hart, 1997; DeLong et al., 2011; Felis et al., 2003, 2004). To assess the inter-colony variability of Tahiti corals and quantify the representativeness of coral climate archives for mean climate conditions at Tahiti, the five modern *Porites* records were compared (Figures 2a and 3a). Variability in Sr/Ca seasonality was 0.023 mmol mol$^{-1}$ (1$\sigma$) that equals 0.5$^{\circ}$C (Figure 2a) and variability in mean Sr/Ca was 0.051 mmol mol$^{-1}$ (1$\sigma$) that equals 1.0$^{\circ}$C (Figure 3b). Additionally, for assessment of intra-colony Sr/Ca variability, the geochemical signal from annual bands of last deglacial *Porites* coral 20A-23R-1 (0–78 cm) that were deposited synchronously were compared (Figure 2b; Figure S19 in Supporting Information S1). Here, variability in Sr/Ca seasonality was 0.007 mmol mol$^{-1}$ (1$\sigma$) that equals 0.1$^{\circ}$C and variability in mean Sr/Ca was 0.090 mmol mol$^{-1}$ (1$\sigma$) that equals 1.8$^{\circ}$C. For the seasonality reconstruction, the intra-colony variability is substantially smaller than the inter-colony variability as the first is included in the latter. Inter-colony variability in $\delta^{18}$O was 0.04 $\permil$ (1$\sigma$; seasonality) and 0.07 $\permil$ (1$\sigma$; mean). Intra-colony $\delta^{18}$O variability was 0.07 $\permil$ (1$\sigma$; seasonality) and 0.24 $\permil$ (1$\sigma$; mean). Similarly, inter-colony variability in $\delta^{13}$C was 0.07 $\permil$ (1$\sigma$; seasonality) and 0.22 $\permil$ (1$\sigma$; mean), and intra-colony $\delta^{13}$C variability was 0.19 $\permil$ (1$\sigma$; seasonality) and 0.59 $\permil$ (1$\sigma$; mean).

## 3.3. Proxy Calibration

The slope between the composite Sr/Ca record from the five modern Tahiti corals and OISST v2.1 is -0.050 ± 0.002 mmol mol$^{-1}\^{\circ}$C$^{-1}$ and the correlation between both variables is significant (Figure 3c; $r^2 = 0.72, p < 0.01$). Results are in agreement with previous Sr/Ca-SST calibrations (Corrège, 2006). An ordinary least squares (OLS) regression-based calibration in Felis et al. (2012) between Sr/Ca values from one Tahiti *Porites* colony and proximal IRD SST resulted in a similar slope of -0.057 ± 0.005 mmol mol$^{-1}\^{\circ}$C$^{-1}$ giving additional
<figure><img src="image_8.png" /><figcaption>**Figure 5.** Monthly time series of modern and IODP Expedition 310 fossil Tahiti coral Sr/Ca-based reconstructions of SST (using the calibration of this study: -0.050 mmol mol⁻¹C⁻¹). Note the different scale of the SST y-axis of panel u. In u, different colors indicate different microsampling transects. See Table 1 for more details.</figcaption></figure>

confidence in the calibration presented here. Subsequently, the slope of -0.050 ± 0.002 mmol mol⁻¹C⁻¹ is used for translations of Sr/Ca values to SST estimates in this study.

To further test the fidelity of the OISST v2.1 product, SST time series were compared with in situ measurements. SST was recorded between 1979 and 1991 at Papeete, Tahiti (IRD; Figure 3d). Furthermore, a composite record of SST loggers deployed at various water depths (0–55 m) at Tiahura, Moorea, by CRIOBE provided SST measurements between 1992 and 2018 (Figure 3d). OISST v2.1 and in situ measurements revealed a high level of agreement ($r^2$ = 0.89, *p* < 0.01). Mean SST over similar time intervals was 27.2 ± 0.9°C (SE, IRD) and 27.2 ± 1.0°C (SE, OISST v2.1), as well as 27.7 ± 0.9°C (SE, CRIOBE) and 27.7 ± 1.0°C (SE, OISST v2.1). Similarly, SST seasonality was 2.2 ± 0.2°C (SE, IRD) and 2.6 ± 0.2°C (SE, OISST v2.1), as well as 2.2 ± 0.1°C (SE, CRIOBE) and 2.4 ± 0.1°C (SE, OISST v2.1).
# 4. Discussion

