Data-Model Comparisons of Tropical Hydroclimate Changes Over the Common Era
manuscript submitted to Paleoceanography and Paleoclimatology
1 Data-Model Comparisons of Tropical Hydroclimate Changes Over the Common Era
2
3 A. R. Atwood1*, D. S. Battisti2, E. Wu2, D. M. W. Frierson2, J. P. Sachs3
4 1Florida State University, Dept. of Earth Ocean and Atmospheric Science, Tallahassee, FL, USA 5 2University of Washington, Dept. of Atmospheric Sciences, Seattle, WA 98195, USA 6 3University of Washington, School of Oceanography, Seattle, WA 98195, USA 7 8 Corresponding author: Alyssa Atwood ([email protected])
9
10 Key Points: 11 x A synthesis of 67 tropical hydroclimate records from 55 sites indicate several regionally- 12 coherent patterns of change in the Common Era 13 x Robust patterns include a regional drying event from 800-1000 CE and a range of tropical 14 hydroclimate changes ~1400-1700 CE 15 x Transient model simulations provide several new insights to the proxy reconstructions 16
manuscript submitted to Paleoceanography and Paleoclimatology
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 reveals 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 dry and/or isotopically enriched conditions in 26 the Asian summer monsoon region, wet and/or isotopically depleted conditions in the central 27 Andes and southern Amazon in South America, and fresher and/or isotopically depleted 28 conditions in the Maritime Continent. We find notable dissimilarities between the regional 29 hydroclimate changes and global- and hemispheric-scale temperature reconstructions, indicating 30 that more work needs to be done to understand the mechanisms of the widespread tropical 31 hydroclimate changes during the LIA. Apropos to previous interpretations of large-scale 32 reorganization of tropical Pacific climate during the LIA, we do not find support for a large-scale 33 southward shift of the Pacific Intertropical Convergence Zone, while evidence for a strengthened 34 Pacific Walker Circulation and/or an equatorward contraction of the monsoonal Asian-Australian 35 rain belt exists from limited geographic regions but require additional paleoclimate constraints. 36 Transient climate model simulations exhibit weak forced long-term tropical rainfall changes over 37 the last millennium but provide several important insights to the proxy reconstructions. 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 year-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 century) and unevenly distributed 54 precipitation records around the globe (Biasutti et al., 2018). However, the efficacy of the 55 climate models used to project future precipitation can be tested against past records. In 56 particular, the last 2,000 years, 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 present climate.
manuscript submitted to Paleoceanography and Paleoclimatology
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, which was largely driven by increased volcanic activity (Atwood et al., 2016; 64 Schurer et al., 2014), though solar, greenhouse gas, land use, and orbital forcings, and internal 65 variability are also thought to have played important roles in driving temperature variations over 66 the last millennium (Otto-Bliesner et al., 2016; Schurer et al., 2013; Swingedouw et al., 2011). 67 Recent compilations of temperature-sensitive proxy data suggest that global mean surface 68 temperatures cooled by approximately 0.2 °C over the last millennium (Consortium, 2013; 69 Consortium, 2019) with slightly larger cooling in the Northern Hemisphere (Hegerl et al., 2007; 70 Mann et al., 2009; Marcott et al., 2013; Moberg et al., 2005) than the Southern Hemisphere 71 (Neukom et al., 2014). Embedded within this global cooling trend, the Little Ice Age (LIA) is 72 characterized as a period of regionally-variable cooling during the second half of the last 73 millennium ~1450-1850 (Consortium, 2013; Cunningham et al., 2013; Kobashi et al., 2011; 74 Masse et al., 2008; Neukom et al., 2014). The spatial and temporal heterogeneity of the LIA 75 reflects the complex array of forcings, feedbacks, and internal variability operating in the climate 76 system (Fernandez-Donado et al., 2013; Kaufman et al., 2009; Lehner et al., 2013). Evidence for 77 amplified cooling in central Greenland (exceeding 1.0 °C) has supported interpretations of Arctic 78 amplified cooling during the LIA (Kobashi et al., 2011), which is also seen in transient climate 79 model simulations of the last millennium (Atwood et al., 2016). 80 In the tropics, widespread changes in rainfall patterns are thought to have occurred during 81 the LIA. In Asia, weakened summer monsoon rainfall and a prevalence of extreme droughts 82 around 1300-1700 CE have been inferred from oxygen isotopes in cave deposits in India and 83 China, from tree ring records from Southeast Asia, and from relative abundances of foraminifera 84 from the Arabian Sea (Anderson et al., 2002; Buckley et al., 2010; Sinha et al., 2011; Wang et 85 al., 2005; Zhang et al., 2008). Some studies have suggested that these rainfall changes 86 contributed to the collapse of several major civilizations in India, China and Southeast Asia 87 (Buckley et al., 2010; Sinha et al., 2011; Zhang et al., 2008). On the opposite side of the Pacific 88 basin in the Southern Hemisphere, an intensified South American summer monsoon ca. 1300– 89 1500 CE and continuing through 1800 CE, has been inferred from records of oxygen isotopes in 90 cave deposits, glacial ice, and authigenic calcite in lake sediments (Vuille et al., 2012). 91 Outside of these major monsoon systems, the nature of LIA hydroclimate changes in 92 other regions of the tropics are more heavily disputed, particularly in and around the tropical 93 Pacific Ocean. For example, different compilations of hydroclimate records from the Pacific 94 basin and surrounding regions have been interpreted as representing either: 1) a southward 95 migration of the mean annual position of the Pacific Intertropical Convergence Zone (ITCZ) 96 (Nelson and Sachs, 2016; Richey and Sachs, 2016; Sachs et al., 2009) and Atlantic ITCZ (Haug 97 et al., 2001), 2) an equatorward contraction of the Asian-Australian monsoons (Denniston et al., 98 2016; Griffiths et al., 2016; Yan et al., 2015), and/or 3) a strengthened Pacific Walker 99 Circulation (Griffiths et al., 2016; Yan et al., 2011; Yan et al., 2015). Importantly, however, the 100 definition of the LIA and the reference period varies widely across these studies; in some cases, 101 the interpretation was based on trends during the LIA (e.g. Richey and Sachs, 2016), in other 102 cases LIA conditions were referenced to the Medieval Climate Anomaly (MCA) in the early part 103 of the last millennium (Haug et al., 2001; Nelson and Sachs, 2016) or to the modern period 104 (Sachs et al., 2009), and in yet other cases, the MCA and modern periods were used 105 interchangeably, depending on data availability (Yan et al., 2015). Standardizing these
manuscript submitted to Paleoceanography and Paleoclimatology
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 ITCZ are highly zonally variable even under a large interhemispheric asymmetry in forcing 117 (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. Here, we collect public data to reconstruct robust spatiotemporal 121 patterns of tropical hydroclimate change over the Common Era using a larger suite of 122 hydroclimate records than has previously been considered. Furthermore, we compare the 123 reconstructed hydroclimate patterns to a large suite of transient climate model simulations of the 124 last millennium to shed light on the mechanisms of the long-term hydroclimate changes, 125 including distinguishing between forced and unforced modes of variability.
manuscript submitted to Paleoceanography and Paleoclimatology
126
127 2. Materials and Methods
128 2.1. Proxy Record Selection and Treatment
129 Tropical hydroclimate records spanning the last millennium were selected from the 130 NCDC/NCEI (https://www.ncdc.noaa.gov/) and PANGAEA (https://www.pangaea.de/) 131 paleoclimate data repositories. A total of 67 speleothem, lake and marine sediment, and tropical 132 glacier records were selected based on the following criteria: 1) the original authors interpreted 133 the proxies as representative of hydroclimate conditions, 2) the records were located between 134 35°S to 35°N, and 3) the records spanned at least 70% of the last millennium (850-1850 CE) 135 with a minimum temporal resolution of 200 years and contained at least one age control point 136 within that period. More stringent record selection criteria were adopted for the cluster analysis 137 and Empirical Orthogonal Function analysis, as detailed in Sections 2.2.1 and 2.2.3. For all 138 analyses, the records were truncated outside of the time period 350-1950 CE and standardized to 139 Z-scores by subtracting the mean and dividing by the standard deviation of the data during this 140 interval. The Z-scores from each record (with zero mean and unit variance) were binned into 141 100-yr windows to focus on the centennial to millennial scale patterns in the data, and to limit 142 the sensitivity of the results to the age model uncertainties in the reconstructions. The sign of the 143 inferred relationship between rainfall amount and the measured values of the proxies was 144 adopted from the original publication of each record. A lake minerology record from Kiritimati 145 Atoll (Higley et al., 2018) was omitted from our analyses due to a substantial revision in the age 146 model between 1200 CE and the core top (J. Conroy, personal communication). Table 1, Fig. 1, 147 and Table S1 provide a summary of the proxy records.
148 Of the 67 records in the compilation, 24% (16 out of 67) are non-water isotope records, 149 consisting of sediment grain size, bulk density, magnetic susceptibility, color, elemental 150 composition (Ti/Ca, percent Mg in calcite), δ13C of organic matter, and relative diatom 151 abundances in lake and marine sediments (see Table S1). The remainder of the records (51 out of 152 67; 76%) are water isotope-based archives, which reflect the isotopic composition (δ2H or δ18O) 153 of precipitation, groundwater, lake water, or near-surface seawater. The stable isotopic 154 composition of water is a powerful tracer of the global hydrologic cycle. Lower (higher) δ2H and 155 δ18O values in meteoric water generally occur during periods of regionally wetter (drier) 156 conditions, as described by the “amount effect”, which is an empirical relationship in which the 157 isotopic composition of precipitation in the tropics is negatively correlated with the amount of 158 precipitation on monthly or longer timescales (Dansgaard, 1964; Risi et al., 2008). Many 159 different components of tropical hydroclimate are represented in our multi-proxy compilation. 160 The isotopic records from closed lake systems generally reflect the local water balance of the 161 lake (i.e. precipitation minus evaporation), while open lake, speleothem, and glacier ice records 162 generally reflect the isotopic composition of precipitation, which integrates information about 163 numerous aspects of the water cycle, including precipitation and evaporation rates, seasonality of 164 precipitation, moisture sources, transport, and water recycling over land. Reconstructions of 165 surface seawater isotope ratios provide a means of characterizing the surface exchange of 166 freshwater and water mass mixing in the ocean, and therefore provide constraints on ocean 167 salinity. Future work should incorporate isotope-enabled model simulations to explicitly track
manuscript submitted to Paleoceanography and Paleoclimatology
168 water isotope ratios in different components of the hydrosphere and thereby provide more 169 quantitative comparisons between the water isotope-based proxy records and model simulations.
