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1 Data-Model Comparisons of Tropical Hydroclimate Changes Over the Common

2 A. R. Atwood1*, D. S. Battisti2, E. Wu2, D. M. W. Frierson2, J. P. Sachs3

3 1Florida State University, Dept. of Earth Ocean and Atmospheric , Tallahassee, FL, USA 4 2University of Washington, Dept. of Atmospheric , Seattle, WA 98195, USA 5 3University of Washington, School of Oceanography, Seattle, WA 98195, USA 6 7 Corresponding author: Alyssa Atwood ([email protected])

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9 Key Points: 10 • A synthesis of 67 tropical hydroclimate records from 55 sites indicate several regionally- 11 coherent patterns of change during the Common Era 12 • Robust patterns include a regional drying from 800-1000 CE and a range of tropical 13 hydroclimate changes ~1400-1700 CE 14 • Poor agreement between the centennial-scale changes in the reconstructions and transient 15 model simulations of the last 16

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17 Abstract 18 We examine the evidence for large-scale tropical hydroclimate changes over the Common Era 19 based on a compilation of 67 tropical hydroclimate records from 55 sites and assess the 20 consistency between the reconstructed hydroclimate changes and those simulated by transient 21 model simulations of the last millennium. Our synthesis of the proxy records reveal several 22 regionally-coherent patterns on centennial timescales. From 800-1000 CE, records from the 23 eastern Pacific and northern Mesoamerica indicate a pronounced drying event relative to 24 background conditions of the Common Era. In addition, 1400-1700 CE is marked by pronounced 25 hydroclimate changes across the tropics, including an inferred strengthening of the South 26 American summer monsoon, weakened Asian summer monsoons, and fresher conditions in the 27 Maritime Continent. We find notable dissimilarities between the regional hydroclimate changes 28 and large-scale temperature reconstructions with regard to the reversal of the Little Ice 29 (LIA) hydroclimate trends around 1600 CE. Apropos to previous interpretations of largescale 30 reorganization of the tropical Pacific climate system during the LIA, we do not find support for a 31 large-scale southward shift of the Pacific Intertropical Convergence Zone, while evidence for an 32 equatorward contraction of the monsoonal Asian-Australian rainbelt and/or a strengthened 33 Pacific Walker Circulation exists from limited geographic regions. Transient climate model 34 simulations exhibit weak forced long-term tropical rainfall changes over the last millennium that 35 poorly agree with the reconstructions, suggesting that either the reconstructed tropical 36 hydroclimate changes were unforced, or the models are unable to capture the forcings, 37 feedbacks, or long-term hydroclimate responses over the last millennium. 38

39 1. Introduction 40 The large-scale circulation of the tropical atmosphere represents an essential component 41 of the global climate system. The mean meridional Hadley circulation, which governs the zonal 42 mean distribution of tropical rainfall, transports heat and moisture between the tropics and 43 subtropics and plays an integral role in regulating hemispheric energy budgets on seasonal to 44 orbital timescales (Donohoe et al., 2013; Kang et al., 2008). Tropical rainfall patterns are also 45 governed by the zonal Walker and monsoonal circulations, which exist due to the presence of 46 continents. These zonal asymmetries render the dynamics of precipitation in regions of deep 47 convection (e.g. the Western Pacific Warm Pool) to be largely distinct from regions 48 characterized by strong sea surface temperature (SST) gradients and narrow rainbands (i.e., the 49 Intertropical Convergence Zones [ITCZs]). The interplay of these large-scale circulation features 50 dictate the seasonal and -to-year variations in tropical rainfall that billions of people around 51 the globe depend on for food and water. 52 How tropical rainfall patterns will respond to anthropogenic climate change is still 53 uncertain, due to the short (spanning roughly half a ) and unevenly distributed 54 precipitation records around the globe (Biasutti et al., 2018). However, the efficacy of the 55 climate models used to project precipitation can be tested against records. In 56 particular, the last 2,000 , known as the Common Era, has been a major reconstruction 57 target for the paleoclimate science community over the last decade (Consortium, 2013; 58 Consortium, 2019; PAGES 2k-PMIP3 group, 2015; Konecky et al., 2020; Tierney et al., 2015) 59 given its relative abundance of data and the similarity of the Earth’s boundary conditions to our 60 climate.

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61 Global compilations of temperature reconstructions of the Common Era generally 62 indicate a long-term cooling trend over the last millennium that extended into the middle of the 63 nineteenth century and was largely driven by increased volcanic activity (Atwood et al., 2016; 64 Schurer et al., 2014), though greenhouse gas, solar, land use, and orbital forcings, and internal 65 variability are also thought to have played an important role in driving temperature variations 66 over the last millennium (Otto-Bliesner et al., 2016; Schurer et al., 2013; Swingedouw et al., 67 2011). Indeed, the spatial and temporal heterogeneity of the LIA points to a complex array of 68 forcings, feedbacks, and internal variability operating in the climate system at that 69 (Fernandez-Donado et al., 2013; Kaufman et al., 2009; Lehner et al., 2013). Recent compilations 70 of temperature-sensitive proxy data suggest that global mean surface temperatures cooled by 71 approximately 0.2 °C over the last millennium (Consortium, 2013; Consortium, 2019) with 72 larger cooling in the Northern Hemisphere (Hegerl et al., 2007; Mann et al., 2009; Marcott et al., 73 2013; Moberg et al., 2005) than the Southern Hemisphere (Neukom et al., 2014). Embedded 74 within this global cooling trend, the Little Ice Age (LIA) has been characterized as a period of 75 regionally-variable cooling during the second half of the last millennium ~1450-1850 76 (Consortium, 2013; Cunningham et al., 2013; Kobashi et al., 2011; Masse et al., 2008; Neukom 77 et al., 2014). Evidence for amplified cooling in central Greenland (exceeding 1.0 °C) has 78 supported interpretations of Arctic amplified cooling during the LIA (Kobashi et al., 2011), 79 which is also seen in transient climate model simulations of the last millennium (Atwood et al., 80 2016). 81 In the tropics, widespread changes in rainfall patterns are thought to have occurred during 82 the LIA. In Asia, weakened summer monsoon rainfall and a prevalence of extreme droughts 83 around 1300-1700 CE have been inferred from oxygen isotopes in cave deposits in India and 84 , from tree ring records from Southeast Asia, and from relative abundances of foraminifera 85 from the Arabian Sea (Anderson et al., 2002; Buckley et al., 2010; Sinha et al., 2011; Wang et 86 al., 2005; Zhang et al., 2008). Some studies have suggested that these rainfall changes 87 contributed to the collapse of several major civilizations in India, China and Southeast Asia 88 (Buckley et al., 2010; Sinha et al., 2011; Zhang et al., 2008). On the opposite side of the Pacific 89 in the Southern Hemisphere, an intensified South American summer monsoon ca. 1300–1500 90 CE, continuing through 1800 CE, has been inferred from records of oxygen isotopes in cave 91 deposits, glacial ice, and authigenic calcite in lake sediments (Vuille et al., 2012). 92 Outside of these major monsoon systems, the nature of LIA hydroclimate changes in 93 other regions of the tropics are more heavily disputed, particularly in and around the tropical 94 Pacific Ocean. For example, different compilations of hydroclimate records from the Pacific 95 basin and surrounding regions have been interpreted as representing either: 1) a southward 96 migration of the mean annual position of the Pacific Intertropical Convergence Zone (ITCZ) 97 (Nelson and Sachs, 2016; Richey and Sachs, 2016; Sachs et al., 2009) and Atlantic ITCZ (Haug 98 et al., 2001), 2) an equatorward contraction of the Asian-Australian monsoons (Denniston et al., 99 2016; Griffiths et al., 2016; Yan et al., 2015), and/or 3) a strengthened Pacific Walker 100 Circulation (Griffiths et al., 2016; Yan et al., 2011; Yan et al., 2015). Importantly, however, the 101 definition of the LIA and the reference period varies widely across these studies; in some cases, 102 the interpretation was based on trends during the LIA (e.g. Richey and Sachs, 2016), in other 103 cases LIA conditions were referenced to the MCA (Haug et al., 2001; Nelson and Sachs, 2016) 104 or to the modern period (Sachs et al., 2009), and in yet other cases, the MCA and modern periods 105 were used interchangeably, depending on data availability (Yan et al., 2015). Standardizing these

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106 comparisons is critical to developing more robust constraints on the regional signatures of 107 tropical hydroclimate changes during the last millennium. 108 Further complicating the interpretation of the LIA hydroclimate changes, theoretical 109 frameworks to support the proposed rainfall changes are not well established. One theory links a 110 southward migration of the zonal mean Hadley circulation to greater cooling of the Northern 111 Hemisphere than the Southern Hemisphere during the LIA (Chiang and Bitz, 2005; Hwang et al., 112 2013; Kang et al., 2008). However, the applicability of this zonal mean energetic framework (i.e. 113 the response of the zonal mean atmospheric circulation to changes in the hemispheric energy 114 budget) to the regional scale of the paleoclimate records is unclear, given the relatively weak 115 interhemispheric temperature changes during this time, and the mounting evidence that shifts of 116 the zonal mean ITCZ are highly zonally variable even under a large interhemispheric asymmetry 117 in forcing (Atwood et al., 2020; Roberts et al., 2017). 118 Over the past decade there has been a substantial increase in the number of high- 119 resolution, hydroclimate-sensitive proxy records that are publicly accessible under improved 120 paleo-data stewardship efforts, an effort that has accelerated over the past few years, resulting in 121 several data sets that post-dated the data collection and analysis aspects of the current study 122 (Atsawawaranunt et al., 2018; Konecky et al., 2020). Here, we collected public data to 123 reconstruct robust spatiotemporal patterns of tropical hydroclimate change over the Common Era 124 using a larger suite of hydroclimate records than has previously been considered. Furthermore, 125 we compare the reconstructed hydroclimate patterns to a large suite of transient climate model 126 simulations of the last millennium to shed light on the mechanisms of the long-term 127 hydroclimate changes, including distinguishing between forced and unforced modes of 128 variability.

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130 2. Materials and Methods

131 2.1. Proxy Record Selection and Treatment

132 Tropical hydroclimate records spanning the last millennium were selected from the 133 NCDC/NCEI (https://www.ncdc.noaa.gov/) and PANGAEA (https://www.pangaea.de/) 134 paleoclimate data repositories. A total of 67 speleothem, lake and marine sediment, and tropical 135 glacier records were selected based on the following criteria: 1) the original authors interpreted 136 the proxies as representative of hydroclimate conditions, 2) the records were located between 137 35°S to 35°N, and 3) the records spanned at least 70% of the last millennium (850-1850 CE) 138 with a minimum temporal resolution of 200 years and contained at least one age control point 139 within that period. More stringent record selection criteria were adopted for the cluster analysis 140 and Empirical Orthogonal Function analysis, as detailed in Sections 2.2.1-2.2.2. For all analyses, 141 the records were truncated outside of the time period 350-1950 CE and standardized to Z-scores 142 by subtracting the mean and dividing by the standard deviation of the data during this interval. 143 The Z-scores from each record (with zero mean and unit variance) were binned into 100-yr 144 windows to focus on the centennial to millennial scale patterns in the data, and to limit the 145 sensitivity of the results to the age model uncertainties in the reconstructions. The sign of the 146 inferred relationship between rainfall amount and the measured values of the proxies was 147 adopted from the original publication of each record. A lake minerology record from Kiritimati

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148 Atoll (Higley et al., 2018) was omitted from our analyses due to a substantial revision in the age 149 model between 1200 CE and the core top (J. Conroy, personal communication). Table 1, Fig. 1, 150 and Table S1 provide a summary of the proxy records.