## 4.1. Last Deglacial SST Seasonality

For the modern *Porites* corals, Sr/Ca seasonality is 0.139 ± 0.010 mmol mol⁻¹ (2.8 ± 0.2°C; SE) and the mean OISST v2.1-derived SST seasonality over the same period is 2.6 ± 0.1°C (SE). Thus, the OISST v2.1 SST seasonality is within the uncertainty of the Sr/Ca-derived modern SST seasonality (Figure 2a) which supports the reconstructions of paleo-seasonality from deglacial coral records. Although Sr/Ca- and OISST v2.1-derived SST seasonality are statistically identical, the slope used in this study (−0.050 ± 0.002 mmol mol⁻¹°C⁻¹) shows the tendency to overestimate SST seasonality from Sr/Ca (Figure 3d). However, note that the relative differences in Sr/Ca-based calculations of seasonality between records remain similar, independently from the slope that is used for estimating SST from Sr/Ca values (Table 1).

The large inter-colony variability in coral Sr/Ca seasonality and mean complicates reconstructions of these parameters in absolute SST from coral records. To reduce the uncertainty associated with our reconstructions we therefore averaged parameters over multiple annual bands and colonies. Thus, uncertainties of all results are indicated as SE and additionally, as CE that incorporates the inter-colony-variability as observed between the five modern corals in the uncertainty estimates. This procedure has been applied previously for similar studies of coral Sr/Ca-derived seasonality and absolute SST reconstructions (Brocas et al., 2018; Giry et al., 2012). Given the large CE of the seasonality reconstructions no statistically significant difference between modern and last-deglacial corals is detectable (Figure 6a). However, when considering the SE some significant variations become apparent. While the reconstructed seasonality for the Early Holocene with 0.130 ± 0.008 mmol mol⁻¹ (2.6 ± 0.2°C; SE), the Bølling–Allerød with 0.123 ± 0.008 mmol mol⁻¹ (2.5 ± 0.2°C; SE), and Heinrich Stadial 1 with 0.133 ± 0.008 mmol mol⁻¹ (2.7 ± 0.2°C; SE) remain statistically indistinguishable from the modern Sr/Ca seasonality, the seasonality for the Younger Dryas is with 0.171 ± 0.017 mmol mol⁻¹ (3.4 ± 0.3°C; SE) significantly elevated, and for MWP1A with 0.105 ± 0.015 mmol mol⁻¹ (2.1 ± 0.3°C, SE) significantly reduced compared to modern (Figure 6a). A reduced SST seasonality during MWP1A may reflect the reduction in insolation seasonality during that time (Berger, 1978; Figure 6d). Furthermore, the observed increase in Sr/Ca seasonality in Tahiti corals from the Younger Dryas suggests that the local SST seasonality was further modulated by local processes. Warming of the surface ocean during summer depends largely on the mixed layer depth as it determines the volume of water over which the incoming atmospheric heat flux is integrated (Alexander et al., 2018; Jo et al., 2022). Therefore, an increase in ocean stratification associated with a shoaling of the surface mixed layer results in an increase in SST seasonality, particularly an increase in summer SST. A stronger stratification of the surface ocean at Tahiti during the Younger Dryas implies that Tahiti was likely not located within the SPCZ at that time. Thus, the center of the South Pacific Subtropical Gyre which is characterized by a stronger stratification of the upper ocean and is today located in the southeast of Tahiti was likely located or had expanded further northwards. Previous studies have demonstrated that the location of the SPCZ has varied on a decadal to centennial time scale affecting mean climate conditions of Southwest Pacific tropical islands (Maupin et al., 2014; Partin et al., 2013) and others suggested that the SPCZ was likely located further northwards or had even partially merged with the Intertropical Convergence Zone (ITCZ) during the Younger Dryas (Asami et al., 2009; Corrège et al., 2004). Furthermore, studies found that the ITCZ in the Eastern and Central tropical Pacific may have been located further southward during the Younger Dryas in response to AMOC slowdown and associated cooling of the tropical North Atlantic that intensified the northeasterly trade winds across the Central American isthmus (Jacobel et al., 2016; Timmermann, Okumura, et al., 2007). It can be speculated that a more southward-located ITCZ may have partially merged with a further northward-located SPCZ. However, these studies also suggest that the annual SST cycle in the tropical Eastern Pacific decreased during the Younger Dryas and Heinrich Stadial 1 as a result of a further south-ward located ITCZ contrasting our results from Tahiti in the tropical Central Pacific. The elevated annual SST cycle reconstructed from Tahiti corals from the Younger Dryas, therefore, supports results from other studies that found that the southward migration of the ITCZ in the Pacific during the Younger Dryas has been only limited (Kumar et al., 2021; Xie et al., 2008). Overall, these results suggest that during most of the last deglaciation, the reduced insolation seasonality affected the SST seasonality at Tahiti only by a small amount that is below the detection limits of coral Sr/Ca-thermometry. The Younger Dryas is an exception as during this interval a shoaling of the mixed layer depth associated with a northward migration or even partial inactivity of the SPCZ likely elevated SST seasonality. It is also worth mentioning that between records no significant correlation (*p* > 0.05) in mean Sr/Ca and Sr/Ca seasonality was found that would point to a
<figure><img src="image_9.png" /><figcaption>**Figure 6.** Comparison of mean parameters between modern and IODP Expedition 310 fossil Tahiti coral age groups: (a) coral Sr/Ca-based SST seasonality with modern OISST v2.1 seasonality (red line) with SE, (b) coral δ¹⁸O seasonality, (c) coral δ¹³C seasonality, and (d) insolation seasonality at the latitude of Tahiti from Berger (1978). Thick black error bars represent the respective SE and thin error bars with caps represent the respective CE. Uncertainties are at 1σ level.</figcaption></figure>