170 2.2. Proxy Data Analyses
171 2.2.1. Hierarchal Cluster Analysis
172 Hierarchal cluster analysis (HCA; Eisen et al., 1998) was used to organize the 173 paleoclimate records into groups with similar temporal patterns. In HCA, a binary, hierarchical 174 cluster tree (dendrogram) links pairs of records that are in close proximity based on a specified 175 distance metric; binary clusters are then grouped into larger clusters, forming a hierarchical tree. 176 The height of the linkages in the dendrogram represents the distance between objects. HCA has 177 been successfully used to divide diverse multi-proxy paleoclimate datasets into groups with 178 shared temporal patterns (Kaufman et al., 2016; Shuman et al., 2018). We constructed the 179 hierarchical cluster tree in MATLAB (clusterdata.m) based on Euclidian pairwise distances 180 between records and Ward’s minimum variance method with a maximum of three clusters 181 specified. For studies in which more than one dataset was generated to reconstruct hydroclimate 182 at a given site, all relevant datasets from that study were averaged together. Records with 183 missing data in more than two 100-year windows across the 350-1950 CE period of analysis 184 were excluded from the cluster analysis; records with missing data in d 2 windows were linearly 185 interpolated and extrapolated. The Sletten et al. (2013) speleothem G18O record from Namibia 186 and the Stansell et al. (2012) ostrocod G18O record from Nicaragua were omitted to improve the 187 stability of the cluster analysis. After these adjustments, 41 unique records remained in the 188 cluster analysis. Composite records for each cluster were calculated by averaging all binned data 189 from each cluster. The 95% confidence intervals were calculated based on 1000 iterations of 190 bootstrap sampling with replacement.
191 2.2.2. Regional Composites
192 In regions of high data density, regional composites of the proxy data were constructed to 193 identify the robust regional hydroclimate signals in the reconstructions (Fig. 1). Within each of 194 the selected regions, regional composites were constructed from the Z-scores of the 100-yr 195 binned proxy data. For studies that invoked more than one dataset to reconstruct hydroclimate at 196 a given proxy site, the relevant datasets from that study were averaged together. Some records 197 were omitted from the regional composites based on the original authors interpretation of those 198 records. In particular, we omitted the records from Laguna Pallacocha in the Ecuadorian Andes 199 from the South American composite as these records were interpreted by the original authors as 200 representing ENSO-driven episodes of alluvial deposition, in contrast to the other records from 201 this region.
202 2.2.3. EOF/PC Analysis
203 Empirical Orthogonal Function (EOF)/Principle Component (PC) Analysis was 204 employed to characterize the dominant modes of variability in both the proxy records and models 205 and to compare the patterns of variability between the two. These analyses were conducted on 206 the full tropical dataset, as well as selected regional subsets of the data. The record selection 207 criteria matched that of the cluster analyses in Section 2.2.1. To facilitate data-model
manuscript submitted to Paleoceanography and Paleoclimatology
208 intercomparison, the proxy data were truncated between 850-1850 CE, converted to Z-scores, 209 and binned into 100-yr windows (as in the cluster analysis) before calculating the PCs and EOF 210 loadings from the proxy data. Regional subsets selected for EOF analysis (based on proxy data 211 density) were Asia, South America, and the western Pacific. For the model simulations, 212 precipitation and surface air temperature data from the individual models and ensemble and 213 multi-model means were binned into 100-yr windows to mimic treatment of the proxy data and 214 the EOFs and PCs were calculated from regions bounded by the subsets of proxy data.
215 2.2.4. Epoch 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 (hereafter GISS 121 and GISS 124; 227 Schmidt et al., 2006), which differ in their prescribed solar forcing, MPI-ESM-P (Marsland et al., 228 2003; Raddatz et al., 2007), CSIRO Mk3L version 1.2 (Phipps et al., 2012; Rotstayn et al., 229 2012), and HadCM3 (Collins et al., 2001; Pope et al., 2000; Schurer et al., 2013). The second set 230 of simulations are from the Community Earth System Model (CESM) Last Millennium 231 Ensemble (LME) version 1.1 (with version 5 of the Community Atmosphere Model; CAM5), 232 consisting of a set of 13 experiments with all last millennium climate forcings included (orbital, 233 volcanic, solar, and greenhouse gas), and the following single-forcing experiments: volcanic- 234 only (n=5), solar-only (n=4), orbital-only (n=3), greenhouse gas-only (n=3) and land-use only 235 (n=3), as detailed in Otto-Bliesner et al. (2016). The prescribed forcings are described in Schmidt 236 et al. (2011) and Otto-Bliesner et al. (2016); the volcanic and orbital forcings are also shown in 237 Fig. S1 and S2 of this study.
238 Two different reconstructions of volcanic aerosols and three different methods for 239 implementing aerosols were used in the CMIP5/PMIP3 and CESM simulations analyzed in this 240 study. The volcanic aerosol reconstructions were taken from either (1) the Gao et al. (2008) 241 dataset of stratospheric sulfate loading, where the loading is a function of month, latitude in 10 242 bands, and height from 9 to 30 km at 0.5-km resolution, or (2) the Crowley et al. (2008) dataset 243 of aerosol optical depth (AOD) (Schmidt et al., 2011). Substantial differences exist between the 244 Gao et al. (2008) and Crowley et al. (2008) reconstructions with respect to the global mean and 245 interhemispheric asymmetry of the aerosol forcing (Fig. S1). The CCSM4 and CESM 246 simulations prescribed sulfate loading from Gao et al. (2008), and stratospheric aerosols are 247 prescribed in the models as a fixed single-size distribution in the three layers in the lower 248 stratosphere above the tropopause (Landrum et al., 2013; Otto-Bliesner et al., 2016). The GISS 249 121, GISS 124, MPI, and HadCM3 simulations are forced by a prescribed aerosol effective
manuscript submitted to Paleoceanography and Paleoclimatology
250 radius and AOD from Crowley et al. (2008), which varies a function of height and latitude in the 251 atmosphere. In the CSIRO model, volcanic forcing was imposed as a globally averaged 252 perturbation to total solar irradiance (TSI) that scales with the Crowley et al. (2008) dataset of 253 AOD (Masson-Delmotte, 2013).
254 2.4. Energy Budget Analysis
255 We quantified the hemispheric asymmetry of the forcings and feedbacks in the last 256 millennium simulations during the LIA relative to the MCA by employing the Approximate 257 Partial Radiative Perturbation (APRP) method (Taylor et al., 2007). The APRP method is based 258 on a single-layer, shortwave radiative model of the atmosphere and has been successfully used to 259 quantify the strength of climate forcings and feedbacks in a wide variety of paleoclimate, 260 historical, and future climate simulations (e.g. Atwood et al., 2016; Crucifix, 2006; Hwang et al., 261 2013). In the APRP method, the influence of changes in surface albedo, shortwave absorption, 262 and scattering on the TOA energy budget are diagnosed at every model grid cell using all-sky 263 and clear-sky model output. Changes in top-of-the-atmosphere (TOA) shortwave energy fluxes 264 were decomposed into individual climate forcing and feedback terms as in Atwood et al. (2016), 265 except that here the analysis is performed separately for each hemisphere. Energy fluxes are 266 reported as NH minus SH, where negative terms indicate more NH cooling and positive terms 267 indicate more SH cooling. Radiation data for CSIRO and the CESM simulations were not 268 available at the time of analysis and thus these simulations were omitted from the APRP 269 analyses.
270
271 3. Results
272 3.1. Cluster Analysis of Proxy Records
273 The hierarchal cluster analysis indicates that the most common pattern, found in 41% of 274 the tropical hydroclimate records analyzed, is a millennial-scale isotopic depletion trend 275 beginning around 1000 CE and reversing around 1800 CE (Cluster 1 in Fig. 2A-B), which 276 generally tracks global temperature trends over the last millennium (Fig. 3H; Consortium, 2019). 277 This cluster includes proxies from nearly every region represented in the proxy compilation (Fig. 278 2C). We will demonstrate that this signal is not the dominant pattern in any regional compilation, 279 however, and thus it likely does not reflect regional changes in precipitation. Instead, it may 280 reflect tropic-wide changes in the isotopic composition of water vapor (Falster et al., 2019), 281 possibly driven by the cooling of tropical sea surface temperatures over the last millennium.
282 The next most common pattern, found in 37% of the records in the cluster analysis, 283 including the majority of records from the Maritime Continent and the southwestern tropical 284 South America (including the central Andes and the southern Amazon), is dominated by a drying 285 (and/or isotopic enrichment) trend from 1600 CE to present (Cluster 2; Fig. 2D,E). An opposing 286 trend is observed in the third cluster, which contains 22% of the records, including those from 287 northern and eastern South America, and parts of Mesoamerica and East Asia (Cluster 3; Fig. 288 2F,G). In this cluster, the dominant pattern is drying (and/or isotopic enrichment) beginning ca.
manuscript submitted to Paleoceanography and Paleoclimatology
289 1300 CE and peaking during the LIA around 1600-1700 CE, followed by a wetting trend to 290 present.
291 3.2. Regional Composites of Proxy Records
292 When the records are compiled by region, several robust regionally-coherent 293 hydroclimate patterns emerge that provide more insight into the patterns found in the cluster 294 analysis. From 800-1000 CE, encompassing the early part of the MCA, three lake records from 295 two islands in the eastern equatorial Pacific indicate notably dry (and/or isotopically enriched) 296 conditions relative to the average over the Common Era (Fig. 3A), a pattern that is also found in 297 half (5/10) of the records from Mesoamerica, especially those clustered in the northern part of 298 the region (Fig. 3B). The remaining 5/10 Mesoamerican records, which extend father south, 299 indicate a distinctly different centennial pattern, characterized by a steady drying (and/or isotopic 300 enrichment) trend from 1200 CE to present (Fig. 3C).
301 During the second half of the last millennium, records from the major monsoon systems 302 of Asia and South America demonstrate antiphased behavior. In South and East Asia, nearly all 303 records demonstrate anomalously dry (and/or isotopically enriched) conditions ca. 1400-1700 304 CE followed by a coherent wetting (and/or isotopic depletion) trend to present (Fig. 3F). In 305 contrast, nearly all records from the central Andes and Southern Amazon in South America 306 indicate anomalously wet (and/or isotopically depleted) conditions ca. 1500-1700 CE, followed 307 by a drying trend to the present (Fig. 3E). During the LIA, the South American records are 308 characterized by an opposing pattern between the southwestern sites and the northern and eastern 309 sites: southwestern South America experienced a wetting (and/or isotopic depletion) trend from 310 1300-1600 CE while northern South America and Nordeste Brazil experienced a drying (and/or 311 isotopic enrichment) trend (Fig. 3D,E). From 1600 CE to present, northern South America and 312 Nordeste Brazil remained anomalously dry while southwestern South America experienced a 313 drying trend.