151 Of the 67 records in the compilation, 24% (16 out of 67) are non-water isotope records, 152 consisting of sediment grain size, bulk density, magnetic susceptibility, color, elemental 153 composition (Ti/Ca, percent Mg in calcite), δ13C of organic matter, and relative diatom 154 abundances in lake and marine sediments (see Table S1). The remainder of the records (51 out of 155 67; 76%) are water isotope-based archives, which reflect the isotopic composition (δ2H or δ18O) 156 of ground water, lake water, or near-surface seawater. The stable isotopic composition of water 157 is a powerful tracer of the global hydrologic cycle, as hydrologic processes such as evaporation 158 and precipitation, advective transport in the atmosphere and ocean, and land processes affect the 159 isotope composition of water. In this study, we combine isotopic and non-isotopic archives of the 160 hydrologic cycle, with the assumption that the water isotope archives reflect the regional 161 hydrologic budget and that lower (higher) δ2H and δ18O values occur during periods of 162 regionally wetter (drier) conditions. This assumption is supported by the sensitivity of meteoric 163 water to the regional (and sometimes local) hydrologic balance, and by the “amount effect”, 164 which is an empirical relationship in which the isotopic composition of precipitation in the 165 tropics is negatively correlated with the amount of precipitation on monthly or longer timescales 166 (Dansgaard, 1964; Risi et al., 2008). We recognize that many different components of the 167 hydroclimate system are represented in our diverse multi-proxy network. The isotopic records 168 from closed lake systems likely reflect the water balance of the lake (i.e. precipitation minus 169 evaporation), while open lake, speleothem, and glacier ice records likely reflect the isotopic 170 composition of precipitation. Seawater isotope records in the near-surface ocean, meanwhile, 171 behave similarly to salinity and provide a means of characterizing the surface exchange of 172 freshwater and water mass mixing. Future work should incorporate isotope-enabled model 173 simulations to explicitly track water isotope ratios in different components of the hydrosphere 174 and thereby provide more quantitative comparisons between the water isotope-based proxy 175 records and model simulations.

176 2.2. Proxy Data Analyses

177 2.2.1. Hierarchal Cluster Analysis

178 Hierarchal cluster analysis (HCA; Eisen et al., 1998) was used to organize the 179 paleoclimate records into groups with similar temporal patterns. In HCA, a binary, hierarchical 180 cluster tree (dendrogram) links pairs of records that are in close proximity based on a specified 181 distance metric; binary clusters are then grouped into larger clusters, forming a hierarchical tree. 182 The height of the linkages in the dendrogram represents the distance between objects. HCA has 183 been successfully used to divide diverse multi-proxy paleoclimate datasets into groups with 184 shared temporal patterns (Kaufman et al., 2016; Shuman et al., 2018). We constructed the 185 hierarchical cluster tree in MATLAB (clusterdata.m) based on Euclidian pairwise distances 186 between records and Ward’s minimum variance method with a maximum of three clusters 187 specified. For studies in which more than one dataset was generated to reconstruct hydroclimate 188 at a given site, all relevant datasets from that study were averaged together. Records with 189 missing data in more than two 100-year windows across the 1600-year (350-1950 CE) period of 190 analysis were excluded from the cluster analysis; records with missing data in £ 2 windows were

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191 linearly interpolated and extrapolated. The Sletten et al. (2013) speleothem d18O record from 192 Namibia and the Stansell et al. (2012) ostrocod d18O record from Nicaragua were omitted to 193 improve the stability of the cluster analysis. After these adjustments, 41 unique records remained 194 in the cluster analysis. Composite records for each cluster were calculated by averaging all 195 binned data from each cluster. The 95% confidence intervals were calculated based on 1000 196 iterations of bootstrap sampling with replacement.

197 2.2.2. EOF/PC Analysis

198 Empirical Orthogonal Function (EOF)/Principle Component (PC) Analysis was 199 employed to characterize the dominant modes of variability in both the proxy records and models 200 and to compare the patterns of variability between the two. These analyses were conducted on 201 the full tropical dataset, as well as selected regional subsets of the data. The record selection 202 criteria matched that of the cluster analyses in Section 2.2.1. To facilitate data-model 203 intercomparison, the proxy data were truncated between 850-1850 CE, converted to Z-scores, 204 and binned into 100-yr windows (as in the cluster analysis) before calculating the PCs and EOF 205 loadings from the proxy data. Regional subsets selected for analysis were Asia, South America, 206 and the western Pacific. For the model simulations, precipitation and surface air temperature data 207 from the individual models and multi-model means were binned into 100-yr windows and the 208 EOFs and PCs calculated from the regions bounded by the regional subsets of proxy data.

209 2.2.3. Regional Composites

210 Regional composites of the proxy data were constructed to identify the robust regional 211 hydroclimate signals in the reconstructions. For selected regions, Z-scores of the 100-yr binned 212 proxy data were plotted with the average and standard deviation of the regional composites. For 213 studies that invoked more than one dataset to reconstruct hydroclimate at a given proxy site, all 214 relevant datasets from that study were averaged together.

215 2.2.4. Analysis

216 To compare the proxy-reconstructed hydroclimate changes over the last millennium to 217 climate model output, all data points in each standardized proxy record were averaged across the 218 LIA (1450-1850 CE) and MCA (950-1200 CE). An unpaired 2-sample student t-test was 219 performed to assess the significance of the change between these time periods (with degrees of 220 freedom equal to the number of data points in each period).

221 2.3. Model Simulations

222 We include two sets of model simulations in our analyses. The first set of simulations are 223 six forced transient last millennium simulations that were performed in association with the 224 Coupled Model Intercomparison Project Phase 5 (CMIP5)/Paleoclimate Intercomparison Project 225 Phase 3 (PMIP3). These models are CCSM4 (Gent et al., 2011; Landrum et al., 2013); GISS-E2- 226 R forcing ensemble members r1i1p121 and r1i1p124 (Schmidt et al., 2006), MPI-ESM-P 227 (Marsland et al., 2003; Raddatz et al., 2007), CSIRO Mk3L version 1.2 (Phipps et al., 2012; 228 Rotstayn et al., 2012), and HadCM3 (Collins et al., 2001; Pope et al., 2000; Schurer et al., 2013). 229 The second set of simulations are from the Community Earth System Model (CESM) Last

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230 Millennium Ensemble (LME) version 1.1 (with version 5 of the Community Atmosphere Model; 231 CAM5), and include a set of 13 experiments with all last millennium climate forcings included 232 (orbital, volcanic, solar, and greenhouse gas), and the following single-forcing experiments: 233 volcanic-only (n=5), solar-only (n=4), orbital-only (n=3), greenhouse gas-only (n=3) and land- 234 use only (n=3).

235 Two different reconstructions of volcanic aerosols and three different methods for 236 implementing aerosols in the models were used in the CMIP5/PMIP3 and CESM simulations 237 analyzed in this study. The volcanic aerosol reconstructions were taken from either (1) the Gao et 238 al. (2008) dataset of sulfate loading, where the sulfate loading is a function of month, latitude in 239 10 bands, and height from 9 to 30 km at 0.5-km resolution, or (2) the Crowley et al. (2008) 240 dataset of aerosol optical depth (AOD) (Schmidt et al., 2011). The CCSM4 and CESM 241 simulations prescribed sulfate loading from Gao et al. (2008), and stratospheric aerosols are 242 prescribed in the models as a fixed single-size distribution in the three layers in the lower 243 stratosphere above the tropopause (Landrum et al., 2013; Otto-Bliesner et al., 2016). The GISS 244 121, GISS 124, MPI, and HadCM3 simulations are forced by a prescribed aerosol effective 245 radius based on the AOD from Crowley et al. (2008), which is a function of height and latitude 246 in the atmosphere. In the CSIRO model, volcanic forcing was imposed as a globally averaged 247 perturbation to total solar irradiance (TSI) that scales with the Crowley et al. (2008) dataset of 248 AOD (Masson-Delmotte, 2013).

249 In comparing reconstructed to simulated hydroclimate changes over the last millennium, 250 we generally avoid making point-to-point comparisons between the proxy data and models, 251 given the structural uncertainties in simulated precipitation climatology associated with the 252 coarse spatial resolution and large tropical mean state biases in the models (e.g. PAGES Hydro2k 253 Consortium, 2017). In addition, most of the proxy records included in this analysis are based on 254 water isotopes, hence they record regional, as well as local, hydroclimate variations (e.g. water 255 isotope records from the Andes reflect changes in rainfall and upstream rainout over the Amazon 256 Basin (Vuille et al., 2012), while water isotope records from Chinese speleothems are sensitive 257 to changes in moisture source, trajectory, and upstream rainout over South Asia (Pausata et al., 258 2011)). For these reasons, we compare networks of proxy reconstructions to simulated 259 precipitation over a regional scale whenever possible.

260 2.4. Energy Budget Analysis

261 In order to quantify the hemispheric asymmetry of the forcings and feedbacks in the last 262 millennium simulations during the LIA relative to the MCA, we employed the Approximate 263 Partial Radiative Perturbation (APRP) method (Taylor et al., 2007). The APRP method is based 264 on a single-layer, shortwave radiative model of the atmosphere and has been successfully used to 265 quantify the strength of climate forcings and feedbacks in a wide variety of paleoclimate, 266 historical, and future climate simulations (e.g. Atwood et al., 2016; Crucifix, 2006; Hwang et al., 267 2013). In the APRP method, the influence of changes in surface albedo, shortwave absorption, 268 and scattering on the TOA energy budget are diagnosed at every model grid cell using all-sky 269 and clear-sky model output. Changes in top-of-the-atmosphere (TOA) shortwave energy fluxes 270 were decomposed into individual climate forcing and feedback terms as in Atwood et al. (2016), 271 except that here the analysis is performed separately for each hemisphere. Energy fluxes are 272 reported as NH minus SH, where negative terms indicate more NH cooling and positive terms

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273 indicate more SH cooling. Radiation data for CSIRO and the CESM LME simulations were not 274 available at the time of analysis and thus these simulations were omitted from the APRP 275 analyses.