general dependency between mean climate state and seasonality as reported in another study from the South China Sea (Yan et al., 2015).

The seasonal patterns in coral δ¹⁸O are largely similar to the Sr/Ca and SST seasonality with lower δ¹⁸O values in summer compared to winter (Figures 7a–7f; Figure S75 in Supporting Information S1). On a multi-millennial scale, calculations in δ¹⁸O seasonality for the last deglaciation are statistically indistinguishable from the δ¹⁸O seasonality observed for the modern corals (Figure 7a). Nevertheless, the δ¹⁸O seasonality for the Younger Dryas is significantly higher (SE) compared to the Bølling–Allerød and Meltwater Pulse 1A. Although coral skeletal δ¹⁸O from Tahiti mostly reflects paleo-SST variations (Cahyarini et al., 2008) this proxy is additionally influenced by variations in δ¹⁸Oₗseawater and therefore also reflects variations in SSS and rainfall (Ren et al., 2003). Potential causes for an elevated δ¹⁸Oₗseawater seasonality especially during the Younger Dryas are a higher discrepancy between rainfall in summer and winter or a higher discrepancy between seawater and meteoric δ¹⁸O.

<figure><img src="image_10.png" /><figcaption>**Figure 7.** Mean seasonal cycles in Sr/Ca, δ¹⁸O and δ¹³C for modern and IODP Expedition 310 fossil Tahiti corals. A composite (weighted mean) annual cycle for each age group is also shown (thick line) to illustrate that a similar pattern is found among all age groups despite differences between single records. Corals shown are: Modern (a, g, m) TIAREI-1 (blue line), TIAREI 3B (gray), TIAREI-4 (purple), MARAA-7 (green), MARAA-8 (orange), Early Holocene (b, h, n) 310-7B-11R-2 (0–14 cm; blue), 310-7A-18R-1A (gray), 5A-12R-1A (purple), 7B-21R-1A (red), 7A-13R-1A (green), 7A-15R-1A (orange), Younger Dryas (c, i, o) 310-23A-5R-1A (red), 310-23A-6R-1 (3–20 cm; black), Bølling–Allerød (d, j, p) 310-18A-18R-1A (blue), 310-18A-19R-1A (gray), 310-9D-7R-1 (15–50 cm; green), 310-9A-6R-1A (orange), MWP1A (e, k, q) 310-20A-23R-1B (blue), 319-9E-9R-1A (gray), 20A-22R-2A (green), and 310-20A-23R-1A (orange), Heinrich Stadial 1 (f, l, r) 18A-22R-1A (blue), 20A-24R-1A (gray), 310-24A-11R-2 (1–62 cm; purple), 24A-13R-1A (red), 310-25B-9R-1A (green), 310-9D-9R-1A (orange).</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>**Figure 8.** Comparison of mean parameters between modern and IODP Expedition 310 fossil Tahiti coral age groups: (a) Mean coral Sr/Ca-based SST corrected for variations in paleo-seawater Sr/Ca assuming that seawater Sr/Ca was increased by 0.5% relative to today during the last glacial maximum, (b) mean coral $\Delta\delta^{18}\text{O}$ corrected for the ice volume effect relative to modern, and (c) mean coral linear extension rate. Thick black error bars represent the respective SE and thin gray error bars with caps represent the respective CE, both at $1\sigma$ level.</figcaption></figure>

at Tahiti. Jasechko et al. (2015) reported that a majority of records for glacial $\delta^{18}\text{O}_{\text{rainfall}}$ show reduced values compared to modern. At the same time, glacial $\delta^{18}\text{O}_{\text{seawater}}$ (Figure 8b) was elevated due to the ice-volume effect (Shackleton, 1967). Furthermore, elevated $\delta^{18}\text{O}_{\text{seawater}}$ at Tahiti during the Younger Dryas aligns with a stronger influence of the South Pacific Subtropical Gyre as this region is characterized by higher $\delta^{18}\text{O}_{\text{seawater}}$ (LeGrande & Schmidt, 2006). For the Early Holocene, the annual $\delta^{18}\text{O}$ cycle of the two youngest corals (310-7A-13R-1 (23–37 cm), 9.06 ± 0.01 ka BP; 310-7A-15R-1 (23–32 cm), 9.37 ± 0.02 ka BP) offset by 3–4 months relative to the annual cycles of older Early Holocene corals covering the time interval of 9.52 ± 0.03 ka BP to 11.06 ± 0.04 ka BP (Figure 7h). This may indicate a shift in the hydroclimatological regime during this period.