314 Hydroclimate records from the Maritime Continent generally oppose the hydroclimate 315 patterns observed in Asia. Namely, the composite of marine records from the Maritime 316 Continent indicates isotopically depleted precipitation and seawater (and by inference lower 317 salinity) from 1500-1700 CE (Fig. 3G), followed by an isotopic enrichment trend to present. The 318 coherence is especially pronounced in the four marine records from the Makassar Strait and 319 south of the Philippine Islands (Fig. S3D). A single speleothem record from northern Australia 320 reflects the Asian monsoon trends to some extent, with drier (and/or isotopically enriched) 321 conditions ca. 1400-1600 CE, relative to much of the prior 500 years, followed by a wetting 322 (and/or isotopic depletion) trend to present (Fig. S3C). These features provide some evidence of 323 coherence between the northern and southern margins of the Asian-Australian rain belt. 324 However, while the LIA conditions in Asia are anomalous with respect to the average conditions 325 of the Common Era, this is not the case for the northern Australia record. Additional 326 hydroclimate records from northern Australian are needed to confirm the regional patterns in this 327 area, as only a single record that spanned the last millennium was publicly available at the time 328 of analysis.
329 3.3. Simulated Tropical Precipitation Changes over the Last Millennium
manuscript submitted to Paleoceanography and Paleoclimatology
330 To investigate the mechanisms of the robust hydroclimate changes inferred from the 331 proxy records, including distinguishing between forced and unforced variability, we compared 332 the reconstructions to a large suite of transient climate model simulations of the last millennium. 333 The forced response of tropical rainfall is weak on centennial timescales in the models, as 334 indicated by the small (< 10%) LIA precipitation changes relative to the models’ climatology in 335 the CMIP5/PMIP3 multi-model mean (Fig. 4B) and in the CESM ensemble mean (Fig. 4D). The 336 weak forced responses are due to large differences in tropical rainfall anomalies during the LIA 337 across different models (Fig. S4A) and across different ensemble members of the same model 338 (CESM; Fig. S4B).
339 However, the precipitation changes in the CMIP5/PMIP3 multi-model average show 340 several areas of broad agreement with the ensemble average changes in CESM (Fig. 4), 341 suggesting similar forced precipitation responses across models in these regions. For example, 342 during the LIA relative to the MCA, nearly all CMIP5/PMIP3 models and all CESM ensemble 343 members simulate large-scale drying in South and East Asia, drying over the Sahel, wetting in 344 the equatorial Atlantic Ocean that extends to the northeastern edge of South America, drying 345 along the southwestern edge of the South Pacific Convergence Zone, and drying of the mid- and 346 high latitudes (Fig. S4). These features are qualitatively similar between the first and second half 347 of the LIA, but are more pronounced in the second half of the LIA (Fig. S5). The CESM single- 348 forcing experiments indicate that these tropical rainfall changes are largely driven by orbital 349 forcing in the models, with drying over Asia and wetting along the northeast coast of South 350 America also driven by volcanic forcing (Fig. 5). Further details of the tropical hydroclimate 351 responses to orbital forcing are provided in the Supplemental Information. Notably, in tropical 352 South America, the models do not simulate a strengthening of the South American Summer 353 Monsoon (SASM) during the LIA, and few coherent rainfall changes occur over South America 354 in the simulations during the height of the monsoon season in austral summer (DJF; Fig. S6A,C).
355 In the tropical Pacific, the precipitation changes are notably different between the CESM 356 ensemble average and CMIP5/PMIP3 multi-model average, tracking differences in the tropical 357 cooling patterns in the two sets of model simulations over the last millennium (Fig. S7). Namely, 358 the CMIP5/PMIP3 multi-model mean precipitation response features a robust meridional pattern, 359 with drying in the northern branch of the Pacific ITCZ, while the northern subtropical Pacific 360 and southeastern Pacific become wetter, consistent with a weakened local Hadley circulation in 361 this region (Fig. 4; Fig. S7A). This meridional rainfall pattern is consistent with the change in 362 tropical surface temperatures which features enhanced cooling in the northern tropical Pacific, 363 including under the northern branch of the Pacific ITCZ (Fig. S7B). In contrast, the forced 364 precipitation trend in the CESM simulations features an enhanced Walker Circulation, with 365 drying that extends to the far western Pacific (Fig. S7C-D). The internal variability in the Indo- 366 Pacific is generally larger than the forced signal in CESM however (Fig. S8), and thus the long- 367 term strengthening of the Walker circulation is only weakly significant (see the lack of stippling 368 in much of the Indo-Pacific in Fig. 4C,D). The internal variability is characterized by centennial 369 variations in the Walker Circulation, while the forced response is characterized by a weak 370 equatorial Pacific drying trend and wetting in the Maritime Continent west of Papua New 371 Guinea. The forced response in CESM reflects an enhanced Walker Circulation that is driven by 372 a combination of orbital and volcanic forcing, and drying in the West Pacific/wetting in the 373 Maritime Continent that is driven by land use forcing (Fig. 5). However, this tropical Pacific 374 response to the last millennium forcings appears to be unique to CESM since none of the
manuscript submitted to Paleoceanography and Paleoclimatology
375 CMIP5/PMIP3 models simulate a strengthening of the Walker Circulation over the last 376 millennium (Fig. S4; Fig. 6; though multiple ensemble members of the CMIP5/PMIP3 377 simulations would be required to identify the true forced response in those models). The forced 378 response and the large unforced centennial variations in the Walker Circulation in CESM may be 379 related to stronger tropical Pacific ocean-atmosphere coupling in that model, as suggested by the 380 larger interannual to centennial variability in the tropical Pacific in the CESM last millennium 381 simulations as compared to the CMIP5/PMIP3 last millennium simulations (Fig. S10). The 382 magnitude of tropical Pacific climate variability varies widely across models, ranging from weak 383 variability in the GISS and CSIRO last millennium simulations, in which the variance of Niño 3 384 SST anomalies is around 30% of that observed over the last 40 years, to large variability in 385 CESM, in which the variance of Niño 3 SST anomalies is around 150% of that observed (Fig. 386 S11).
387 3.4. Data-Model Intercomparison
388 Point-to-point comparisons between the proxy data and models are generally not 389 meaningful because of the structural uncertainties in simulated precipitation climatology 390 associated with the coarse spatial resolution and large tropical mean state biases in the models 391 (e.g. PAGES Hydro2k Consortium, 2017). In addition, most of the proxy records included in this 392 analysis are based on water isotopes, hence they record regional, as well as local, hydroclimate 393 variations. For these reasons, we focus on comparing the proxy reconstructions to simulated 394 precipitation over a regional scale.
395 The simulations highlight three large-scale hydroclimate features that are worthy of 396 consideration when interpreting the proxy records. The first is the large-scale drying pattern over 397 South and East Asia from the MCA to LIA, which appears in both the proxies and models (Fig. 398 4) and is driven by volcanic and orbital forcing in the models (Fig. 5). Importantly, however, the 399 temporal evolution of the drying trend from the LIA to present is notably different between the 400 models and proxies; a distinction that becomes more evident when an EOF analysis is performed 401 only on the Asian records and precipitation fields (Fig. S12). The proxy data (comprised of 402 Asian speleothem and lake sediment records) features a coherent drying trend that peaks around 403 1600 CE, followed by a reversal (wetting trend) through 1850 CE (Fig. S12A). In contrast, the 404 models, particularly the CMIP5/PMIP3 models, are characterized by a persistent drying trend 405 over Asia through the last millennium and enhanced drying from 1600-1850 CE (Fig. S12B-C). 406 This drying trend is apparent in 5/6 CMIP5/PMIP3 models, while in the CESM ensemble 407 members, large internal variability in this region generally overwhelms the forced response (data 408 not shown).
409 A second large-scale hydroclimate feature in the models that may be relevant to the proxy 410 reconstructions is increased rainfall in the South American monsoon entrance region, upstream 411 of the SASM. The proxy records in the central Andes and southern Amazon indicate isotopically 412 depleted conditions ca. 1500-1700 CE, consistent with a greater degree of upstream rainout 413 (Vuille and Werner, 2005). Though there is no indication of enhanced SASM activity in the 414 models, as has previously been inferred from the South American proxy records, precipitation 415 increases in the equatorial Atlantic and extends into the monsoon entrance region in nearly all 416 models (Fig. 4; Fig. S4), driven by orbital and volcanic forcing in CESM (Fig. 5). The increase 417 in equatorial Atlantic rainfall under orbital forcing occurs alongside a contraction in the Atlantic
manuscript submitted to Paleoceanography and Paleoclimatology
418 ITCZ (i.e. a reduced seasonal migration of the Atlantic ITCZ), which we discuss in more detail 419 in Section 4.3.3.
420 The third large-scale hydroclimate feature that is shared between the proxy data and 421 select CESM simulations is wetter LIA conditions in the Maritime Continent coupled with dry 422 conditions in the central tropical Pacific (Fig. 4; Fig. S4). In CESM, this feature occurs in 8/13 423 ensemble members (Fig. S4) and also occurs the ensemble mean response (Fig. 4), associated 424 with an enhanced Pacific Walker Circulation (Fig. S8). However, central Pacific drying is only 425 indicated by a single record and there is pronounced data-model mismatch in the eastern 426 equatorial Pacific (the three records from the Galápagos Islands indicate wetter conditions during 427 the LIA, while the models indicate drying or no change; Fig. 4; Fig. 3A). Thus, additional 428 hydroclimate records from the Pacific cold tongue are needed to support any inference of a 429 change in the Walker Circulation during the LIA.
430 3.5. Small but Robust Zonally-Averaged Tropical Rainfall Response in Models
431 Despite large differences in regional precipitation across models, some coherence is 432 observed in the zonally averaged context: all models demonstrate a small zonal mean southward 433 shift in tropical precipitation (20ºS:20ºN) during the LIA relative to the MCA (Fig. 7). The 434 change in the precipitation asymmetry ('PNH-SH; the precipitation averaged between 0-20°N 435 minus 0-20°S) is small (d 0.3 mm/day), representing a -15% to +10% change in precipitation 436 asymmetry across models. To put the magnitude of these changes into context, the multi-model 437 mean change in the tropical precipitation centroid (defined as the median of zonal average 438 precipitation from 20ºS to 20ºN) is -0.03º latitude in the last millennium simulations (data not 439 shown), which is approximately 6% as large as the -0.53º multi-model mean shift in the PMIP2 440 and PMIP3 Last Glacial Maximum simulations and 1% as large as the -2.48º shift in 1 Sv North 441 Atlantic meltwater simulations (Atwood et al., 2020).