276

277 3. Results

278 3.1. Cluster Analysis of Proxy Records

279 The hierarchal cluster analysis indicates that the most common pattern, found in 41% of 280 the tropical hydroclimate records analyzed, is a millennial-scale isotopic depletion trend 281 beginning around 1000 CE and reversing around 1800 CE (Cluster 2; Fig. 2B-C). This pattern 282 appears in nearly every region represented in the proxy compilation and tracks global 283 temperature trends over the last millennium (Fig. 3H; Consortium, 2019). An EOF analysis of 284 the reconstructions also yields this pattern in EOF1/PC1 (Fig. S1A). We will demonstrate that 285 this signal is not the dominant pattern at regional scales, and thus it is likely that it does not 286 solely reflect precipitation changes but includes an important contribution from changes in 287 tropical evaporation rates as well.

288 The next most common pattern, found in 37% of the records in the cluster analysis, 289 including the majority of records from the Maritime Continent and the majority of records from 290 southwestern tropical South America (including the central Andes and the southern Amazon), is 291 dominated by a drying (and/or isotopic enrichment) trend from 1600 CE to present (Cluster 1; 292 Fig. 2D,E). An opposing trend is observed in the third cluster, which contains 22% of the 293 records, including those from northern and eastern South America, and parts of Mesoamerica 294 and East Asia (Cluster 3; Fig. 2F,G). In this cluster, the dominant pattern is drying (and/or 295 isotopic enrichment) beginning ca. 1300 CE and peaking during the LIA around 1600-1700 CE, 296 followed by a wetting trend to present.

297 3.2. Regional Composites of Proxy Records

298 When the records are compiled by region, several robust regionally-coherent 299 hydroclimate patterns emerge that provide more insight into the patterns found in the cluster 300 analysis. From 800-1000 CE, encompassing the early part of the MCA, three lake records from 301 two islands in the eastern equatorial Pacific indicate notably dry (and/or isotopically enriched) 302 conditions (Fig. 3A), a pattern that is also found in half (5/10) of the records from Mesoamerica, 303 especially those clustered in the northern part of this region (Fig. 3B). The remaining 5/10 304 Mesoamerican records, which extend father southward, indicate a distinctly different centennial 305 pattern, characterized by a steady drying (and/or isotopic enrichment) trend from 1200 CE to 306 present (Fig. 3C).

307 During the second half of the last millennium, records from the major monsoon systems 308 of Asia and South America demonstrate strong regional coherency. The South American records 309 are characterized by an opposing pattern between the southwestern sites (i.e. the central Andes 310 and southern Amazon) and the northeastern sites. Nearly all records from the southwest indicate 311 anomalously wet (and/or isotopically depleted) conditions ca. 1500-1700 CE, followed by a

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312 drying trend to the present (Fig. 3E), while the northeastern records indicate anomalously dry 313 conditions beginning around 1500 CE and extending into the present (Fig. 3D).

314 Meanwhile, in the South and East Asian monsoon regions, nearly all records demonstrate 315 anomalously dry (and/or isotopically enriched) conditions ca. 1400-1700 CE followed by a 316 coherent wetting (and/or isotopic depletion) trend to present (Fig. 3F). An opposing pattern is 317 found in a composite of the marine records from the Maritime Continent, which indicates 318 isotopically-depleted seawater (and by inference lower salinity) from 1500-1700 CE (Fig. 3G). A 319 single speleothem record from northern also reflects the Asian monsoon trends to some 320 extent, with drier (and/or isotopically enriched) conditions ca. 1400-1600 CE, relative to much of 321 the prior 500 years, followed by a wetting (and/or isotopic depletion) trend to present (Fig. S2C). 322 These features provide some evidence of coherency between the northern and southern margins 323 of the Asian-Australian rainbelt. However, while the LIA conditions in Asia are anomalous with 324 respect to the background conditions of the Common Era, this is not the case for the northern 325 Australian record. More hydroclimate records from the northern Australian are needed to 326 confirm these regional patterns in this area, as only a single record that spanned the last 327 millennium was publicly available at the time of analysis.

328 3.3. Data-Model Intercomparison

329 To investigate the mechanisms of the robust regional hydroclimate patterns observed in 330 the proxy records, we compared the reconstructions to a large suite of transient climate model 331 simulations of the last millennium. While the models reproduce the reconstructed hemispheric- 332 scale, long-term temperature trends with good fidelity over the last millennium (Atwood et al., 333 2016), the dominant patterns of centennial hydroclimate variability in the models poorly agree 334 with the proxy data, when individual models are considered (Fig. 6) and in the multi-model mean 335 (Fig. 4; Fig. S1).

336 The only broadly similar feature between the models and proxy data is the long-term 337 drying trend in South and East Asia over the last millennium. However, the temporal evolution 338 of this feature is notably different between the proxies and models. The coherent large-scale 339 drying pattern in the Asian proxy records (namely speleothem and lake sediment records) over 340 the last millennium becomes more apparent when the EOF analysis is performed on the Asian 341 records exclusively (Fig. S3A). The drying pattern in EOF1 of the Asian records (which explains 342 46% of the variance) peaks around 1600 CE, followed by a reversal to the modern era. In 343 contrast, in 5/6 PMIP3 models (data not shown) and the multi-model mean (Fig. S3B), EOF1 344 (which explains 64% of the variance) is characterized by a persistent drying trend over Asia 345 throughout the last millennium. This persistent drying trend in South and East Asia in the models 346 is also apparent in EOF1 of the full tropical precipitation field (Fig. S1B).

347 Epoch analysis of the LIA (1450-1850) and MCA (950-1200 CE) differences further 348 underscore the lack of agreement in long-term precipitation changes in the models and the 349 hydroclimate reconstructions in the regions where proxy records exist (Fig. 4), with the 350 exception of drying in southern and eastern Asia. The data/model discrepancies are most 351 pronounced over tropical South America and the Maritime Continent. Over South America, the 352 simulated and reconstructed hydroclimate changes are both robust, but show little agreement 353 with one another (Fig. 4). The robust LIA antiphased dipole pattern over South America in the

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354 reconstructions contrasts strongly with the last millennium simulations, which are generally 355 characterized by increased rainfall in the northeast and little change over the central Andes and 356 southern Amazon (Fig. 4; Fig. 6A,B). Notably, the models do not simulate a strengthening of the 357 South American Summer Monsoon (SASM) during the LIA, and few coherent rainfall changes 358 are simulated over South America during the height of the monsoon season in austral summer 359 (DJF; Fig. S4A,C).

360 The CMIP5/PMIP3 and CESM simulations also provide little insight into the Pacific 361 hydroclimate reconstructions during the last millennium. There is poor agreement between the 362 proxy data and models in the Maritime Continent and in the Pacific cold tongue. In addition, the 363 simulated tropical Pacific rainfall changes in the CMIP5/PMIP3 simulations are notably different 364 from the CESM simulations (Fig. 4). These differences are associated with differing long-term 365 changes in the tropical Pacific sea surface temperatures in the two models (Fig. S5). In the 366 CMIP5/PMIP3 models, the rainfall changes are consistent with a weakened local Hadley 367 circulation, with drying along the Pacific ITCZ and wetting in parts of the subtropical and 368 equatorial Pacific (Fig. 4). In contrast, the CESM LME simulations are generally characterized 369 by a strengthened Pacific Walker Circulation, though the lack of stippling in much of the tropical 370 Pacific (and Indian Ocean) in the CESM ensemble mean indicates that the internal variability of 371 Indo-Pacific precipitation is generally larger than the forced signal in that model.

372 The forced response of tropical rainfall to the last millennium forcings is weak on 373 centennial timescales in the models, as indicated by the substantial differences in tropical rainfall 374 changes across different models and across different ensemble members of the same model, 375 particularly in the tropical Pacific (Fig. 4A-D; Fig. S6). These differences result in small (< 10%) 376 LIA precipitation changes relative to the models’ climatology in the PMIP3 multi-model mean 377 (Fig. 4B) and in the CESM ensemble mean (Fig. 4D).

378 Despite the weak amplitude of the response, in some regions of the tropics, the models 379 demonstrate a similar forced precipitation response. During the LIA relative to the MCA, nearly 380 all CMIP5/PMIP3 models and CESM ensemble members simulate large-scale drying in South 381 and East Asia, an equatorward shift in precipitation over the Sahel and the equatorial Atlantic 382 Ocean, wetting in the far northeast of tropical South America, drying in parts of the southwest 383 tropical Pacific (in the region of the South Pacific Convergence Zone), and drying of the mid- 384 and high latitudes (Fig. 4). These features are qualitatively similar between the first and second 385 half of the LIA, but are more pronounced in the second half of the LIA (Fig. S7). The CESM 386 single-forcing experiments indicate that the tropical rainfall changes are largely driven by orbital 387 forcing in the models, with drying over Asia also driven by volcanic forcing (Fig. 5). However, 388 the poor agreement between data and models indicates that the sensitivity of tropical rainfall to 389 orbital forcing in the models likely has little-to-no basis in the proxy reconstructions. These 390 findings highlight a divergence between the large-scale temperature response and the tropical 391 rainfall response to forcing in the models: while the models reasonably reproduce the long-term 392 global temperature trends over the last millennium (Atwood et al., 2016; Neukom et al., 2014),

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393 the simulations strongly diverge from the reconstructions when it comes to changes in tropical 394 precipitation.

395

396 4. Discussion

397 Three robust signatures of tropical hydroclimate change over the Common Era have 398 emerged from our analyses of the centennial patterns in the proxy records. These are: (1) a pan- 399 tropical isotopic depletion trend in the water isotope based records over the past millennium, (2) 400 anomalously dry conditions in the eastern tropical Pacific Ocean and Mesoamerica between 800- 401 1000 CE, and (3) a range of hydroclimate anomalies across the tropics from 1400-1700 CE, 402 including drier and/or isotopically enriched conditions in South and East Asia, wetter and/or 403 isotopically depleted marine conditions in the Maritime Continent, and a dipole pattern in South 404 America. We discuss each of these robust regional patterns in more detail, below.

405 4.1. Isotopic Depletion Trend in Pan-Tropical Records Over the Last Millennium

406 The dominant pattern of centennial hydroclimate variability, found in 41% of the tropical 407 hydroclimate records analyzed (14 out of 17 of which are based on water isotopes) and appearing 408 in nearly every region of the tropics represented in the proxy compilation, is an isotopic 409 depletion trend beginning around 1000 CE and reversing around 1800 CE (Fig. 2A-C), which 410 tracks global temperature trends over the Common Era (Fig. 3H; PAGES2k Consortium, 2013; 411 PAGES2k Consortium, 2019; Emile-Geay et al., 2017). The global cooling trend over the last 412 millennium, which ended at the start of the industrial era, is thought to largely be the result of 413 increased volcanic activity, with a secondary role for changes in solar irradiance, greenhouse gas 414 concentrations, and land use changes (Atwood et al., 2016; Schurer et al., 2014).