In most records, coral $\delta^{13}\text{C}$ shows a seasonal cycle with lower values in summer and higher values in winter with a seasonality ranging between 0.86 ± 0.22 ‰ during the Younger Dryas and 0.31 ± 0.08 ‰ today (Figures 7g–7l; Figure 6c; Figure S76 in Supporting Information S1). Boisseau et al. (1998) found a similar seasonal cycle in modern Tahiti coral $\delta^{13}\text{C}$ and identified variations in photosynthetic activity as the primary driver. The positive relationship between coral $\delta^{13}\text{C}$ and both, photosynthetic activity and irradiance results from the preferential sequestration of $^{12}\text{C}$ by coral zooxanthellae increasing the fraction of $^{13}\text{C}$ in the internal dissolved inorganic carbon (DIC) pool that is used for coral biomineralization (Grottoli & Wellington, 1999). Boisseau et al. (1998) argued that in Tahiti the increased cloud cover during the rainy season in austral summer would result in a reduction of photosynthesis during that period despite total irradiance being maximal in summer. However, a more recent publication (Hopuare et al., 2019) has shown that irrespective of seasonal variations in cloud cover, available solar energy in Tahiti follows the seasonal cycle in irradiance, insolation respectively (Figure 2b). Therefore, we believe it is unlikely that the seasonal cycle in coral $\delta^{13}\text{C}$ at Tahiti is primarily driven by variations in photosynthetic activity. Other potential drivers of coral $\delta^{13}\text{C}$ variability include seasonal variations in calcification rates with lower $\delta^{13}\text{C}$ values coinciding with increased calcification rates, changes in ambient seawater DIC $\delta^{13}\text{C}$, and variations in the relative levels of heterotrophy (Allison et al., 1996; Felis et al., 1998). No calcification rates with sub-seasonal resolution are available for Tahiti corals. However, corals from this study revealed a weak but significant negative correlation between annual linear extension rates and skeletal $\delta^{13}\text{C}$ ($r^2 = 0.14, p < 0.01$; Figure S25b in Supporting Information S1). The carbon isotopic composition of DIC from terrestrial runoff (Moyer & Grottoli, 2011) and upwelling (Watanabe et al., 2017) is characterized by $\delta^{13}\text{C}$ lower than coastal seawater and increased heterotrophic energy supply initiated through increased upwelling and associated enhanced abundances of zooplankton results in lower skeletal $\delta^{13}\text{C}$ values (Felis et al., 1998). Indeed, coral $\delta^{13}\text{C}$ in Tahiti shows lower values during the late rainy season when the discharge of meteoric waters is expected to be highest. Overall, the exact driver of seasonal variations in Tahiti coral $\delta^{13}\text{C}$ remains unclear but is likely related to seasonal variations in calcification rates or the local hydrological regime, rather than variations in irradiance. Seasonality in $\delta^{13}\text{C}$ shows the tendency of being increased in last-deglacial corals relative to modern corals (Figure 6c), with a significantly elevated $\delta^{13}\text{C}$ seasonality during the Younger Dryas. Following the above, it can be speculated that the elevated $\delta^{13}\text{C}$ seasonality in deglacial corals most likely results from stronger kinetic