442 This robust zonal mean precipitation shift occurs in association with anomalous 443 northward cross-equatorial atmospheric energy transport in all models that is primarily driven by 444 volcanic forcing (i.e. greater volcanic aerosol loading in the northern hemisphere) and land use 445 changes (i.e. cropland expansion in Eurasia), and is amplified by the surface albedo feedback 446 (e.g. through the growth of Arctic sea ice in response to orbital forcing) and longwave feedbacks 447 (Fig. S13). While these zonal mean changes are small ('PNH-SH d 0.3 mm/day; 'AHT)=0 < 0.2 448 W/m2), they demonstrate a perceptible influence of the hemispherically-asymmetric forcings 449 (namely volcanic forcing and Eurasian cropland expansion) and feedbacks (e.g. the surface 450 albedo feedback) on the zonal mean tropical precipitation response in the models.
451 The fact that the zonal mean precipitation changes are fairly consistent across models 452 indicates that the net hemispheric asymmetry in the forcings and feedbacks are generally 453 consistent across models and also indicates that the zonal mean precipitation response is less 454 sensitive to structural uncertainties in the models than the regional rainfall responses are. 455 However, the zonal mean rainfall changes are clearly irrelevant to our analysis of regional 456 rainfall patterns, since they are too small to be of consequence for interpreting the changes seen 457 in the proxy records and since zonal mean changes have little bearing on regional rainfall 458 patterns (e.g. Atwood et al., 2020).
manuscript submitted to Paleoceanography and Paleoclimatology
459
460 4. Discussion
461 Three robust signatures of tropical hydroclimate change over the Common Era have 462 emerged from our analyses of the centennial patterns in the proxy records. These are: (1) a pan- 463 tropical isotopic depletion trend in the water isotope based records over the past millennium 464 (Cluster 1), (2) anomalously dry conditions in the eastern tropical Pacific Ocean and 465 Mesoamerica between 800-1000 CE, and (3) a range of hydroclimate anomalies across the 466 tropics from 1400-1700 CE, including drier and/or isotopically enriched conditions in South and 467 East Asia, wetter and/or isotopically depleted marine conditions in the Maritime Continent, and a 468 dipole pattern in South America. We discuss each of these robust regional patterns in more 469 detail, below.
470 4.1. Isotopic Depletion Trend in Pan-Tropical Records Over the Last Millennium
471 The dominant pattern of centennial hydroclimate variability, found in 41% of the tropical 472 hydroclimate records analyzed (14 out of 17 of which are based on water isotopes) and appearing 473 in nearly every region of the tropics represented in the proxy compilation, is an isotopic 474 depletion trend beginning around 1000 CE and reversing around 1800 CE (Fig. 2A-C), which 475 tracks global temperature trends over the Common Era (Fig. 3H; PAGES2k Consortium, 2013; 476 PAGES2k Consortium, 2019; Emile-Geay et al., 2017). In model simulations, global cooling of 477 0.2-0.3 °C over the last millennium (Fig. S14A) is accompanied by a perceptible but very weak 478 decrease in tropical area-averaged precipitation (Fig. S14B; 'PLIA - MCA = -0.28 ± 0.08% [1V] and 479 -0.23 ± 0.04% for the CMIP5/PMIP3 and CESM runs, respectively). However, the magnitude of 480 these precipitation changes would be far too small to be detectable in non-water isotope based 481 proxy archives (which indirectly reflect changes in precipitation), or in water isotope based 482 records via the amount effect. Water isotope tracers, however, are sensitive to both precipitation 483 and evaporative fluxes (the latter of which supply the atmospheric water vapor reservoir from 484 which all precipitation is derived). Thus, the millennial-scale isotopic depletion trend observed in 485 many of the proxy records is likely not associated with changes in precipitation, but rather with 486 decreased evaporation rates (and increased fractionation during evaporation) under cooler 487 tropical sea surface temperatures (Risi et al., 2010). This hypothesis is supported by the models’ 488 robust pan-tropical cooling trend over the last millennium (Fig. S8; Fig. S9) and is currently 489 being explored using model simulations with water isotope tracers (Falster et al., 2019). The 490 response of the global water cycle to cooling over the last millennium—and how those changes 491 are recorded in water isotope tracers—is worthy of further investigation using such tools.
492 4.2. Dry Conditions in the Eastern Pacific and Mesoamerica ~800-1000 CE
493 One of the robust regional hydroclimate signals to emerge from our analysis is the 494 pronounced drying event from 800-1000 CE recorded in a series of lake sediment records from 495 the Galápagos Islands in the eastern equatorial Pacific (Fig. 3A; 1˚S, 90-91˚W), and in many lake 496 and speleothem records from Mesoamerica (Fig. 3B). This signal is evident in a variety of 497 different types of hydroclimate records, including both minerology and water isotope based 498 sedimentary records, suggesting that it is driven by local changes in precipitation in these 499 regions. This event is apparent in the regional composites from the eastern equatorial Pacific
manuscript submitted to Paleoceanography and Paleoclimatology
500 (Fig. 3A) and northern Mesoamerica (Fig. 3B) as well as in the cluster analysis (Fig. 2A-C). 501 Evidence of a pronounced drying event during this interval has also been found in a sediment 502 minerology record from a brackish lake in Kiritimati in the central equatorial Pacific (Higley et 503 al., 2018).
504 A period of extended drought across Mesoamerica from 800-1050 CE has been inferred 505 in earlier studies and implicated in major social disruptions in pre-Columbian Mesoamerican 506 civilizations (e.g. Bhattacharya et al., 2017; Hodell et al., 2005). The consistency in timing 507 between the dry intervals in Mesoamerica and the tropical Pacific cold tongue is suggestive of a 508 dynamical linkage between the eastern Pacific climatology and the Mesoamerican monsoon 509 during this time.
510 4.3. Widespread Tropical Hydroclimate Changes during the LIA ~1400-1700 CE
511 The second half of the last millennium is characterized by a set of pronounced and 512 synchronous hydroclimate changes throughout the tropics from 1400-1700 CE, including: 1) 513 isotopically enriched (drier) conditions in South and East Asia, 2) isotopically depleted (fresher) 514 surface ocean conditions in the Maritime Continent, and 3) a dipole pattern in tropical South 515 America. We investigate each of these regional signals and discuss their possible interpretations, 516 below.
517 4.3.1. Dry/Isotopically Enriched Conditions in South and East Asia
518 Hydroclimate reconstructions from South and East Asia generally indicate isotopically 519 enriched (drier) conditions from 1400-1700 CE relative to present and to the background 520 conditions of the Common Era, followed by progressively depleted (wetter) conditions to present 521 (Fig. 3F). LIA drying in this region is also apparent in the LIA-MCA epoch analysis (Fig. 4), and 522 in the EOF/PC analysis (Fig. S12A). Because all of the proxy records from Asia are water 523 isotope records (ten speleothem and one lake carbonate record) (Fig. 3F), which integrate climate 524 information on large spatial scales, the widespread signature of isotopic enrichment in these 525 records likely reflects reduced upstream rainout associated with a weakened South, and possibly 526 East Asian, monsoon during the LIA.
527 This interpretation is consistent with past studies that have characterized the LIA as a 528 time of weakened South and East Asian summer monsoon strength, marked by a series of 529 monsoon droughts in India, China, and Southeast Asia (Sinha et al., 2011; Zhang et al., 2008). 530 Proposed mechanisms for the weakened monsoons include widespread extratropical NH cooling 531 that drove a southward shift of the rain belt in the Indian Ocean (Sinha et al., 2011) and reduced 532 land-sea temperature contrast between Asia and the tropical Indo-Pacific (Zhang et al., 2008). 533 We do indeed see weakening of the South and East Asian monsoons over the last millennium in 534 the models under enhanced cooling over land, particularly in the NH, that is driven by orbital and 535 volcanic forcing (Fig. 4; Fig. 5; Fig. S7). However, the temporal pattern of these changes is 536 notably different between the reconstructions and models. In the reconstructions, aridity in the 537 Asian records peaks around 1600 CE, followed by a wetting trend to the present era, while in the 538 models, the drying trend extends to at least 1800 CE (Fig. S12). Thus, though the models 539 indicate that the weakening of the Asian monsoons during the LIA may have been driven in part
manuscript submitted to Paleoceanography and Paleoclimatology
540 by orbital and/or volcanic forcing, there are clear limitations in the models’ ability to capture this 541 important feature of tropical hydroclimate change over the last millennium.
542 4.3.2. Wet/Isotopically Depleted Conditions in the Maritime Continent
543 Further evidence that the period from 1400-1700 CE was characterized by widespread 544 tropical hydroclimate changes comes from the Maritime Continent. In this region, marine records 545 generally indicate isotopically depleted precipitation and surface ocean conditions around 1400- 546 1700 CE, relative to present and to the background conditions over the Common Era (Fig. 3G). 547 One of the six marine records is a hydrogen isotope leaf wax record from the West Sulawesi 548 margin, which is a proxy for the isotopic composition of precipitation in the region (Tierney et 549 al., 2010). The remaining records are seawater G18O reconstructions based on the G18O and 550 Mg/Ca ratios in planktonic foraminifera. Four out of the six marine records from the Maritime 551 Continent demonstrate especially strong coherence over the last millennium, characterized by a 552 steady trend to more depleted (fresher/wetter) conditions from 1200 to 1700 CE, followed by a 553 drying trend to present (Fig. S3D). That this pattern is not shared by all marine records in this 554 region is perhaps expected given the complexity of the surface ocean circulation in the 555 Indonesian Archipelago. The inferred freshening across a broad sector of the Maritime Continent 556 generally opposes the drying pattern seen in the South and East Asian record composite (Fig. 3F) 557 and also somewhat opposes the weak drying pattern seen in the single record from northern 558 Australia (Fig. S3C).