415 In model simulations, a perceptible but very weak decrease in tropical area-averaged 416 precipitation occurs over the last millennium (Fig. S8B; DPLIA - MCA = -0.28 ± 0.08% [1s] and 417 -0.23 ± 0.04% for the CMIP5/PMIP3 LM and CESM LME runs, respectively) in association 418 with global cooling of 0.2-0.3 °C (Fig. S8A; Atwood et al., 2016). This decrease in simulated 419 tropical precipitation would be far too small to be detectable in non-water isotope based proxy 420 archives (which indirectly reflect changes in precipitation), or in water isotope based records via 421 the amount effect. However, the water isotope tracers are sensitive to evaporative fluxes (which 422 are the source of the atmospheric water vapor pool from which all precipitation is derived) in 423 addition to precipitation. Thus, the millennial-scale isotopic depletion trend observed in many of 424 the proxy records may be associated with decreased evaporation rates (and increased 425 fractionation during evaporation) under cooler tropical sea surface temperatures (Risi et al., 426 2010). This hypothesis could be tested using model simulations with water isotope tracers, and 427 the response of the global water cycle to global and regional cooling over the last millennium—

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428 and how those changes are recorded in water isotope tracers—is worthy of further investigation 429 using such tools.

430 4.2. Dry Conditions in the Eastern Pacific and Mesoamerica ~800-1000 CE

431 One of the largest regional hydroclimate signals to emerge from our analysis is the 432 pronounced drying event from 800-1000 CE recorded in a series of lake sediment records from 433 the Galápagos Islands in the eastern equatorial Pacific (Fig. 3A; 1˚S, 90-91˚W), and in many lake 434 and speleothem records from Mesoamerica (Fig. 3B). This signal is evident a variety of different 435 types of hydroclimate records, including both minerology and water isotope based sedimentary 436 records, suggesting that it is driven by a change in precipitation. Lake sediment records are 437 sensitive to precipitation minus evaporation (P-E) while speleothem records are largely 438 controlled by the isotopic composition of precipitation, which are often interpreted in terms of 439 the ‘amount effect’ (Konecky et al., 2020). Evidence of pronounced drying during this interval 440 has also been found in a sediment minerology record from a hypersaline lake in Kiritimati in the 441 central equatorial Pacific (Higley et al., 2018). This event can be seen in the cluster analysis in 442 Cluster 2 (Fig. 2A,B), which contains a number of records from the eastern Pacific and 443 Mesoamerica (Fig. 2C).

444 This dry interval is consistent with prior studies that have found robust evidence of 445 extended drought across Mesoamerica from 800-1050 CE, which has been implicated in major 446 social disruptions in pre-Columbian Mesoamerican civilizations (e.g. Bhattacharya et al., 2017; 447 Hodell et al., 2005a). The consistency in timing between the dry intervals in Mesoamerica and 448 the eastern tropical Pacific suggests a dynamical linkage between them, possibly associated with 449 a coordinated response of the eastern Pacific ITCZ and Mesoamerican monsoon to the interbasin 450 Atlantic-Pacific SST gradient during this time. Indeed, Bhattacharya et al. (2017) and 451 Bhattacharya and Coats (2020) suggest that the dry interval in Mesoamerica may have been 452 driven by cooler northern tropical Atlantic sea surface temperatures, strengthened easterly trade 453 , and increased moisture divergence over the Caribbean and Mesoamerica. If this cooling 454 signal extended into the eastern tropical Pacific via the -evaporation-SST feedback (Xie and 455 Carton, 2004), such tropical Atlantic cooling could provide a plausible mechanism by which the 456 prolonged drought conditions in the eastern Pacific and Mesoamerica were linked.

457 4.3. Widespread Tropical Hydroclimate Changes during the LIA ~1400-1700 CE

458 The second half of the last millennium is characterized by a set of pronounced and 459 synchronous hydroclimate changes throughout the tropics around 1400-1700 CE, including: 1) 460 isotopically-enriched (drier) conditions in South and East Asia, 2) isotopically-depleted (fresher) 461 surface ocean conditions in the Maritime Continent, and 3) a dipole pattern in South America. 462 We investigate each of these regional signals and discuss their possible interpretations, below.

463 4.3.1. Dry/Isotopically-Enriched Conditions in South and East Asia

464 Hydroclimate reconstructions from South and East Asia generally indicate dry 465 (isotopically-enriched) conditions from 1400-1700 CE relative to present and to the background 466 conditions of the Common Era, followed by progressively wetter (isotopically-depleted) 467 conditions to present (Fig. 3F). LIA drying in this region is also apparent in the LIA-MCA epoch

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468 analysis (Fig. 4), and in the EOF/PC analysis (EOFs 1-2 in Fig. S3A). Because all of the proxy 469 records from Asia are water isotope (speleothem and lake) records (Fig. 3F), which integrate 470 climate information on large spatial scales, the widespread signature of isotopic enrichment in 471 these records likely reflective of reduced upstream rainout associated with a weakened South, 472 and possibly East Asian, monsoon during the LIA.

473 This interpretation is consistent with past studies that have characterized the LIA as a 474 time of weakened South and East Asian summer monsoon strength, marked by a series of 475 monsoon droughts in India, China, and southeast Asia (Sinha et al., 2011; Zhang et al., 2008). 476 Proposed mechanisms for the weakened monsoons include widespread extratropical NH cooling 477 that drove a southward shift of the rain belt in the Indian Ocean (Sinha et al., 2011) and reduced 478 land-sea temperature contrast between Asia and the tropical Indo-Pacific (Zhang et al., 2008). 479 We do indeed see a weakening of the South and East Asian monsoons in the model simulations 480 of the last millennium in association with enhanced cooling over land, particularly in the NH 481 (Fig. 4; Fig. S9). In fact, largescale drying over South and East Asia is one of the only 482 hydroclimate features in the reconstructions that is captured to some extent by the climate 483 models. However, the temporal pattern of these changes is notably different between the 484 reconstructions and models. In the reconstructions, the drying signal peaks around 1600 CE, 485 before reversing to the present era, while in the models, the cooling and drying patterns extend 486 throughout the last millennium (Fig. S3). We revisit this discrepancy in the timing of the LIA 487 hydroclimate and temperature anomalies in Section 4.3.4. Mechanisms of the weakened Asian 488 monsoons in the models are explored further in Sections 4.5 and 4.6.

489 4.3.2. Wet/Isotopically-Depleted Conditions in the Maritime Continent

490 Further evidence that the period 1400-1700 CE was characterized by widespread tropical 491 hydroclimate changes comes from the Indonesian Straits in the Maritime Continent. In this 492 region, marine records generally indicate isotopically depleted surface ocean conditions around 493 1400-1700 CE, relative to present and to background conditions over the Common Era (Fig. 3G; 494 Fig. S10A). Four of the six marine records from the Maritime Continent demonstrate strong 495 coherency in their centennial-scale variability over the last millennium, characterized by a steady 496 trend to more depleted seawater d18O values from 1200 to 1700 CE (EOF1 in Fig. S10A). That 497 this pattern is not shared by all marine records in this region is perhaps expected given the 498 complexity of the surface ocean circulation in the Indonesian Archipelago. The inferred 499 freshening across a broad sector of the Maritime Continent generally opposes the drying pattern 500 seen in the South and East Asian composite (Fig. 3F) and also somewhat opposes the weak 501 drying pattern seen in the single record from northern Australia (Fig. S2C).

502 The reconstructed hydroclimate trends in the Maritime Continent region are sensitive to 503 the type of proxy archive (Fig. S2); namely, the LIA signal is largely absent in terrestrial records 504 from the region, including lake sediment records (Fig. S2F) and speleothem records (Fig. S2G) 505 from Indonesia and Borneo. These differences suggest that large-scale changes in moisture 506 balance in the region may have been integrated into a regionally-coherent seawater d18O signal, 507 while terrestrial records reflect more localized precipitation changes and/or changes in moisture 508 source and trajectories (e.g. Carolin et al., 2013). The considerable spatial heterogeneity in 509 rainfall throughout the Indonesian archipelago, both in the annual mean and seasonally, has been 510 well-documented (As-syakur et al., 2013) and attributed to the complex interplay between the

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511 ocean, islands, topography, and the multitude of regional circulation patterns that impact this 512 region.

513 4.3.3. Dipole Pattern in South American Hydroclimate

514 The final piece of evidence that the LIA was marked by widespread changes in tropical 515 hydroclimate comes from tropical South America, where the reconstructions show a distinct 516 antiphased dipole pattern between the southwestern and northeastern sites from 1400-1700 CE. 517 Records from the central Andes and the southern Amazon generally indicate wet (and/or 518 isotopically depleted) conditions during this period, followed by a pronounced drying trend to 519 the present (Fig. 3E), while an opposing pattern is seen in the composite from northern South 520 America and Nordeste (Fig. 3D). This dipole pattern is also apparent in the cluster 521 analysis (Fig. 2D-G), the epoch analysis of the LIA to MCA differences (Fig. 4), and the 522 EOF/PC analysis (EOFs 1-2 in Fig. S11A).

523 Because nine out of the eleven proxy records from the central Andes and southern 524 Amazon are water isotope records (Fig. 3E), which integrate climate information on large spatial 525 scales, the widespread signature of isotopic depletion in these records during the LIA is likely 526 reflective of increased upstream rainout under enhanced SASM activity, rather than changes in 527 local rainfall amount (Vuille et al., 2012). These interpretations are consistent with the review of 528 the 2,000-year of the South American monsoon by Vuille et al. (2012), which identified 529 three major excursions in the strength of the monsoon from the background climate of the 530 Common Era: the LIA (~1400-1800 CE) when the SASM was strong, and the MCA (~900-1100) 531 and the 20th century when the SASM was weak. The cause of the inferred changes in monsoon 532 intensity is not known, though (Vuille et al., 2012) hypothesize that cooler North Atlantic 533 temperatures during the LIA relative to the MCA led to a southward shift of the Atlantic ITCZ 534 and strengthening of the SASM. We discuss the timing of these monsoonal signals relative to 535 large-scale temperature reconstructions over the last millennium in the next section.

536 Antiphased hydroclimate changes between northern South America/Nordeste Brazil and 537 the central Andes/southern Amazon has been well-documented during the last glacial period and 538 over the in association with precessional forcing (Cruz et al., 2009). A dipole pattern in 539 d18O of precipitation over South America has been demonstrated by the coupled climate model 540 simulations of Liu et al. (2015) under precessional forcing, whereby a decrease in insolation in 541 austral summer (DJF) gives rise to isotopically enriched d18O of precipitation over the Andes (in 542 association with a reduced intensity of the SASM in response to reduced land heating) and 543 isotopically depleted precipitation over northeastern Brazil (in association with a southward shift 544 of the Atlantic ITCZ driven largely by cooling over southern Africa).

545 Another proposed dynamical mechanism of the dipole pattern that may be more relevant 546 to the LIA involves convective heating over the southwest Amazon during of enhanced 547 SASM activity that produces large-scale subsidence and decreased summer precipitation over 548 northeast Brazil (Lenters and Cook, 1997; Novello et al., 2012; Vuille et al., 2012). This 549 mechanism has been invoked to explain antiphased precipitation in tropical South America on 550 seasonal to orbital time scales (Cruz et al., 2009; Lenters and Cook, 1997) and is a plausible

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551 cause of the inferred antiphased precipitation changes observed between northeastern and 552 southwestern South America during the LIA (Novello et al., 2012; Vuille et al., 2012).