effects in skeletal δ¹³C due to reduced deglacial calcification rates (linear extension rates; Figure 8c). An enhanced discrepancy between summer and winter rainfall during the Younger Dryas was suggested as a potential cause of the observed increase in skeletal δ¹⁸O seasonality. Similarly, the increased discrepancy between summer and winter runoff may have contributed to the substantially elevated δ¹³C seasonality during the Younger Dryas.

## 4.2. Last Deglacial Changes in Mean Environmental Conditions

Mean coral skeletal Sr/Ca at Tahiti (Figures 4b and 8a) was significantly elevated (reconstructed SST reduced) relative to the mean Sr/Ca from the modern *Porites* corals during the Early Holocene by +0.129 ± 0.049 mmol mol⁻¹ (−2.6 ± 1.0°C; SE), the Younger Dryas by +0.134 ± 0.012 mmol mol⁻¹ (−2.7 ± 0.2°C; SE), and Heinrich Stadial 1 by +0.122 ± 0.020 mmol mol⁻¹ (−2.4 ± 0.4°C; SE). During the Bølling–Allerød, (+0.031 ± 0.021 mmol mol⁻¹; −0.6 ± 0.4°C; SE) and MWP1A (+0.040 ± 0.039 mmol mol⁻¹; −0.8 ± 0.8°C; SE) the mean values in Sr/Ca are statistically indistinguishable from the modern colonies (Figure 8a). Overall, these results align with previous SST reconstructions from Tahiti (Asami et al., 2009; Cohen & Hart, 2004; Felis et al., 2012) and suggest that last deglacial temperature variations, such as the Younger Dryas cooling, have had a global impact (Wang et al., 2001) and have also influenced the tropical South Pacific climate (Corrège et al., 2004). However, our results imply that the negative excursions in SST during the last deglaciation were less pronounced at Tahiti than at other locations in the Pacific. For instance, a negative excursion of ~−4°C relative to today has been reported for the Younger Dryas reconstructed from coral Sr/Ca records from Vanuatu in the western tropical South Pacific (Corrège et al., 2004). Furthermore, it needs to be mentioned that the slope of −0.050 ± 0.002 mmol mol⁻¹ C⁻¹ that is used in our study likely overestimates absolute SST values from skeletal Sr/Ca as a comparison between reconstructed SST from modern coral Sr/Ca seasonality and OISST v2.1 seasonality has shown (Figures 2a and 3d; Figure 6a). Furthermore, the coral Sr/Ca-SST relationship of our study (−0.050 mmol mol⁻¹ C⁻¹) is derived from a seasonal calibration. Other studies have applied rescaled seasonal calibrations for studying changes in mean SST, calibrations derived from bulk coral values for reef settings with different average temperatures, and annual mean calibrations using coral time series and have found steeper calibration slopes (−0.084 mmol mol⁻¹ C⁻¹, Gagan et al., 2012; −0.089 mmol mol⁻¹ C⁻¹, DeLong et al., 2010; −0.140 mmol mol⁻¹ C⁻¹, Felis et al., 2009). If these slopes are applied, warmer absolute values for last deglacial SST are reconstructed (Figure 4b). Therefore, absolute negative excursions in reconstructed deglacial SST presented in this study should be considered a lower limit of the magnitude of cooling.