559 Notably, the reconstructed hydroclimate trends in the Maritime Continent are sensitive to 560 the type of proxy archive (Fig. S3); namely, this LIA signal is absent in many terrestrial records 561 from the region, including two out of three lake sediment records (Fig. S3E) and three out of 562 three speleothem records (Fig. S3F) from Indonesia and Borneo. These differences suggest that 563 large-scale changes in moisture balance in the region may have been integrated into a regionally- 564 coherent seawater G18O signal, while terrestrial records reflect more localized precipitation 565 changes and/or changes in moisture source and trajectories (e.g. Carolin et al., 2013). There is 566 considerable spatial heterogeneity in rainfall throughout the Indonesian archipelago, both in the 567 annual mean and seasonal rainfall (As-syakur et al., 2013), which has been attributed to a 568 complex interplay between the ocean, islands, topography, and the multitude of regional 569 atmospheric circulation patterns in this region. Despite this complexity, marine reconstructions 570 from the Maritime Continent provide a robust regional signature of tropical hydroclimate change 571 during the LIA.
572 4.3.3. Dipole Pattern in South American Hydroclimate
573 The final piece of evidence that the LIA was marked by widespread changes in tropical 574 hydroclimate comes from tropical South America, where the reconstructions show a distinct 575 dipole pattern between the southwestern and northeastern sites between the LIA and MCA (Fig. 576 4). Records from the central Andes and the southern Amazon generally indicate wet (and/or 577 isotopically depleted) conditions from 1400-1700 CE, followed by a pronounced drying trend to 578 the present (Fig. 3E), while records from northern South America and Nordeste Brazil display a 579 coherent drying (and/or isotopic enrichment) trend from 1300-1700 CE, followed by persistently 580 dry conditions (Fig. 3D). This dipole pattern is also apparent in the cluster analysis (Fig. 2D-G), 581 the epoch analysis of the LIA versus MCA (Fig. 4), and the EOF/PC analysis (Fig. 6A).
manuscript submitted to Paleoceanography and Paleoclimatology
582 Because all nine proxy records in the southwestern composite are water isotope records 583 (Fig. 3E), which integrate climate information over large spatial scales, the widespread signature 584 of isotopic depletion in these records during the LIA is likely reflective of increased upstream 585 rainout (either under enhanced SASM activity or increased rainfall in the monsoon entrance 586 region), rather than changes in local rainfall amount (Vuille et al., 2012). The model results 587 presented here do not support enhanced rainfall over the Amazon basin as the cause of the LIA 588 isotopic depletion pattern in the central Andes and southern Amazon. Instead, they support 589 upstream rainout in the monsoon entrance region in northeastern South America, as has also 590 been found in isotope-enabled simulations with CESM (Orrison and Vuille, 2020).
591 Another possibility is that precipitation changes in tropical South America were caused 592 by changes in tropical Pacific SSTs. However, if ENSO is the template, then the hydroclimate 593 anomalies in northern South America and the Nordeste should be in phase with those in the 594 central Andes, e.g. the warm phase of ENSO causes a drying in far northern South America and 595 the Nordeste (Davey et al., 2014) as well as drying and isotopic enrichment in the central Andes 596 and Amazon basin (Vuille and Werner, 2005), which is inconsistent with the proxy data.
597 An alternative explanation of the tropical South American records during the LIA is a 598 southward shifted Atlantic ITCZ, as the ITCZ serves as the major conduit for moisture transport 599 in the SASM region and responds sensitively to changes in the cross-equatorial SST gradient 600 (Chiang, 2009; Chiang et al., 2002). There is ample evidence showing co-variation in the 601 Atlantic ITCZ and isotopic variations in the central Andes. For example, during the millennial- 602 scale warm and cold events of the last ice age (i.e. Dansgaard-Oeschger oscillations) shifts of the 603 Atlantic ITCZ, inferred from reflectance records from the Cariaco Basin (Deplazes et al., 2013), 604 closely paralleled changes in the strength of the SASM (inferred from speleothem G18O records 605 from the central Andes (Kanner et al., 2012)). However, the paleoclimate data from Nordeste 606 Brazil is difficult to reconcile with the interpretation of a meridional shift of the Atlantic ITCZ 607 during the LIA under northern tropical Atlantic cooling (Novello et al., 2012), as we would 608 expect to see a strong antiphased relationship between the northern sites and Nordeste Brazil in 609 response to a meridional shift of the ITCZ, when in fact we see the opposite.
610 A close in-phase relationship between hydroclimate records from the Cariaco Basin and 611 Nordeste Brazil, such as observed over the last millennium, has also been documented from the 612 last glacial period through the Holocene (Cruz et al., 2009; Novello et al., 2012). In those studies, 613 the authors theorized that during times of enhanced SASM activity, intense convection in the 614 Amazon basin produces compensating large-scale subsidence in surrounding regions (more 615 specifically, enhanced convective heating in the southwest Amazon intensifies the Bolivian 616 High-Nordeste Low pressure system, which leads to increased subsidence and decreased summer 617 precipitation over northeast Brazil) (Lenters and Cook, 1997; Novello et al., 2012).
618 We present a plausible alternative interpretation of the covariation between the northern 619 South America and Nordeste Brazil records, namely that these records depict contractions and 620 expansions of the Atlantic ITCZ, driven by coordinated changes in the northward and southward 621 migrations of the Atlantic Ocean rain belt through the year. The Cariaco Basin sits beneath the 622 northward extent of the western Atlantic ITCZ (Fig. 1; Deplazes et al., 2013), while a substantial 623 fraction of annual rainfall in Nordeste Brazil occurs when the Atlantic ITCZ is at its 624 southernmost position between March and May (Novello et al., 2012), thus rainfall in these
manuscript submitted to Paleoceanography and Paleoclimatology
625 regions should be sensitive to changes in the northward and southward seasonal extremes of the 626 ITCZ, respectively.
627 Indeed, recent modeling studies have indicated that the ocean rain bands expand and 628 contract in response to changes in the interhemispheric temperature gradient under precessional 629 forcing (Singarayer et al., 2017). The mechanism of the expansion/contraction in the models 630 relates to the differential response of the hemispheres to insolation forcing. Since 1000 CE, 631 insolation has decreased in boreal summer/fall when the Atlantic ITCZ is farthest north (Fig. S2), 632 and thus the ITCZ is not pulled as far away from the equator today compared to 1000 CE. In 633 boreal winter/spring, when the ITCZ is farthest south, insolation has increased, and because it 634 does so in both the northern and southern hemispheres, although the southern hemisphere warms, 635 the northern hemisphere warms more because of the greater proportion of land there. 636 Consequently, the boreal winter ITCZ shifts less far south in response. If these principles 637 prevailed over the last millennium, as boreal summer/fall insolation decreased and boreal 638 winter/spring insolation increased (Fig. S2), the annual mean Atlantic ITCZ would be expected 639 to contract, consistent with the drier conditions observed in both the Cariaco Basin and Nordeste 640 Brazil (Fig. 3D).
641 Additional support for this hypothesis is found in the early Holocene, when hydroclimate 642 records indicate wetter conditions in the Cariaco Basin and Nordeste Brazil in the early Holocene 643 relative to today (Cruz et al., 2009). During the early Holocene, boreal summer insolation was 644 higher and boreal winter insolation was lower than present, and thus the boreal summer ITCZ 645 would be expected to shift farther north and the boreal winter ITCZ would be expected to shift 646 farther south, resulting in an expansion of the ITCZ, which is consistent with the wetter 647 conditions recorded in both the Cariaco Basin and Nordeste Brazil during this time (Cruz et al., 648 2009). In model simulations, this behavior of the Atlantic ITCZ has been documented in time 649 slice simulations of the last glacial cycle with the HadCM3 coupled climate model (Singarayer et 650 al., 2017). In the present study, we see this behavior as a contraction of the Atlantic ITCZ over 651 the last millennium in both the full forcing and orbital-only simulations with CESM (Fig. 4; Fig. 652 5; Fig. S4B), as well as in a number of the CMIP5/PMIP3 simulations (Fig. S4A). The opposing 653 seasonal rain band shifts are clear in the solstice seasons of all models, i.e. a northward shift in 654 DJF and southward shift in JJA (see Fig. S6 for the ensemble means; individual models not 655 shown), though they are more muted in the annual mean response of the CMIP5/PMIP3 models 656 (Fig. S4A).
657 While the simulated drying trend along the northern and southern edges of the Atlantic 658 ITCZ does not extend over the South American continent in most models, the large spatial scale 659 of this hydroclimate feature warrants serious consideration in interpreting the proxy 660 reconstructions, particularly given the large uncertainties associated with the coarse spatial 661 resolution and large tropical mean state biases in the models (PAGES Hydro2k Consortium, 662 2017). The clear co-variation of paleo-hydroclimate records from northern South America and 663 Nordeste Brazil over the precessional cycle and during the last millennium, in concert with the 664 simulated contraction/expansion of the Atlantic ITCZ in response to precessional forcing in a 665 diverse array of models, provides compelling evidence that a contraction of the Atlantic ITCZ is 666 a plausible mechanism of the observed hydroclimate changes in northern South America and 667 Nordeste Brazil over the last millennium. Notably, the central Andes and southern Amazon 668 records more closely track the Asian monsoon records (with an antiphase relationship) than they
manuscript submitted to Paleoceanography and Paleoclimatology
669 do the records from northern South America and Nordeste Brazil, especially during the reversal 670 of the LIA hydroclimate trends around 1600 CE, suggesting that the tropical Atlantic rain band 671 evolved somewhat differently from rainfall in monsoon regions over the last millennium.
672 4.3.4. Correspondence Between the Tropical Hydroclimate and Temperature Changes
673 As discussed above, the period from 1400-1700 CE stands out as a time of pronounced 674 tropical hydroclimate change across the tropics. These changes include drier (and/or isotopically 675 enriched conditions) in the Asian monsoon region, wetter (and/or isotopically depleted) 676 conditions in the central Andes and southern Amazon, and fresher (and/or isotopically depleted) 677 conditions in the Maritime Continent. Previous studies have highlighted the general 678 correspondence between these hydroclimate changes and changes in the large-scale (global or 679 northern hemisphere mean) temperature or changes in the interhemispheric temperature gradient 680 (e.g. Sinha et al., 2011; Vuille et al., 2012; Yan et al., 2011). However, we find notable 681 dissimilarities between the timing of these hydroclimate changes and the evolution of global 682 mean temperature over the last millennium based on recent temperature reconstructions (Fig. 3H; 683 PAGES2k Consortium, 2019). In particular, the hydroclimate composites from Asia (Fig. 3F), 684 southwestern South America (Fig. 3E), and the Maritime Continent (Fig. 3G) all indicate a 685 reversal in the LIA hydroclimate trends around 1600-1700 CE (Fig. 3E-G; Fig. S12A; Fig. S3D), 686 while the reversal of global cooling occurs substantially later, around 1800 CE, in the ensemble- 687 median of the PAGES2k Consortium (2019) reconstructions (Fig. 3H). There is substantial 688 uncertainty across the statistical methods used to reconstruct global mean surface temperature, 689 however, and one of the seven reconstruction methods (the Bayesian hierarchical model) 690 indicates an early termination to the LIA around 1600-1700 CE, mirroring the pattern observed 691 in the hydroclimate reconstructions.