553 4.3.4. Correspondence Between Tropical Hydroclimate and Large-scale Temperature 554 Changes

555 As discussed above, the period from 1400-1700 CE stands out as a time of pronounced 556 tropical hydroclimate changes across the tropics. These changes include a strengthening of the 557 South American summer monsoon, weakened Asian summer monsoons, and fresher conditions 558 in parts of the Maritime Continent. Previous studies have highlighted the general correspondence 559 between these hydroclimate changes and changes in the large-scale (global or NH mean) 560 temperature or changes in the interhemispheric temperature gradient (e.g. Sinha et al., 2011; 561 Vuille et al., 2012; Yan et al., 2011). However, we find notable dissimilarities between the 562 timing of these hydroclimate changes and the evolution of global mean temperature and the 563 interhemispheric temperature gradient over the last millennium based on recent temperature 564 reconstructions (PAGES2k Consortium, 2019; Neukom et al., 2014). In particular, the 565 hydroclimate composites from Asia (Fig. 3F), southwestern South America (Fig. 3E), and the 566 Maritime Continent (Fig. 3G) all indicate a reversal in the LIA hydroclimate trends around 1600 567 CE, while the reversal of global cooling occurs substantially later, around 1800 CE in the 568 ensemble-median of the reconstructions (Fig. 3H). The reconstructed interhemispheric 569 temperature gradient displays a more pronounced increase from 1600 CE to present, however the 570 interhemispheric gradient and hydroclimate trends diverge during the earlier half of the last 571 millennium, e.g. during the temperature gradient minimum around 1200-1300 CE (Fig. 3H). The 572 timing of the reversal of the LIA hydroclimate trends around 1600 CE is remarkably consistent 573 across the regional composites of South and East Asia, South America, and the Maritime 574 Continent, and is also reflected in Clusters 1 and 3 from the cluster analysis (Fig. 2D,F). While 575 the age model uncertainty in the reconstructions must be taken into account (particularly in the 576 14C-dated marine and lake sediment records with few age control points; Table S1), the Asian 577 and South American composites are dominated by well-dated speleothem records, and the 578 synchronicity of this feature across a wide range of proxy archives from different regions of the 579 tropics supports the inference of a widespread reversal in LIA hydroclimate trends around 1600 580 CE.

581 It is perhaps unsurprising that the LIA hydroclimate changes were not simply related to 582 global cooling trends or a weakened interhemispheric temperature gradient, given that little 583 spatial coherence has been found in regional cooling patterns over the LIA, with the half-century 584 of greatest cooling ranging from the 15th to 19th century in different regions of the world 585 (Neukom et al., 2019). The hydroclimate changes in South America, Asia, and the Maritime 586 Continent may be better understood as a response to regional temperature variations, though the 587 coordinated nature of the hydroclimate anomalies across these disparate regions of the tropics 588 remains to be explained. Regardless, it is clear that more work needs to be done to understand 589 the mechanisms of the widespread tropical hydroclimate changes around 1400-1700 CE. A 590 targeted comparison of the regional temperature changes in the tropics could also help 591 disentangle the large data-model discrepancies in tropical hydroclimate, as the model simulations

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592 have been shown to display more regional coherency than the reconstructions in temperature 593 variations over the last millennium (PAGES 2k-PMIP3 group, 2015).

594 4.4. Re-evaluation of Tropical Pacific Climate Changes Over the Last Millennium

595 Based on the above findings in Section 4.3, we now re-examine a number of longstanding 596 hypotheses about the nature of tropical Pacific hydroclimate changes during the LIA. These 597 hypotheses include: 1) a southward shift of the mean annual position of the Pacific ITCZ (Nelson 598 and Sachs, 2016; Richey and Sachs, 2016; Sachs et al., 2009), 2) an equatorward contraction of 599 the western Pacific and Asian-Australian rainband (Denniston et al., 2016; Griffiths et al., 2016; 600 Yan et al., 2015), and 3) a strengthened Pacific Walker Circulation (Griffiths et al., 2016; Yan et 601 al., 2011; Yan et al., 2015).

602 4.4.1. A Largescale Southward Shift of the Pacific ITCZ During the LIA?

603 A number of previous studies have inferred a coordinated, large-scale southward shift of 604 the Pacific ITCZ during the relatively cool epoch of the LIA, arguing that this interpretation is 605 consistent with dynamical theory that links the zonal mean ITCZ latitude to the inter- 606 hemispheric energy budget and temperature gradient (Nelson and Sachs, 2016; Richey and 607 Sachs, 2016; Sachs et al., 2009). This explanation has also been invoked in the inference of a 608 southward shift of the Atlantic ITCZ during the LIA (Haug et al., 2001). Some of the earliest 609 evidence to support a southward shift of the Pacific ITCZ came from a d2H record from lake 610 sediments in Washington Island in the central equatorial Pacific (Sachs et al., 2009), which has a 611 perennially wet climate due to its location under the southern margin of the Pacific ITCZ (Fig. 1; 612 4.7° N, 160.8°W). An abrupt shift to drier conditions around 1400 CE (as inferred from 613 isotopically-enriched total lipid extracts and the presence of microbial mat communities similar 614 to those in hypersaline ponds from Kiritimati Atoll at 1°52’N, 157°20’W) led to the 615 interpretation of an abrupt transition to drier conditions as the rain band shifted south during the 616 onset of the LIA (Sachs et al., 2009).

617 However, we do not find broad support for a large-scale southward shift of the Pacific 618 ITCZ during the LIA. First, the theoretical framework for this theory is lacking. Comprehensive 619 climate models indicate that large-scale shifts in the tropical rain belt occur only under strong 620 hemispheric asymmetry in radiative forcing, and even in those cases, there are large zonal 621 variations such that even the tropical Pacific rain band doesn’t shift uniformly (e.g. Atwood et 622 al., 2020). Second, no clear evidence of a largescale southward shift in the Pacific rain band is 623 observed when all available records from the central and eastern Pacific (where the Pacific ITCZ 624 is well defined) are considered. While the abrupt transition to isotopically-enriched values, 625 indicating dry conditions, is clear in the Washington Island record around 1500 CE in our 626 compilation (Fig. 3A; Fig. S2B), given its location under the southern margin of the Pacific 627 ITCZ, dry conditions at that location would be more likely to be associated with a more northerly 628 location of the Pacific ITCZ (Fig. 1), as associated with an enhanced Pacific Walker Circulation 629 during La Niña events. Notably, there is also no synchronous shift in the eastern equatorial 630 Pacific compilation around this time (Fig. 3A), which lie just south of the southern extent of the 631 rain band (Fig. 1), and thus would be expected to moisten under a southward shift of the Pacific 632 ITCZ. However, any robust conclusions regarding changes in the Pacific rainband during the 633 LIA will require more data from the central and eastern Pacific, where the Pacific ITCZ is well-

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634 defined. In addition, interpretations of the Galápagos records during the LIA are still being 635 refined. While all Galápagos records indicate the pronounced drying event around 800-1000 CE, 636 the records show less consistency in the latter part of the last millennium (Fig. 3A) with the 637 highland records (Atwood and Sachs, 2014; Conroy et al., 2008) agreeing with one another but 638 showing distinct behavior from the single lowland record (Nelson and Sachs, 2016). Ongoing 639 efforts to develop additional reconstructions and calibrate proxy data against modern 640 instrumental data should yield new insights in this data-sparse region of the tropics (e.g. Martin 641 et al., 2018).

642 As it stands, without more evidence for a coordinated large-scale southward shift of the 643 annual mean Pacific rain band during the LIA, support for this interpretation is lacking. As 644 additional paleoclimate records are developed from the central and eastern tropical Pacific, 645 alternate interpretations of the inferred rainfall changes captured by such decadally-resolved lake 646 sediment records should be considered, such as changes in rain band width, intensity, and 647 seasonal migration (Bartos et al., 2018; Byrne et al., 2018; Singarayer et al., 2017; Wodzicki and 648 Rapp, 2016), changes in the strength of the Pacific Walker Circulation, and changes in the 649 variability and spatial characteristics of the El Niño/Southern Oscillation (ENSO).

650 4.4.2. Contraction of the Monsoonal Asian-Australian Rainbelt During the LIA?

651 Another proposed reorganization of tropical climate during the LIA comes from the 652 western edge of the Pacific basin, where previous compilations of records from the Maritime 653 Continent, Australia, and East Asia have been interpreted as an equatorward contraction of 654 rainfall at various periods within the LIA (ranging from 1400-1640 CE (Denniston et al., 2016), 655 1400-1850 CE (Yan et al., 2015), and 1500-1900 CE (Griffiths et al., 2016)). Characterized by 656 drying in the Asian and Australian summer monsoons and wetting near the equator, our findings 657 generally align with these prior studies, though there are some aspects of this interpretation that 658 remain tenuous due to limited data availability. In particular, support for an equatorward 659 contraction and/or weakening of the Australian monsoon during the LIA presented in Denniston 660 et al. (2016), Griffiths et al. (2016), and the present study, all rely on a single record from the 661 Australian monsoon region (Denniston et al., 2016). Our analyses indicate that this record is 662 consistent with a weakened Australian monsoon from 1400-1600 CE followed by an 663 intensification to modern times, though notably it does not provide support for anomalous LIA 664 conditions relative to background conditions of the Common Era (Fig. S2C). Although eight 665 records from the southern edge of the Australian Summer Monsoon were included in the analysis 666 of Yan et al. (2015), only three of these records spanned the last millennium, and of these, only 667 one record (Denniston et al., 2016) was publicly available and thus used in the present analysis. 668 In addition, in the analyses presented in Yan et al. (2015), the reference period varied between 669 the MCA and the last 150 years for different records in their compilation, depending on record 670 length. We see large differences between the MCA and the last 200 years in many records in our 671 analysis, and thus we strongly advocate for a standardized treatment across all proxy records so 672 that the trends over the last few hundred years are separable from the changes between the MCA 673 and LIA.

674 While we find general coherency in the pattern of LIA drying in Asia and northern 675 Australia and freshening of the surface ocean in the Maritime Continent, the dominant pattern of 676 centennial variability across records in the Asian-Australian rainbelt is notably more complex

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677 than a simple equatorward contraction of rainfall (Fig. S12A). To robustly test the hypothesis of 678 a coordinated contraction of the monsoonal Asian-Australian rainbelt over the last millennium, 679 additional records from the Australian monsoon region are needed, as well as an improved 680 understanding of the regional signatures in the terrestrial and marine records from the Maritime 681 Continent (described in Section 4.3.2). Further, this hypothesis would benefit from a plausible 682 set of physical mechanisms that could explain a coordinated response of these two highly 683 seasonal and out-of-phase monsoonal rainfall regimes.

684 4.4.3. Strengthened Pacific Walker Circulation During the LIA?