Our results support that the Northern Hemisphere last deglacial temperature variations, such as the Younger Dryas cooling, have likely not similarly affected all regions of the South Pacific (Kiefer & Kienast, 2005). For instance, opposed to findings from this study for Tahiti and a previous study for Vanuatu (Corrège et al., 2004), coral-based-SST reconstructions for the Great Barrier Reef (Brenner et al., 2020; Felis et al., 2014), reconstructions of glacier retreat in the New Zealand Southern Alps (Koffman et al., 2017), and stable isotope stalagmite records from Borneo (Partin et al., 2007) revealed no evidence for the Younger Dryas cooling. The cause of these regional differences in South Pacific last deglacial climate variability is unclear. It is possible that climate conditions at the Great Barrier Reef and New Zealand were more strongly influenced by changes in the Southern Subtropical Front and less by changes in the South Pacific Subtropical Gyre, while the opposite holds for Tahiti and Vanuatu (Brenner et al., 2020; Strand et al., 2019). However, our results support that Northern Hemisphere climate signals that were mainly related to changes in the AMOC strength were rapidly propagated to the central tropical South Pacific by atmospheric teleconnections (McManus et al., 2004).

At Tahiti, Early Holocene coral mean Sr/Ca is relatively high (absolute SST low), similar to the Younger Dryas (Figure 8a). Low absolute SST reconstructions from coral Sr/Ca for the Early Holocene have been reported previously and have been explained by local cooling through upwelling (DeLong et al., 2010), local changes in ocean circulation, or a basin-wide negative SST excursion (Beck et al., 1997). Holocene temperature variability is widely debated and has likely not been globally uniform (Cartapanis et al., 2022; Kaufman et al., 2020; Liu et al., 2014). Therefore, although the precise mechanism is unknown it is possible that Early Holocene Tahitian climate variability did diverge from the temperature variability reconstructed from, for instance, ice core records (Figures 4a and 4c).
While seasonal variations in δ¹⁸Oseawater are not resolvable by applying the paired coral δ¹⁸O-Sr/Ca proxy at Tahiti (Cahyarini et al., 2008) this method was used to estimate changes in mean last deglacial paleo-SSS from the calculated δ¹⁸Oseawater contribution to coral skeletal δ¹⁸O (Δδ¹⁸O; Figure 8b). Following Cahyarini et al. (2008) we calculated Δδ¹⁸O from mean coral δ¹⁸O corrected for the ice-volume effect (Shackleton, 1967) and Sr/Ca for each age interval by using the Sr/Ca-SST slope from this study (−0.050 mmol mol⁻¹C⁻¹) and the δ¹⁸O-SST slope from Gagan et al. (1998; −0.180 ‰C⁻¹) that reflects the dependency between coral skeletal δ¹⁸O and SST without any contribution from δ¹⁸Oseawater. As no local δ¹⁸Oseawater-SSS calibration for Tahiti exists SSS was estimated from Δδ¹⁸O using the slope from Fairbanks et al. (1997; 0.273 ‰ PSU⁻¹) that has been universally applied across the tropical Pacific. However, it must be noted that the δ¹⁸Oseawater-SSS slope varies from 0.09 ‰ PSU⁻¹ in the eastern to 0.32 ‰ PSU⁻¹ in the western Pacific (Conroy et al., 2017) and has also changed substantially through time (Rohling & Bigg, 1998). Thus, the absolute SSS estimates presented here may be considered only an upper estimate of the deglacial salinity increase. Mean last deglacial skeletal