692 In the proxy records, the reversal of the LIA hydroclimate trends around 1600-1700 CE is 693 remarkably consistent across the regional composites of South and East Asia, South America, 694 and the Maritime Continent, and is also reflected in Clusters 2 and 3 from the cluster analysis 695 (Fig. 2D,F). While the age model uncertainty in the reconstructions must be taken into account 696 (particularly in the 14C-dated marine and lake sediment records with few age control points; 697 Table S1), the Asian and South American composites are dominated by well-dated speleothem 698 records, and thus the synchronicity of this feature across a wide range of proxy archives from 699 different regions of the tropics supports the inference of a widespread reversal in LIA 700 hydroclimate trends around 1600-1700 CE.
701 The fact that Cluster 1 from the hydroclimate records tracks the reconstructed global 702 temperature trends over the last millennium is consistent with our interpretation of this cluster as 703 reflecting reduced evaporation rates under pan-tropical cooling. That the reconstructed regional 704 hydroclimate anomalies do not track the global temperature trends is perhaps unsurprising, given 705 that changes in tropical rainfall patterns would be expected to track the interhemispheric 706 temperature gradient and/or tropical temperature patterns more closely than global temperature 707 trends. Recent reconstructions of the interhemispheric temperature gradient (Neukom et al., 708 2014) do indeed mirror the hydroclimate trends from 1400 CE to present, however the 709 interhemispheric gradient diverges from the hydroclimate trends earlier in the millennium, e.g. 710 during the temperature gradient minimum around 1200-1300 CE (Fig. 3H). These discrepancies 711 may reflect uncertainty in the reconstructions, or they may indicate that the regional
manuscript submitted to Paleoceanography and Paleoclimatology
712 hydroclimate anomalies during the LIA were not simply related to changes in the inter- 713 hemispheric temperature gradient. Regardless, it is clear that more work needs to be done to 714 understand the mechanisms of the widespread tropical hydroclimate changes around 1400-1700 715 CE. A targeted assessment of the regional temperature changes in the tropics over the last 716 millennium would help better constrain the mechanisms.
717 4.4. Re-evaluation of Tropical Pacific Climate Changes Over the Last Millennium
718 We now re-examine a number of longstanding hypotheses about the nature of tropical 719 Pacific hydroclimate changes during the LIA. These hypotheses include: 1) a southward shift of 720 the mean annual position of the Pacific ITCZ (Nelson and Sachs, 2016; Richey and Sachs, 2016; 721 Sachs et al., 2009), 2) an equatorward contraction of the western Pacific and Asian-Australian 722 rainband (Denniston et al., 2016; Griffiths et al., 2016; Yan et al., 2015), and 3) a strengthened 723 Pacific Walker Circulation (Griffiths et al., 2016; Yan et al., 2011; Yan et al., 2015). We briefly 724 evaluate the evidence for each and highlight the areas where additional corroboration is needed.
725 4.4.1. A Large-scale Southward Shift of the Pacific ITCZ During the LIA?
726 A number of previous studies have postulated that a coordinated, large-scale southward 727 shift of the ITCZ occurred during the relatively cool epoch of the LIA, arguing that this 728 interpretation is consistent with dynamical theory that links the zonal mean ITCZ latitude to the 729 inter-hemispheric energy budget and temperature gradient (Broecker and Putnam, 2013; 730 Lechleitner et al., 2017; Nelson and Sachs, 2016; Richey and Sachs, 2016; Sachs et al., 2009). 731 This theory has been invoked in the inference of a southward shift of both the Pacific rain band 732 (Sachs et al., 2009) and Atlantic rain band (Haug et al., 2001) during the LIA.
733 We do not find broad support for a large-scale southward shift of the Pacific ITCZ during 734 the LIA. First, no clear evidence of a large-scale southward shift in the Pacific rain band is 735 observed when all available records from the central and eastern Pacific (where the Pacific ITCZ 736 is well defined) are compiled. While an abrupt transition to arid conditions around 1450 CE in 737 Washington Island is clear in our compilation CE (Fig. 3A; Fig. S3B), given its location under 738 the southern margin of the Pacific ITCZ in the present climate, dry conditions at that location 739 would be more likely to be associated with a northerly shift of the Pacific ITCZ (Fig. 1), as 740 occurs in association with an enhanced Pacific Walker Circulation during La Niña events. 741 Furthermore, there is also no synchronous shift in the eastern equatorial Pacific compilation 742 around this time (Fig. 3A), which lie just south of the southern extent of the eastern Pacific rain 743 band (Fig. 1), and thus would be expected to experience wetter conditions under a southward 744 shift of the Pacific ITCZ. Second, the theoretical framework for a southward shift of the Pacific 745 ITCZ during the LIA is lacking. Comprehensive climate models indicate that large-scale shifts in 746 the tropical rain belt occur only under strong hemispheric asymmetry in radiative forcing, and 747 even in those cases, the tropical Pacific rain band doesn’t shift uniformly (e.g. Atwood et al., 748 2020). Ultimately, robust conclusions regarding changes in the Pacific rainband during the LIA 749 will require more proxy data from the central and eastern Pacific, where the Pacific ITCZ is 750 well-defined. Modern process studies would also help refine the interpretation of some existing 751 records. For example, while all records from the Galápagos Islands indicate the pronounced 752 drying event around 800-1000 CE, the records show less consistency in the latter part of the last 753 millennium (Fig. 3A) with the highland lake records (Atwood and Sachs, 2014; Conroy et al.,
manuscript submitted to Paleoceanography and Paleoclimatology
754 2008) agreeing with one another but showing distinct behavior from the coastal lake records 755 (Nelson and Sachs, 2016). Ongoing efforts to develop additional hydroclimate reconstructions 756 from the Galápagos Islands and calibrate proxy data against modern instrumental data should 757 yield important insights in this data-sparse region of the tropical Pacific (e.g. Martin et al., 2018).
758 In summary, the evidence does not support a large-scale southward shift of the Pacific 759 rain band during the LIA. As additional paleoclimate records are developed from the central and 760 eastern tropical Pacific, all mechanisms of decadal rainfall variations in this region should be 761 considered in the interpretation of the records, including changes in the width, intensity, and 762 seasonal cycle of the Pacific rain band (Bartos et al., 2018; Byrne et al., 2018; Singarayer et al., 763 2017; Wodzicki and Rapp, 2016), as well as changes in the variability and spatial characteristics 764 of the El Niño/Southern Oscillation.
765 4.4.2. Strengthened Pacific Walker Circulation During the LIA?
766 Other studies have compiled various hydroclimate records from the tropical Pacific and 767 surrounding regions to infer a strengthened Pacific Walker Circulation during the LIA relative to 768 the MCA (Griffiths et al., 2016; Yan et al., 2011; Yan et al., 2015). The marine records from the 769 Maritime Continent indicating isotopically depleted (fresher) conditions during the LIA 770 generally support this interpretation (Fig. 3G), as does the Washington Island record in the 771 central tropical Pacific, which indicates an abrupt shift to drier conditions around 1450 CE (Fig. 772 3A; Fig. S3B). However, there is no evidence of drier conditions in the eastern equatorial Pacific 773 in association with a strengthened Pacific Walker Circulation during the LIA (Fig. 3A). This 774 apparent discrepancy may be explained by the fact that the Galápagos Islands are situated in the 775 already dry zone of the Pacific cold tongue and thus may experience little change in average 776 rainfall under an enhanced Walker Circulation. The available data thus supports LIA freshening 777 in parts of the Maritime Continent, with a possible linkage to central Pacific drying (based on a 778 single record), but with no clear connection to the eastern Pacific. We therefore conclude that 779 until further paleoclimate records are developed from the tropical Pacific, inferences of a 780 strengthened Pacific Walker Circulation during the LIA remain speculative. Additional proxy 781 sites near the Pacific rainbands could be especially helpful in this regard, as the rainbands would 782 be expected to expand poleward (equatorward) under an enhanced (weakened) Pacific Walker 783 Circulation as occurs in association with the El Niño/Southern Oscillation (Wallace et al., 1998).
784 4.4.3. Contraction of the Monsoonal Asian-Australian Rain Belt During the LIA?
785 Another hypothesized large-scale reorganization in tropical climate during the LIA has 786 been proposed based on proxy data from the western edge of the Pacific basin. Various 787 compilations of records from the Maritime Continent, Australia, and East Asia have been 788 interpreted as an equatorward contraction of rainfall at various periods within the LIA (ranging 789 from 1400-1640 CE (Denniston et al., 2016), 1400-1850 CE (Yan et al., 2015), and 1500-1900 790 CE (Griffiths et al., 2016)). Our findings of drying in the Asian summer monsoon region and 791 wetting in the Maritime Continent near the equator generally align with these prior studies, 792 though there are some aspects of this interpretation that remain tenuous due to limited data 793 availability. In particular, support for an equatorward contraction and/or weakening of the 794 Australian monsoon during the LIA presented in Denniston et al. (2016) and Griffiths et al. 795 (2016) rely on a single record from the Australian monsoon region (Denniston et al., 2016). Our
manuscript submitted to Paleoceanography and Paleoclimatology
796 analyses indicate that this record is consistent with a modest weakening of the Australian 797 monsoon from 1400-1600 CE followed by a pronounced intensification to the modern era, 798 though notably the LIA conditions are not anomalous with respect to background conditions of 799 the Common Era (Fig. S3C). Although eight records from the southern edge of the Australian 800 Summer Monsoon were presented in the analysis of Yan et al. (2015), only three of these records 801 spanned the last millennium, and of these, only one record (Denniston et al., 2016) was publicly 802 available and thus used in our analysis. In addition, in the analyses presented in Yan et al. 803 (2015), the reference period varied between the MCA and the last 150 years for different records 804 in their compilation, depending on record length. We see large differences between the MCA and 805 the last 200 years in many records in our analysis, and thus we strongly advocate for a 806 standardized treatment across all proxy records so that the trends over the last few hundred years 807 are separable from the changes between the MCA and LIA.
808 To robustly test the hypothesis of a coordinated contraction of the monsoonal Asian- 809 Australian rain belt over the last millennium, additional records from the Australian monsoon 810 region are needed. In addition, this hypothesis would benefit from a plausible set of physical 811 mechanisms that could explain a coordinated response of the two highly seasonal and out-of- 812 phase monsoonal rainfall regimes in Asia and Australia.