685 Still other studies have compiled hydroclimate records from the tropical Pacific and 686 surrounding regions to infer a strengthened Pacific Walker Circulation during the LIA relative to 687 the MCA (Griffiths et al., 2016; Yan et al., 2011; Yan et al., 2015). The handful of marine 688 records from the Maritime Continent indicating fresher conditions during the LIA generally 689 support this interpretation (Fig. 3G), as does the Washington Island record in the central tropical 690 Pacific, which indicates an abrupt shift to drier conditions around 1500 CE (Fig. 3A; Fig. S2B). 691 While there is no evidence for drier conditions in the eastern equatorial Pacific (EEP) in 692 association with a strengthened Pacific Walker Circulation during the LIA (Fig. 3A), this 693 apparent discrepancy may be explained by the fact that the Galápagos Islands are situated in the 694 already dry zone of the Pacific cold tongue and thus may experience little change in average 695 rainfall under an enhanced Walker Circulation. The available data thus supports LIA freshening 696 in parts of the Maritime Continent, with a possible linkage to central Pacific drying (based on a 697 single record), but with no clear connection to the eastern Pacific. We therefore conclude that 698 until further paleoclimate records are developed from the tropical Pacific, inferences of a 699 strengthened Pacific Walker Circulation during the LIA remain speculative. Targeting the 700 position of the Pacific rainbands could be especially helpful in this regard, as the rainbands 701 would be expected to expand poleward (equatorward) under an enhanced (weakened) Pacific 702 Walker Circulation as occurs in association with the El Niño/Southern Oscillation (Wallace et 703 al., 1998). In the next section we demonstrate that one model, CESM, simulates a strengthened 704 Pacific Walker Circulation over the last millennium, but that our analyses indicate that this 705 response may be unique to this particular model.

706 4.5. Possible Sources of Data-Model Disagreement

707 In order to shed light on the mechanisms of the robust long-term hydroclimate changes 708 observed in the proxy records over the last millennium, including distinguishing between forced 709 and unforced variability, we compared the reconstructed hydroclimate patterns to a large suite of 710 forced transient climate model simulations of the last millennium. The centennial-scale rainfall 711 changes simulated by the models generally display very little agreement with proxy data in the 712 regions where proxy data exist (Fig. 4, Fig. S1). Even over South and East Asia where both 713 proxies and models seem to indicate drier conditions during the second half of the last 714 millennium (Fig. 4), the temporal evolution of the large-scale drying pattern in the models does 715 not match the reconstructions. In the reconstructions, peak dry conditions (isotopic enrichment) 716 occur around 1600 CE (Fig. S3A). While in the models, a continuous large-scale drying trend

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717 occurs over this region throughout the last millennium (Fig. S3B,C) that tracks the simulated 718 millennial-scale cooling trend in the tropics and extratropics (Fig. S9A,B).

719 Some of the most pronounced data/model discrepancies occur over tropical South 720 America where the simulated and reconstructed hydroclimate changes are both robust, but show 721 little agreement with one another (Fig. 4). While the reconstructions have been interpreted as 722 representing increased upstream rainout under enhanced SASM activity in tandem with 723 subsidence and decreased summer precipitation over northeastern Brazil (Lenters and Cook, 724 1997; Novello et al., 2012; Vuille et al., 2012), the models simulate increased rainfall in the far 725 northeastern corner of South America that is driven by precessional forcing over the last 726 millennium (as discussed in Section 4.6). It is possible that the observed LIA isotopic depletion 727 pattern in the central Andes and southern Amazon was due to increased upstream rainout via an 728 enhanced SASM, but that this rainout occurred over northeastern South America instead of the 729 Amazon. However, such a mechanism does not explain the observed antiphased LIA 730 drying/isotopic enrichment signal in the northeast. The data-model disagreement in this region 731 should be more rigorously assessed in future studies by investigating last millennium simulations 732 with water isotope tracers.

733 The tropical Pacific is another region of pronounced data-model disagreement, especially 734 in the CMIP5/PMIP3 models (Fig. 4). The tropical Pacific proxy records align moderately better 735 with the CESM LME simulations, which are characterized by a strengthened Pacific Walker 736 Circulation with drying in the Pacific cold tongue and along the equatorward edges of the Pacific 737 rainbands, and wetting in parts of the equatorial Maritime Continent (Fig. 4C,D). These features 738 are driven by orbital (and to a lesser extent, volcanic) forcing in CESM. However, this response 739 appears to be unique to that model as none of the six CMIP5/PMIP3 models indicate a 740 strengthened Walker Circulation under the combined last millennium forcings (Fig. S6). These 741 results indicate that while the forced hydroclimate signals in the tropical Pacific may be large 742 enough to rise above the natural climate variability in CESM, the large structural uncertainty 743 amongst models precludes any confidence in this forced response.

744 From the pronounced data-model disagreement, we conclude one or more of the 745 following: (1) the models are not able to accurately simulate the forced rainfall response (due to 746 inaccuracies in either the climate forcings, the local or remote feedbacks, or the tropical 747 hydroclimate responses), (2) the models underestimate the forced response relative to the internal 748 variability, (3) the robust tropical hydroclimate changes observed in the reconstructions were 749 unforced, and/or (4) the proxy records, which are predominately derived from water isotope- 750 based archives, reflect something other than changes in regional mean annual precipitation 751 amount. We discuss these possibilities in more detail below.

752 While a dominant role for internal climate variability can’t be ruled out, the widespread 753 nature of the tropical hydroclimate anomalies from 1400-1700 CE and the broad correspondence 754 with cooling around the globe is suggestive of a forced response If inaccuracies in the climate 755 models are responsible for the data-model discrepancies, the likely culprits are the errors in the 756 radiative forcings and feedbacks, and/or in the hydroclimate responses to the forcings and 757 feedbacks. For starters, there are substantial uncertainties in the radiative forcings over the last 758 millennium. The radiative forcing during the LIA differs by more than a factor of two across the 759 CMIP5/PMIP3 models due to differences in volcanic forcing (associated with both the aerosol

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760 dataset used and the implementation method of the volcanic aerosols; see Methods) (Atwood et 761 al., 2016). These differences contribute to a factor of two spread in the amplitude of global 762 cooling during the LIA across these models. Differences in volcanic forcing also likely lead to 763 differing tropical rainfall responses in the models, since tropical rainfall patterns are known to be 764 sensitive to the timing, magnitude, and spatial footprint of volcanic eruptions (e.g. Colose et al., 765 2016; PAGES Hydro2k Consortium, 2017). Furthermore, because the volcanic aerosol loadings 766 have changed in recently updated last millennium reconstructions (Sigl et al., 2014; Sigl et al., 767 2015), the next iteration of model simulations may be expected to differ in their representation of 768 volcanically-forced tropical precipitation changes over the last millennium. In addition to the 769 uncertainties in radiative forcing, there are also uncertainties in the strength and pattern of 770 climate feedbacks among models. For example, the total effective feedback parameter during the 771 LIA differs by a factor of two across the CMIP5/PMIP3 models (ranging from -1.0 to -2.0 772 W/m2/K), largely due to differences in the shortwave cloud feedback, the lapse rate feedback, 773 and the surface albedo feedback (Atwood et al., 2016).

774 Uncertainties in the radiative forcings and feedbacks aside, however, it is perhaps 775 unsurprising that ocean-atmosphere general circulation models would have difficulty 776 reproducing regional tropical hydroclimate responses to last millennium forcings, due to the 777 highly parameterized nature of convection and cloud physics, and given the limited ability of 778 CMIP5 models to reproduce observed regional hydroclimate dynamics associated with major 779 hydroclimate phenomena such as the Asian and North American monsoons (Flato, 2013; PAGES 780 Hydro2k Consortium, 2017). Furthermore, paleoclimate simulations do not typically include 781 processes such as dynamic vegetation (which require ocean–atmosphere–vegetation general 782 circulation models), although it is known that land surface processes are integral to the cycling of 783 water between the land and atmosphere. Indeed, vegetation changes have been shown to reduce 784 model-data discrepancies in simulations of the mid-Holocene (Braconnot et al., 2012). The 785 importance of land cover changes on last millennium climate is apparent in the single-forcing 786 simulations of CESM (Fig. 5); these land use changes are also a major contributor to the global 787 radiative forcing near the end of the last millennium (Atwood et al., 2016). Future data-model 788 comparisons should consider the role of dynamic vegetation and should also address the ability 789 of the models to simulate the first-order features of the key hydroclimate phenomena of interest, 790 namely in monsoonal Australasia and South America, and the tropical Pacific.

791 Finally, it is likely that the proxy records used in this analysis, many of which are derived 792 from water isotope-based archives, do not solely reflect changes in regional precipitation 793 amount. The stable isotopic composition of meteoric water integrates information about 794 numerous aspects of the water cycle over and time, including precipitation amount, 795 seasonality of precipitation, the sources of moisture, advective transport in the ocean, 796 evaporation, and water recycling over land. Future work should therefore employ climate models 797 that simulate water isotope tracers to provide more quantitative comparisons between proxy 798 reconstructions and model simulations of the last millennium.

799 4.6. Orbital Forcing Dominates Tropical Hydroclimate Changes in CESM

800 In exploring potential sources of the data-model discrepancies, it is useful to investigate 801 the source of the forced precipitation response in the models. The single forcing simulations with 802 CESM indicate that orbital forcing is largely responsible for the robust set of tropical

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803 precipitation changes in the models. Orbital forcing is associated with a 20-day shift in 804 perihelion (from 15 December to 4 January) over the last millennium, leading to an increase in 805 insolation in boreal spring and early summer relative to late boreal summer and fall. Insolation in 806 the NH high latitudes declined from July-September by ~3 Wm-2, and declined in the SH high 807 latitudes from October-December, i.e. late austral spring to summer, by a similar magnitude over 808 the last millennium (Schmidt et al. 2011). In CESM, orbital forcing causes drying in the NH 809 monsoon systems of South Asia and West Africa and an equatorward contraction/intensification 810 of rainfall over the equatorial Atlantic Ocean and northern South America (Fig. 5). These 811 features are present, though more muted, in the all-forcing CESM simulations and in the 812 CMIP5/PMIP3 simulations (Fig. 4, Fig. 5), indicating a dominant role for orbital forcing in 813 driving the forced long-term tropical hydroclimate response over the last millennium in the 814 models. Within the Pacific basin, orbital forcing (and to a lesser extent, volcanic forcing) 815 enhances the Pacific Walker Circulation and intensifies the Pacific rain bands in CESM (Fig. 5). 816 However, this tropical Pacific response to the combined last millennium forcings appears to be 817 unique to CESM (Fig. 4; Fig. S6). Further details of the tropical hydroclimate responses to 818 orbital forcing are provided in the Supplemental Information.

819 These results indicate that while global cooling over the last millennium is largely driven 820 by volcanic forcing in climate models (e.g. Atwood et al., 2016), the robust multi-centennial 821 changes in tropical precipitation in the models are largely driven by orbital (precessional) 822 forcing. The fact that the simulated precipitation changes don’t agree with the reconstructions, 823 however, indicates that a precession-driven control on tropical precipitation over the last 824 millennium, while robust across models, is inconsistent with the proxy data.