Δδ¹⁸O is significantly elevated in all deglacial corals compared to modern corals (Figure 8b): +0.33 ± 0.18 ‰ (+1.2 ± 0.6 PSU; SE) during the Early Holocene, +0.38 ± 0.05 ‰ (+1.4 ± 0.2 PSU; SE) during the Younger Dryas, +0.77 ± 0.08 ‰ (+2.9 ± 0.3 PSU; SE) during the Bølling-Allerød, +0.96 ± 0.15 ‰ (+3.6 ± 0.5 PSU; SE) during MWP1A, and +0.54 ± 0.10 ‰ (+2.0 ± 0.4 PSU; SE) during Heinrich Stadial 1. Increased SSS and accompanied reduced rainfall rates during the last deglaciation align with a northward migration of the SPCZ at Tahiti. Also, a northward expansion of the South Pacific Subtropical Gyre with increased δ¹⁸Oseawater compared to the SPCZ may have contributed to the elevated deglacial δ¹⁸Oseawater. These results contrast with findings from a recent modeling study (Li & Liu, 2022) that has reported little change in rainfall in the (sub-)tropical South Pacific during the Younger Dryas and Heinrich Stadial 1. The elevated Δδ¹⁸O and SSS during the Bølling-Allerød and MWP1A may reflect increased evaporation rates at Tahiti due to elevated SST at that time. These increased evaporation rates may have amplified the increase in SSS due to the northward migration of the SPCZ and associated reduced rainfall rates.

Last deglacial Tahiti corals show a tendency toward reduced linear extension rates compared to modern and early Holocene corals (Figure 8c) which is likely a result of the less favorable thermal conditions during the last deglaciation. The high linear extension rates for Heinrich Stadial 1 are a result of the columnar, rapid growth pattern of coral 24A-11R-2 (Felis et al., 2012).

**5. Conclusions**

Modern *Porites* corals from Tahiti reefs were successfully used for Sr/Ca-SST calibrations (slope = −0.050 ± 0.002 mmol mol⁻¹C⁻¹) and showed considerable inter- and intra-colony variability in proxy seasonality and mean values. The latter demonstrates the necessity of averaging across a sufficient number of colonies for accurate coral-based SST or SSS reconstructions.

Sr/Ca seasonality from last-deglacial corals was in most cases not significantly different from modern corals (0.139 ± 0.010 mmol mol⁻¹; 2.8 ± 0.2°C, SE). This suggests that the reduced insolation seasonality during the last deglaciation had likely only a small impact on the SST seasonality that is below the detection limits of Sr/Ca-thermometry. However, the elevated Sr/Ca seasonality during the Younger Dryas (0.171 ± 0.017 mmol mol⁻¹; 3.4 ± 0.3°C; SE) points to an additional modulation of SST seasonality through regional processes, most likely a shoaling of the surface ocean mixed layer associated with a northward migration of the SPCZ. In the most extreme case, during the Younger Dryas, the SPCZ may not have formed a distinct southeast-oriented branch extending from the ITCZ as today but remained mostly a part of the ITCZ. Furthermore, an elevated coral δ¹⁸O seasonality during the Younger Dryas points to increased amplitudes of the seasonal δ¹⁸Oseawater cycle that may be a result of elevated δ¹⁸Oseawater relative to meteoric δ¹⁸O, or an increased discrepancy between rainfall in summer and winter. A seasonal cycle in skeletal δ¹³C that is likely related to seasonal variations in coral growth rate was observed across all corals with an elevated amplitude during the Younger Dryas. The latter may be explained by an increased discrepancy between summer and winter coastal runoff.