813 4.5. Insights from Data-Model Comparisons
814 Our primary objective in comparing the hydroclimate reconstructions to last millennium 815 climate model simulations is to shed light on the possible mechanisms of the reconstructed 816 hydroclimate changes. There are several ways in which the models lend important insight to the 817 reconstructions. First, the fact that both proxies and models indicate drier conditions over South 818 and East Asia during the second half of the last millennium provides evidence that the inferred 819 weakening of the Asian monsoons during the LIA may have been forced in part by volcanic 820 aerosols and/or orbital changes. However, the data-model mismatch in the timing of the drying 821 trends over Asia (i.e. aridity peaks around 1600 CE in the reconstructions while the drying trend 822 extends to at least 1800 CE in the models) indicates that there are clear deficiencies in the 823 models’ ability to capture this important feature of tropical hydroclimate change over the last 824 millennium.
825 South America is another region where the models potentially provide new insight into 826 the reconstructions. At first glance, tropical South America is seemingly characterized by 827 pronounced data/model disagreement, as the simulated precipitation and the reconstructed 828 hydroclimate changes are both robust but differ from one another in sign at the location of the 829 proxy sites (Fig. 4). The reconstructions show a distinct dipole pattern between the southwestern 830 and northeastern sites between the LIA and MCA, which past studies have interpreted as 831 reflecting increased upstream rainout under enhanced SASM activity during the LIA in tandem 832 with subsidence and decreased summer precipitation over northeastern Brazil (Lenters and Cook, 833 1997; Novello et al., 2012; Vuille et al., 2012). In contrast, the models do not simulate an 834 enhanced SASM during the LIA, however they do simulate increased rainfall in the monsoon 835 entrance region in the far northeastern corner of South America, that is driven by orbital and 836 volcanic forcing. The simulations thus provide an alternative interpretation of the proxy data: 837 that the observed LIA isotopic depletion in the central Andes and southern Amazon records 838 reflects increased upstream rainout in the monsoon entrance region in northeastern South
manuscript submitted to Paleoceanography and Paleoclimatology
839 America (instead of enhanced SASM activity in the Amazon basin). If this was indeed the case, 840 however, the region of enhanced precipitation must have been confined to a narrow band near 841 the equator, given the LIA drying (and/or isotopic enrichment) signal in both the northern and 842 Nordeste records.
843 There is also potential data-model agreement in the large-scale hydroclimate dynamics in 844 the tropical Atlantic Ocean over the last millennium, as a number of models simulate a 845 contraction of the Atlantic ITCZ along one or both edges of the Atlantic rain band (Fig. 4; Fig. 846 S4), while the reconstructions indicate a progressive shift to drier conditions in both the 847 northward edge (Cariaco Basin and Columbian Andes) and southward edge (Nordeste Brazil) of 848 the Atlantic ITCZ (Fig. 3D). To better assess the plausibility of these mechanisms in driving the 849 observed hydroclimate signals in tropical South America, simulations with water isotope tracers 850 should be included in future data-model comparison studies.
851 Finally, the models also potentially provide insight into the tropical Pacific climate 852 reconstructions. The tropical Pacific proxy records are most closely aligned with individual 853 CESM ensemble members that feature unforced centennial-scale changes in the strength of the 854 Pacific Walker Circulation (Fig. S8) and are also aligned with the CESM ensemble mean 855 response that is characterized by a weak equatorial Pacific drying trend and wetting in the 856 Maritime Continent west of Papua New Guinea (Fig. 4C,D). Given the larger magnitude of 857 internal variability relative to the forced response in CESM, and the fact that the pattern of 858 internal variability projects onto the forced response in the regions of the tropical Pacific proxy 859 data, these results suggest that, to the extent that CESM is representative of the real world, the 860 observed tropical Pacific hydroclimate changes during the LIA are likely due to natural 861 variations in the strength of the Walker Circulation. However, because the magnitude of tropical 862 Pacific climate variability varies widely across models (and is largest in CESM where the 863 variance of Niño 3 SST anomalies is around 150% of that observed over the last several 864 decades), we caution that the model uncertainty precludes firm confidence in the interpretation of 865 an unforced strengthening of the Pacific Walker Circulation during the LIA.
866 Outside of these potential areas of data-model agreement, one unresolved aspect of the 867 reconstructed hydroclimate changes that is not reproduced by the models is the timing of 868 widespread LIA hydroclimate changes in Asia, South America, and the Maritime Continent from 869 1400-1700 CE. There are several possible reasons for this data-model mismatch in the timing of 870 the LIA hydroclimate changes: (1) the reconstructed hydroclimate changes were unforced, (2) 871 the observed changes were forced but model uncertainty renders an unrealistic model response, 872 (3) the observed changes were forced but uncertainty in the forcing renders an unrealistic model 873 response, and/or (4) the proxy records, which are predominately derived from water isotope- 874 based archives, reflect something other than changes in regional mean annual precipitation 875 amount. We discuss each of these possibilities in more detail below.
876 With regard to (1), it is entirely possible that the reconstructed hydroclimate changes 877 from 1400-1700 CE were produced by natural variability, although the broadly synchronous 878 nature of the hydroclimate anomalies and hemispheric-scale temperature changes around this 879 time is somewhat suggestive of a forced response in the climate system. However, with regards 880 to (2), it is also expected that coupled ocean-atmosphere general circulation models would have 881 difficulty reproducing the regional tropical hydroclimate response to last millennium forcings,
manuscript submitted to Paleoceanography and Paleoclimatology
882 given the highly parameterized nature of convection and cloud physics, and the limited ability of 883 CMIP5 models to reproduce observed regional hydroclimate dynamics associated with major 884 hydroclimate phenomena such as the Asian and North American monsoons (Flato, 2013; PAGES 885 Hydro2k Consortium, 2017). Future data-model comparisons of the last millennium should 886 address the ability of the models to simulate the first-order features of the key hydroclimate 887 phenomena of interest. It has also been suggested that coupled climate models may 888 underestimate the strength and persistence of the hydroclimate response to volcanic forcing 889 (Tejedora et al., 2021), which if true, would severely limit the ability of such models to 890 realistically capture the long-term hydroclimate responses to volcanic forcing over the last 891 millennium. Finally, paleoclimate simulations do not typically include processes such as 892 dynamic vegetation (which require ocean–atmosphere–vegetation general circulation models), 893 although it is known that land surface processes are integral to the cycling of water between the 894 land and atmosphere. Indeed, vegetation changes have been shown to reduce model-data 895 discrepancies in simulations of the mid-Holocene (Braconnot et al., 2012).
896 With regards to (3), data-model discrepancies may result from errors in the radiative 897 forcings and feedbacks. There are substantial uncertainties in the radiative forcings over the last 898 millennium. For example, the radiative forcing during the LIA differs by more than a factor of 899 two across the CMIP5/PMIP3 models due to differences in volcanic forcing (associated with 900 both the aerosol dataset used and the implementation method of the volcanic aerosols; see 901 Methods) (Atwood et al., 2016). These differences contribute to a factor of two spread in the 902 amplitude of global cooling during the LIA across these models. These differences in volcanic 903 forcing likely lead to differing tropical rainfall responses in the models, since tropical rainfall 904 patterns are known to be sensitive to the timing, magnitude, and spatial footprint of volcanic 905 eruptions (e.g. Colose et al., 2016; PAGES Hydro2k Consortium, 2017). Furthermore, because 906 the volcanic aerosol loadings have changed (in both their magnitude and hemispheric 907 distributions) in recently updated last millennium reconstructions (Fig. S1; Sigl et al., 2014; Sigl 908 et al., 2015; Toohey and Sigl, 2017), the next iteration of model simulations may be expected to 909 differ in their representation of volcanically-forced tropical precipitation changes over the last 910 millennium. In addition to the uncertainties in radiative forcing, there are also uncertainties in the 911 strength and pattern of climate feedbacks among models. For example, the total effective 912 feedback parameter during the LIA differs by a factor of two across the CMIP5/PMIP3 models 913 (ranging from -1.0 to -2.0 W/m2/K), largely due to differences in the shortwave cloud feedback, 914 the lapse rate feedback, and the surface albedo feedback (Atwood et al., 2016). The degree of 915 interhemispheric asymmetry in the feedback strengths also varies across models (Fig. S13). 916 Errors in the interhemispheric asymmetry of the forcings and feedbacks over the last millennium 917 would be expected to lead to erroneous interhemispheric temperature asymmetry in the models. 918 Indeed, the models, which reasonably reproduce the long-term global cooling trends in 919 reconstructions, tend to substantially underestimate changes in the interhemispheric temperature 920 asymmetry over the last millennium (Neukom et al., 2014). Given the evidence that the tropical 921 hydroclimate anomalies may have been linked to the interhemispheric temperature gradient 922 during the LIA, future data-model comparison efforts should target simulations that reasonably 923 capture the interhemispheric temperature asymmetry during the LIA.
924 Finally, with regards to (4), it is likely that the proxy records used in this analysis, many 925 of which are derived from water isotope-based archives, do not solely reflect changes in regional 926 mean annual precipitation amount. The stable isotopic composition of meteoric water integrates
manuscript submitted to Paleoceanography and Paleoclimatology
927 information about many other aspects of the water cycle over space and time, including the 928 seasonality of precipitation, sources of moisture, advective transport, evaporation, and water 929 recycling over land. It is clear that more work needs to be done to understand the mechanisms of 930 the widespread tropical hydroclimate changes around 1400-1700 CE. Future work should 931 therefore employ climate models that simulate water isotope tracers to provide more quantitative 932 comparisons between proxy reconstructions and model simulations of the last millennium.
933
934 5. Conclusions
935 We review the evidence for large-scale tropical hydroclimate changes over the Common 936 Era based on a compilation of tropical hydroclimate records and assess the consistency between 937 the reconstructed hydroclimate changes and those simulated by forced transient model 938 simulations of the last millennium. A hierarchal cluster analysis indicates that 41% of the records 939 distributed throughout the tropics demonstrate a millennial-scale isotopic depletion/wetting trend 940 beginning around 1000 CE and reversing around 1800 CE that generally tracks global 941 temperature trends over the last millennium, which we speculate is associated with the influence 942 of decreased tropical sea surface temperatures on the isotopic composition of water vapor.