825 4.7. Robust Zonally-Averaged Tropical Rainfall Response in Models

826 Despite large differences in regional precipitation across models, some coherency is 827 observed in the zonally averaged context: all models demonstrate a small zonal mean southward 828 shift in tropical precipitation (20ºS:20ºN) during the LIA relative to the MCA (Fig. 7). The 829 changes in precipitation asymmetry are small (DPNH-SH £ 0.3 mm/day), representing a -15% to 830 +10% change in precipitation asymmetry across models. To put the magnitude of these changes 831 into context, the multi-model mean change in the tropical precipitation centroid (defined as the 832 median of zonal average precipitation from 20ºS to 20ºN) is -0.03º latitude in the last millennium 833 simulations (data not shown), which is approximately 6% as large as the -0.53º multi-model 834 mean shift in the PMIP2 and PMIP3 Last Glacial Maximum simulations and 1% as large as the 835 -2.48º shift in 1Sv North Atlantic meltwater simulations (Atwood et al., 2020).

836 This zonal mean precipitation shift occurs in association with anomalous northward 837 cross-equatorial atmospheric energy transport that is primarily driven by volcanic forcing (i.e. 838 greater volcanic aerosol loading in the northern hemisphere), land use changes (i.e. cropland 839 expansion in Eurasia), and the surface albedo feedback forcing (e.g. through the growth of Arctic 840 sea ice in response to orbital forcing; Fig. S13). The net changes in cross-equatorial atmospheric 841 heat transport are modest due to the compensating shortwave energy fluxes that oppose the 842 changes in outgoing longwave radiation (due to greater cooling of the NH). While these zonal 2 843 mean changes are small (DPNH-SH £ 0.3 mm/day; DAHTF=0 < 0.2 W/m ), they demonstrate a 844 perceptible influence of the hemispherically-asymmetric forcings and feedbacks that drive

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845 enhanced NH cooling (namely volcanic forcing and Eurasian cropland expansion) on the tropical 846 precipitation response in the models.

847 The fact that the zonal mean precipitation changes are fairly consistent across models 848 indicates that the net hemispheric asymmetry in the forcings and feedbacks are generally 849 consistent across models and also that the zonal mean precipitation asymmetry is less sensitive to 850 structural uncertainties in the models than regional rainfall (which depends on convective 851 parameterizations in the models). However, the zonal mean rainfall changes are clearly irrelevant 852 to our analysis of regional rainfall patterns, since they are too small to be of consequence for 853 interpreting the changes seen in the proxy records and since zonal mean changes have little 854 bearing on regional rainfall patterns (e.g. Atwood et al., 2020).

855

856 5. Conclusions

857 We review the evidence for large-scale tropical hydroclimate changes over the Common 858 Era based on a compilation of tropical hydroclimate records and assess the consistency between 859 the reconstructed hydroclimate changes and those simulated by forced transient model 860 simulations of the last millennium. A hierarchal cluster analysis indicates that 41% of the records 861 distributed throughout the tropics demonstrate a millennial-scale isotopic depletion/wetting trend 862 beginning around 1000 CE and reversing around 1800 CE that generally tracks global 863 temperature trends over the last millennium, which we speculate is associated with the influence 864 of decreased tropical sea surface temperatures on the isotopic composition of water vapor.

865 Our synthesis of the proxy records reveal a number of regionally coherent hydroclimate 866 patterns. During the onset of the MCA around 800-1000 CE, records from the eastern equatorial 867 Pacific and Mesoamerica indicate a pronounced drying event, relative to background conditions 868 of the Common Era. The second half of the last millennium is characterized by a set of 869 pronounced and synchronous hydroclimate changes across the tropics from 1400-1700 CE, 870 including drier (isotopically enriched) conditions in South and East Asia, fresher (isotopically 871 depleted) surface ocean conditions in the Maritime Continent, and a dipole pattern in tropical 872 South America, with wet (isotopically depleted) conditions in the southwest and dry (isotopically 873 enriched) conditions in the northeast. While previous studies have linked these hydroclimate 874 changes to global or hemispheric-scale cooling patterns, we find notable dissimilarities in the 875 regional hydroclimate changes and global mean temperature reconstructions, particularly in the 876 timing of reversal of the LIA hydroclimate trends around 1600 CE, indicating that more work 877 needs to be done to understand the mechanisms of the widespread tropical hydroclimate changes 878 during the LIA.

879 Considering previous interpretations of largescale reorganization of the tropical Pacific 880 climate system during the LIA, we do not find support for a largescale southward shift of the 881 Pacific ITCZ during the LIA. Consistent with previous studies, we find that the available 882 evidence generally supports an equatorward contraction of the monsoonal Asian-Australian 883 rainbelt during the LIA, though we caution that more work needs to be done to better 884 characterize the hydroclimate changes in the Australian monsoon region and in the Maritime 885 Continent. Regarding evidence for a strengthened the Pacific Walker Circulation during the LIA,

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886 we find that the available data supports LIA freshening in parts of the Maritime Continent, with a 887 possible linkage to central Pacific drying (based on a single record), but with no clear connection 888 to the eastern Pacific. We therefore conclude that until further paleoclimate records are 889 developed from the tropical Pacific, inferences of a strengthened Pacific Walker Circulation 890 during the LIA remain speculative.

891 Lastly, we find that transient climate model simulations of the last millennium exhibit 892 weak forced long-term tropical rainfall changes over the last millennium that generally poorly 893 agree with the reconstructions, suggesting one or more of the following: (1) the models are not 894 able to accurately simulate the forced rainfall response (either due to inaccurate climate forcings 895 or inaccurate climate responses to the forcings), (2) the models underestimate the forced 896 response relative to internal climate variability, (3) the robust regional signals in the 897 reconstructions were unforced, and/or (4) the proxy records, which are predominately derived 898 from water isotope-based archives, reflect something other than changes in regional mean annual 899 precipitation amount (such as the seasonality of precipitation or water recycling over land). 900 Future work to investigate the mechanisms of the robust regional hydroclimate patterns observed 901 in the proxy records should therefore employ climate models that simulate water isotope tracers 902 to provide more quantitative comparisons between proxy reconstructions and model simulations 903 of the last millennium.

904

905 Acknowledgments 906 The authors gratefully acknowledge all of the researchers who made their paleoclimate 907 data available for public access and who participated in the CMIP5/PMIP3 modeling efforts and 908 made the model output publicly available. This material is based upon work supported by the 909 U.S. National Science Foundation Graduate Research Fellowship Program (DGE-1256082) and 910 by the NOAA Climate and Global Change Postdoctoral Program (under fellowships to ARA). 911 Support was provided to EW and DMWF by NSF Awards AGS-0846641, AGS-0936059, AGS- 912 1359464, and AGS-1665247. Support was provided to D.S.B by the Tamaki Foundation, and to 913 J.P.S. by the U.S. National Science Foundation under Grant Nos. 0823503 and 1502417. We 914 thank K. Shell for providing the CAM3 radiative kernels. The proxy data used in this study were 915 uploaded to Figshare (https://figshare.com/articles/Hydroclimate_proxy_records_zip/12085092) 916 and published with doi: 10.6084/m9.figshare.12085092. The records were originally downloaded 917 from the NOAA Paleoclimatology and PANGAEA databases (see dataset URLs in Table 1), or 918 obtained directly from the authors (see references in Table 1). 919