Last deglacial mean skeletal Sr/Ca revealed a considerable temperature variability at Tahiti during the deglacial transition ranging between +0.134 ± 0.012 mmol mol⁻¹ (−2.7 ± 0.2°C; SE) during the Younger Dryas and +0.031 ± 0.021 mmol mol⁻¹ (−0.6 ± 0.4°C; SE) during the Bølling-Allerød compared to modern. Latter demonstrates that the tropical South Pacific was affected by Northern Hemisphere processes, related to variations in AMOC strength, through atmospheric teleconnections. Last deglacial mean Δδ¹⁸O and consequently SSS was
generally elevated compared to the modern corals pointing to dryer conditions at Tahiti during the last deglaciation. Coral linear extension rates were reduced (0.98 ± 0.56 cm year⁻¹) compared to today (1.58 ± 0.31 cm year⁻¹) likely as a result of the reduced deglacial SST.

Overall, these findings demonstrate how last deglacial temperature fluctuations, such as the Younger Dryas, impacted the climate at Tahiti in the tropical South Pacific. Furthermore, inferred reduced deglacial rainfall rates suggest that the local climate at Tahiti was modulated especially by the meridional movement of the SPCZ that was likely located further northward during most of the last deglaciation. Especially, during the Younger Dryas, the SPCZ may have even merged with the ITCZ which is supported by an inferred shoaling of the surface mixed layer that is indicative of an increased influence of the South Pacific Subtropical Gyre at Tahiti. However, more studies are needed to determine the paleo-position of the SPCZ during the last deglaciation. Additionally, an enhanced seasonal coral skeletal δ¹⁸O and δ¹³C cycle during the Younger Dryas both point to a more extreme difference between summer and winter rainfall and associated runoff.

## Conflict of Interest

The authors declare no conflicts of interest relevant to this study.

## Data Availability Statement

Data files are publicly available at the World Data Center Pangaea at Knebel et al. (2024a) (deglacial corals) and Knebel et al. (2024b) (modern corals).

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## Acknowledgments

This research used samples provided by IODP, drilled on a mission-specific platform expedition conducted by the European Consortium for Ocean Research Drilling (ECORD) Science Operator (ESO). Funded by the Deutsche Forschungsgemeinschaft (DFG; German Research Foundation), project numbers 408139156 (T.F.), INST 247/889 FUGG (D.S.) and the Japan Society for the Promotion of Science (JSPS) KAKENHI, project number 22H01291 (R.A.). We thank Henning Kuhnert, Monika Segl, and their team for stable isotope measurements, Karl-Heinz Baumann and Petra Witte for SEM analyses, and Michael Zuther and Christoph Vogt for XRD measurements at MARUM and Faculty of Geoscience, University of Bremen, Germany. We thank Yasufumi Iryu for XRD measurements at Tohoku University and Masafumi Murayama for stable isotope analyses at Kochi University, Japan, M. Großkopf, M. Weber, and V. Blumrich for assistance during U-series dating at the University of Mainz, Germany, and Corinna Oster for compiling SST logger data and for her support during the initial data processing stage. We thank Joel Orempuller (IRD, Tahiti) for modern coral sampling, and Kristine DeLong, Louisiana State University, for providing geochemical data. Service d'Observation CORAIL from CRIOBE (Gilles Siu, Serge Planes) kindly provided temperature logger data from Tiahura (Moorea). Monthly mean data based on daily OISST v2.1 SST data were computed at and provided by the Integrated Climate Data Center (ICDC), CEN, University of Hamburg, Germany. Finally, we thank Judson Partin, University of Texas at Austin, 3 anonymous reviewers, and Associate Editor Kaustubh Thirumalai, University of Arizona, for their constructive suggestions that greatly improved the article. Open Access funding enabled and organized by Projekt DEAL.
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