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920 Tables and Figures 921 Table 1. List of the paleoclimate data used in this study Lat Lon Elevation Cluster #, Avg resolution Region Site Name (deg N) (deg E) (m) Proxy archive Measurement Index # 1 (years) Reference Dataset URL El Junco Lake, San Cristobal Island, Galápagos -0.9 -89.5 670 lake sediment dinosterol !2H C2, 7 67 Atwood et al., 2014 https://doi.org/10.1594/PANGAEA.834103 Eastern Pacific El Junco Lake, San Cristobal Island, Galápagos -0.9 -89.5 660 lake sediment sediment grain size (% sand) C2, 13 11 Conroy et al., 2008 https://www.ncdc.noaa.gov/paleo/study/6109 Poza de las Diablas, Isabela Island, Galápagos -1.0 -91.0 lake sediment !2H lipid biomarkers C2, 23 45 Nelson and Sachs, 2016 https://www.pnas.org/content/suppl/2016/03/09/1516271113.DCSupplemental Central Pacific Washington Island, Kiribati 4.7 -160.4 lake sediment !2H of total lipid extract 42 Sachs et al., 2009 https://www.nature.com/articles/ngeo554#Sec11 !2 South Pacific Harai Lake #1, Rendova Island, Solomon Islands (SHAR1) -8.6 157.4 lake sediment dinosterol H 185 Maloney, 2018 Apu Bay, Tahaa, French Polynesia -16.7 -151.5 0 marine sediment Ti/Ca C2, 33 1 Toomey et al., 2016 https://agupubs.onlinelibrary.wiley.com/doi/full/10.1002/2015PA002870 MD98-2181 6.3 125.8 -2114 marine sediment carbonate !18O and Mg/Ca (forams) C2, 40 34 Stott et al., 2004 https://www.ncdc.noaa.gov/paleo/study/6236 Bukit Cave, Malaysian Borneo 4.3 115.0 250 speleothem carbonate !18O 6 Chen et al., 2016 https://www.ncdc.noaa.gov/paleo/study/19885 Snail Shell Cave, Malaysian Borneo (SSC01) 4.2 114.9 150 speleothem carbonate !18O C3, 12 53 Carolin et al., 2016 https://www.ncdc.noaa.gov/paleo/study/19806 MD98-2176 -0.2 133.4 -2382 marine sediment carbonate !18O and Mg/Ca (forams) C1, 41 42 Stott et al., 2004 https://www.ncdc.noaa.gov/paleo/study/6236 Makassar Strait, Indonesia (31MC, 34GGC) -3.9 119.5 -459 to -503 marine sediment !2H C30 fatty acid C1, 32 33 Tierney et al., 2010 https://www.ncdc.noaa.gov/paleo/study/10438 Makassar Strait, Indonesia (composite: 31MC, MD60, 34GGC, 32GGC) -3.9 119.5 -454 to -503 marine sediment carbonate !18O and Mg/Ca (forams) C2, 26 10 Oppo et al., 2009 https://www.ncdc.noaa.gov/paleo/study/8699 Western Pacific / 18 Maritime Continent Makassar Strait, Indonesia (MD9821-60) -5.2 117.5 -1185 marine sediment carbonate ! O and Mg/Ca (forams) 9 Newton et al., 2006 https://www.ncdc.noaa.gov/paleo/study/5534 Ranu Lamongan, East Java, Indonesia -8.0 113.4 240 lake sediment mol% Mg in calcite 3 Crausbay et al., 2006 https://www.ncdc.noaa.gov/paleo/study/22409 Lake Lading, SW Indonesia -8.0 113.3 324 lake sediment !2H n-alkanoic acid 9 Koneckey et al., 2013 https://www.ncdc.noaa.gov/paleo/study/14129 Lake Logung, eastern Java, Indonesia -8.0 113.3 215 lake sediment !13C of organic material C1, 28 17 Rodysill et al., 2012 https://www.ncdc.noaa.gov/paleo/study/13177 Liang Luar Cave, Flores, Indonesia (splice: LR06-B1, LR06-B3) -8.5 120.4 550 speleothem carbonate !18O C1, 16 12 Griffiths et al., 2009 https://www.ncdc.noaa.gov/paleo/study/8632 MD98-2170 -10.6 125.4 -832 marine sediment carbonate !18O and Mg/Ca (forams) 163 Stott et al., 2004 https://www.ncdc.noaa.gov/paleo/study/6236 Australia Cave KNI-51, Western Australia -15.2 128.4 100 speleothem carbonate !18O 3 Denniston et al., 2016 https://www.ncdc.noaa.gov/paleo/study/20530 Huangye Cave, northern China 33.6 105.1 1650 speleothem carbonate !18O C3, 30 5 Tan et al., 2010 https://www.ncdc.noaa.gov/paleo/study/11177 Jiuxian Cave, central China (C996-1) 33.6 109.1 1495 speleothem carbonate !18O C1, 11 4 Cai et al., 2010a https://www.ncdc.noaa.gov/paleo/study/9742 Jiuxian Cave, central China (C996-2) 33.6 109.1 1495 speleothem carbonate !18O C1, 11 27 Cai et al., 2010b https://www.ncdc.noaa.gov/paleo/study/9742 Wanxiang Cave, central China 33.3 105.0 1200 speleothem carbonate !18O C3, 37 3 Zhang et al., 2008 https://www.ncdc.noaa.gov/paleo/study/8629 Sanbao Cave, central China (sb43) 31.7 110.4 1900 speleothem carbonate !18O 36 Dong et al., 2010a https://www.ncdc.noaa.gov/paleo/study/10445 Sanbao Cave, central China (sb26) 31.7 110.4 1900 speleothem carbonate !18O 26 Dong et al., 2010b https://www.ncdc.noaa.gov/paleo/study/10445 Heshing Cave, SW China 30.5 110.4 294 speleothem carbonate !18O C2, 19 3 Hu et al., 2008 https://www.ncdc.noaa.gov/paleo/study/6095 Dongge Cave, southern China (D-4) 25.3 108.1 680 speleothem carbonate !18O C2, 14 8 Dykoski et al., 2005 https://www.ncdc.noaa.gov/paleo/study/5441 South/East Asia 18 Dongge Cave, southern China (DA) 25.3 108.1 680 speleothem carbonate ! O C2, 35 4 Wang et al., 2005 https://www.ncdc.noaa.gov/paleo/study/5439 Dongge Cave, southern China (DX1) 25.3 108.1 680 speleothem carbonate !18O 1 Zhao et al., 2015 ftp://ftp.ncdc.noaa.gov/pub/data/paleo/speleothem/Asia/China/Dongge 2015.xls Xiang Yun Lake, China 24.4 102.8 1721 lake sediment carbonate !18O (authigenic) C3, 18 3 Hillman et al., 2014 http://ncdc.noaa.gov/paleo/study/17515 South China Sea 20.1 117.4 -1727 marine sediment carbonate !18O (diatoms) C1, 36 17 Wang et al., 1999 https://www.ncdc.noaa.gov/paleo/study/2598 Dandak Cave, India 19.0 82.0 400 speleothem carbonate !18O 1 Berkelhammer et al., 2010 https://www.ncdc.noaa.gov/paleo/study/8639 Jhumar Cave, India 18.9 81.9 600 speleothem carbonate !18O 2 Sinha et al., 2011 https://www.ncdc.noaa.gov/paleo/study/11937 Cattle Pond, Dongdao Island, South China Sea (DY2) 16.7 112.7 lake sediment sediment grain size 14 Yan et al., 2011a https://www.ncdc.noaa.gov/paleo/study/11197 Cattle Pond, Dongdao Island, South China Sea (DY4) 16.7 112.7 lake sediment sediment grain size 21 Yan et al., 2011b https://www.ncdc.noaa.gov/paleo/study/11197 Cattle Pond, Dongdao Island, South China Sea (DY6) 16.7 112.7 lake sediment sediment grain size 6 Yan et al., 2011c https://www.ncdc.noaa.gov/paleo/study/11197 Lake Bosumtwi, Ghana 6.5 -1.4 lake sediment carbonate !18O (authigenic) C2, 29 5 Shanahan et al., 2009 https://www.ncdc.noaa.gov/paleo/study/8657 Lake Edward, Uganda/Democratic Republic of the Congo -0.4 29.8 917 lake sediment % Mg in calcite 4 Russell et al., 2007 https://figshare.com/articles/Proxy_ data/4787575/1 Africa Dante Cave, Namibia -19.4 17.9 speleothem carbonate !18O 11 Sletten et al., 2013 https://www.ncdc.noaa.gov/paleo/study/14268 Lake Sibaya, -27.3 32.6 20 lake sediment % planktonic diatoms C1, 39 18 Stager et al., 2013 https://www.ncdc.noaa.gov/paleo/study/14629 Tzabnah Cave, Yucatan, Mexico 20.7 -89.5 20 speleothem carbonate !18O C2, 4 3 Medina-Elizalde et al., 2010 https://www.ncdc.noaa.gov/paleo/study/9791 Lake Punta, Yucatan, Mexico 20.6 -87.5 14 lake sediment carbonate !18O (Cytheridella ilosvayi) C2, 1 7 Curtis et al., 1996a https://www.ncdc.noaa.gov/paleo/study/5484 Lake Punta, Yucatan, Mexico 20.6 -87.5 14 lake sediment carbonate !18O (Pyrgophorus coronatus) C2, 1 7 Curtis et al., 1996b https://www.ncdc.noaa.gov/paleo/study/5484 Laguna de Juanacatlán, Mexico 20.6 -104.7 2000 lake sediment Ti concentration 1 Metcalfe et al., 2010 https://www.ncdc.noaa.gov/paleo/study/19768 Lake Chichancanab, Yucatan, Mexico 19.9 -88.8 lake sediment lake sediment density C2, 3 4 Hodell et al., 2008 https://www.ncdc.noaa.gov/paleo/study/6108 Aguada X'caamal Lake, Yucatan Mexico 20.6 -89.7 lake sediment carbonate !18O (Pyrgophorus coronatus) C3, 2 4 Hodell et al., 2005 https://www.ncdc.noaa.gov/paleo/study/6107 Mesoamerica Lake Aljojuca, Mexico 19.1 -97.5 2376 lake sediment carbonate !18O (authigenic) C2, 8 18 Bhattacharya et al., 2015 https://www.ncdc.noaa.gov/paleo/study/17735 Juxtlahuaca Cave, Mexico 17.4 -99.2 934 speleothem carbonate !18O C1, 21 2 Lachniet et al., 2012 https://www.ncdc.noaa.gov/paleo/study/12972 Peten-Itza Lake, Guatemala 17.0 -89.8 110 lake sediment magnetic susceptibility C1, 15 5 Escobar et al., 2012 https://www.ncdc.noaa.gov/paleo/study/13675 Peten-Itza Lake, Guatemala 16.9 -89.8 110 lake sediment carbonate !18O (Cochliopina sp.) 25 Curtis et al., 1998a https://www.ncdc.noaa.gov/paleo/study/5482 Peten-Itza Lake, Guatemala 16.9 -89.8 110 lake sediment carbonate !18O (Cytheridella ilosvayi) 14 Curtis et al., 1998b https://www.ncdc.noaa.gov/paleo/study/5482 Peten-Itza Lake, Guatemala 16.9 -89.8 110 lake sediment carbonate !18O (Pyrgophorus sp.) C3, 5 12 Curtis et al., 1998c https://www.ncdc.noaa.gov/paleo/study/5482 El Gancho Lake, Nicaragua 11.9 -85.9 44 lake sediment carbonate !18O (ostrocod) 5 Stansell, 2012 https://www.ncdc.noaa.gov/paleo/study/13195 Cariaco Basin 10.7 -65.2 -893 marine sediment Ti concentration C3, 17 5 Haug et al., 2001 https://www.ncdc.noaa.gov/paleo/study/2560 Laguna de Ubaque, Colombia 4.5 -73.9 2070 lake sediment % lithics C3, 10 10 Bird et al., 2017 https://www.ncdc.noaa.gov/paleo/study/22275 Laguna Pallcacocha, Ecuador -2.8 -79.2 4200 lake sediment red color intensity C1, 22 0.5 Moy et al., 2002 https://www.ncdc.noaa.gov/paleo/study/5488 Laguna Pallcacocha, Ecuador -2.8 -79.2 4200 lake sediment greyscale C2, 27 0.3 Rodbell et al., 1999 https://www.ncdc.noaa.gov/paleo/study/5487 Palestina Cave, -5.9 -77.4 870 speleothem carbonate !18O C1, 38 3 Apaéstegui et al., 2014 https://www.ncdc.noaa.gov/paleo/study/19301 El Condor Cave, Peru -5.9 -77.3 860 speleothem carbonate !18O 71 Cheng et al. 2013 https://www.ncdc.noaa.gov/paleo/study/13836 Tigre Perdido Cave, Peru (NC-A) -5.9 -77.3 speleothem carbonate !18O C1, 6 29 van Breukelen et al., 2008 https://www.ncdc.noaa.gov/paleo/study/9790 South America Cascayunga Cave, Peru -6.1 -77.2 900 speleothem carbonate !18O 2 Reuter et al., 2009 https://www.ncdc.noaa.gov/paleo/study/8630 Lake Pumacocha, Peru -10.7 -76.1 4300 lake sediment carbonate !18O (authigenic) C2, 9 2 Bird et al., 2011 https://www.ncdc.noaa.gov/paleo/study/11174 Huagapo Cave, Peru -11.3 -75.8 3850 speleothem carbonate !18O C2, 20 5 Kanner et al., 2013 https://www.ncdc.noaa.gov/paleo/study/16405 Diva de Maura, Brazil -12.4 -41.6 680 speleothem carbonate !18O C3, 24 5 Novello et al., 2012 https://www.ncdc.noaa.gov/paleo/study/13670 Quelccaya ice cap, Peru -13.9 -70.8 5670 glacial ice ice !18O C1, 31 10 Thompson et al., 2013 https://www.ncdc.noaa.gov/paleo/study/14174 Pau d'Alho Cave, Brazil -15.2 -56.8 speleothem carbonate !18O C2, 25 1 Novello et al., 2016 https://www.ncdc.noaa.gov/paleo/study/20043 922 Cristal Cave, Brazil -24.6 -48.6 130 speleothem carbonate !18O C1, 34 2 Vuille et al., 2012 / Taylor, 2010

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923 924

925 926

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931

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