943 Our synthesis of the proxy records reveals a number of regionally coherent hydroclimate 944 patterns. During the onset of the MCA around 800-1000 CE, records from the eastern equatorial 945 Pacific and Mesoamerica indicate a pronounced drying event, relative to background conditions 946 of the Common Era. The second half of the last millennium is characterized by a set of 947 pronounced and synchronous hydroclimate changes across the tropics from 1400-1700 CE, 948 including drier (isotopically enriched) conditions in South and East Asia, fresher (isotopically 949 depleted) surface ocean conditions in the Maritime Continent, and a dipole pattern in tropical 950 South America, with wet (isotopically depleted) conditions in the southwest and dry (isotopically 951 enriched) conditions in northern South America and Nordeste Brazil. While previous studies 952 have linked these hydroclimate changes to global or hemispheric-scale cooling patterns, we find 953 notable dissimilarities between the regional hydroclimate changes and global- and hemispheric- 954 scale temperature reconstructions, indicating that more work needs to be done to understand the 955 mechanisms of the widespread tropical hydroclimate changes during the LIA.
956 Considering previous interpretations of tropical Pacific climate change during the LIA, 957 we do not find support for a large-scale southward shift of the Pacific ITCZ during the LIA. We 958 find that the available data generally supports a strengthened the Pacific Walker Circulation 959 during the LIA, with freshening in parts of the Maritime Continent, and a possible linkage to 960 central Pacific drying (based on a single lake sediment record), but with no clear connection to 961 the eastern Pacific. We therefore conclude that until further paleoclimate records are developed 962 from the tropical Pacific, inferences of a strengthened Pacific Walker Circulation during the LIA 963 remain speculative. Regarding the evidence for an equatorward contraction of the monsoonal 964 Asian-Australian rain belt during the LIA, we caution that more work needs to be done to 965 characterize the hydroclimate changes in the Australian monsoon region.
966 Lastly, transient climate model simulations of the last millennium provide several 967 important insights to the interpretation of the reconstructions: (1) the observed drying and/or
manuscript submitted to Paleoceanography and Paleoclimatology
968 isotopic enrichment in the South and East Asian monsoon region during the LIA may reflect 969 weakened monsoon activity driven in part by volcanic and orbital forcing, (2) the observed LIA 970 isotopic depletion in the central Andes and southern Amazon may reflect upstream rainout in the 971 monsoon entrance region (rather than a strengthening of the South American summer monsoon) 972 driven by orbital and volcanic forcing, (3) the observed coordinated drying in northern South 973 America and Nordeste Brazil over the last millennium may reflect a contraction of the Atlantic 974 ITCZ, and (4) isotopically depleted conditions in the Maritime Continent during the LIA are 975 consistent with an unforced strengthening of the Pacific Walker Circulation, though the large 976 model uncertainty in the tropical Pacific precludes confidence in this interpretation. One aspect 977 of the reconstructed hydroclimate changes that remains unresolved is the timing of the LIA 978 hydroclimate changes in Asia, South America, and the Maritime Continent from 1400-1700 CE, 979 followed by a reversal of the LIA conditions to the present era. To better assess the mechanisms 980 of the reconstructed LIA hydroclimate changes, future work should employ a suite of climate 981 model simulations with water isotope tracers and realistic tropical climate variability to improve 982 comparisons between proxy reconstructions and model simulations of the last millennium.
983
984 Acknowledgments 985 The authors gratefully acknowledge all of the researchers who made their paleoclimate 986 data available for public access and who participated in the CMIP5/PMIP3 and CESM LME 987 modeling efforts and made the model output publicly available. Computing resources 988 (doi:10.5065/D6RX99HX) were provided by the Climate Simulation Laboratory at NCAR’s 989 Computational and Information Systems Laboratory, sponsored by the National Science 990 Foundation and other agencies. The output from the CMIP5/PMIP3 and CESM LME simulations 991 was obtained from the Earth System Grid (earthsystemgrid.org). Support for this work was 992 provided to ARA by the U.S. National Science Foundation under award AGS-1949605 and by 993 the NSF Graduate Research Fellowship Program (DGE-1256082) and the NOAA Climate and 994 Global Change Postdoctoral Program (under fellowships to ARA). Support was provided to EW 995 and DMWF by NSF Awards AGS-0846641, AGS-0936059, AGS-1359464, and AGS-1665247. 996 Support was provided to D.S.B by the Tamaki Foundation, and to J.P.S. by NSF under Grant 997 Nos. 0823503 and 1502417. We thank K. Shell for providing the CAM3 radiative kernels. The 998 proxy data used in this study were uploaded to Figshare 999 (https://figshare.com/articles/Hydroclimate_proxy_records_zip/12085092)and published with 1000 doi: 10.6084/m9.figshare.12085092. The records were originally downloaded from the NOAA 1001 Paleoclimatology and PANGAEA databases (see dataset URLs in Table 1), or obtained directly 1002 from the authors (see references in Table 1).
manuscript submitted to Paleoceanography and Paleoclimatology
1003
1004 Figure Captions
1005 Figure 1. Observed mean annual precipitation from GPCP v.2.2 for the period 1979-2012 in 1006 mm/day. The location of the proxy records along with the archive type is denoted by the symbols 1007 (circle = speleothem, triangle = lake, square = marine sediment, diamond = ice) and the color 1008 indicates whether the record is based on water isotopes (green) or not (orange). The numbered 1009 boxes enclose the proxy records that were used in the regional composites.
1010 Figure 2. A) Heat map and dendrogram from cluster analysis of standardized tropical 1011 hydroclimate records used in this study. B,D,F) Average (black line) and 95% bootstrap 1012 confidence intervals with 1,000 iterations (shading) of all records in each cluster. Proxy records 1013 were converted into Z-scores and averaged over 100-year bins during the period 350-1950 CE. 1014 For all records, higher Z-scores indicate wetter conditions, based on interpretation in the original 1015 study from Table 1. C,E,G) Location of proxy records, where the symbols correspond to the type 1016 of proxy archive indicated in Fig. 1. Symbols with a yellow astrisk indicate that the record is not 1017 based on water isotopes.
1018 Figure 3. Figure 3. (A-G) Regional compilations of the tropical hydroclimate records, converted 1019 into Z-scores and averaged over 100-year bins. For all records, higher Z-scores indicate wetter 1020 conditions, based on interpretation in the original study (from Table 1). The average and standard 1021 deviation of all records in each region is depicted by the black line and shading, respectively. 1022 The maps indicate the location of the records, where the symbol denotes the type of proxy 1023 archive (circle = speleothem, triangle = lake, square = marine sediment, diamond = ice) and the 1024 color indicates whether the record is based on water isotopes (green) or not (orange with 1025 asterisk). The legend indicates the relevant reference for each record (see Table 1). Overlapping 1026 symbols have been slightly offset for visual clarity. (H) The reconstructed global mean 1027 temperature anomaly (relative to 1000-1200 CE) from PAGES2k Constortium (2019) is shown 1028 in black, where the solid line is the 31-yr low-pass filtered ensemble median, and the shading is 1029 the 2.5th and 97.5th percentiles of all reconstruction members; the Neukom et al. (2014) 1030 reconstructed interhemispheric (NH-SH) temperature difference (30-year loess-filtered ensemble 1031 mean). The averages over 100-yr bins are also shown for both temperature reconstructions.
1032 Figure 4. Multi-model mean change in precipitation during the LIA (1450-1850 CE) relative to 1033 the MCA (950-1200 CE). (A-B) ∆P in the six PMIP3/CMIP5 LM simulations in units of (A) 1034 mm/day and (B) percent, relative to mean annual precipitation during the MCA. (C-D) ∆P in the 1035 13 all-forcing CESM LME simulations in units of (C) mm/day and (D) percent, relative to mean 1036 annual precipitation during the MCA. Unfilled contours are mean annual precipitation during the 1037 MCA (with contours at 4 and 8 mm/day). Stippling indicates where ≥83% of the models agree 1038 on the sign of the change (5/6 models in panels A,B and 11/13 simulations in panels C,D). The 1039 markers indicate the paleohydroclimate records in Table 1, where the type of symbol denotes the 1040 proxy archive (circle = speleothem, triangle = lake sediment, square = marine sediment, diamond 1041 = ice), the size of the marker scales with log of the difference in Z-scores, and the color 1042 represents the sign of the inferred precipitation change (blue = positive/wetter, red = 1043 negative/drier, grey = no significant change) during the LIA relative to the MCA, based on an 1044 unpaired two-sample Student’s t-test.
manuscript submitted to Paleoceanography and Paleoclimatology
1045 Figure 5. Multi-model mean precipitation change (colors) in the CESM individual forcing 1046 simulations during the LIA (1450-1850 CE) relative to the MCA (950-1200 CE). Unfilled 1047 contours indicate the mean annual precipitation during the MCA (contours are 4 and 8 mm/day). 1048 Stippling indicates where 80% of the models agree on the sign of the change.
1049 Figure 6. Normalized EOF 1 loadings and PC 1 of (A) the proxy Z-scores, (B) simulated tropical 1050 precipitation in the CESM ensemble mean, (C) simulated tropical precipitation in the 1051 CMIP5/PMIP3 multi-model mean, and (D-H) simulated tropical precipitation in the first 1052 ensemble member of the CESM LME and in each of the CMIP5/PMIP3 last millennium 1053 simulations. The proxy and model data were binned into 100-yr windows prior to EOF analysis. 1054 Higher values of the loadings indicate wetter (or more isotopically depleted) conditions.
1055 Figure 7. Change in zonal mean interhemispheric tropical precipitation asymmetry between 1056 20°S:20°N (∆PNH-SH) in the LIA (1450-1850 CE) relative to the MCA (950-1200 CE) versus the 1057 change in cross-equatorial atmospheric energy transport (∆AHTΦ=0) in the PMIP3/CMIP5 last 1058 millennium simulations. Positive values of ∆AHTΦ=0 represent anomalous northward heat 1059 transport. ∆PNH-SH averaged over the 13-member CESM last millennium ensemble is shown by 1060 the dashed horizontal line (the radiation data needed to calculate ∆AHTΦ=0 was not obtained for 1061 the CESM LME).
1062 1063
manuscript submitted to Paleoceanography and Paleoclimatology
1064 References 1065 1066 Anderson, D.M., Overpeck, J.T., Gupta, A.K., 2002. Increase in the Asian southwest monsoon during the past four 1067 centuries. Science 297, 596-599. 1068 As-syakur, A.R., Tanaka, T., Osawa, T., Mahendra, M.S., 2013. Indonesian rainfall variability observation using 1069 TRMM multi-satellite data. Int. J. Remote Sens. 34, 7723-7738. 1070 Atwood, A.R., Donohoe, A., Battisti, D.S., Liu, X., Pausata, F.S.R., 2020. Robust longitudinally-variable responses 1071 of the ITCZ to a myriad of climate forcings. Geophys. Res. Lett. n/a, e2020GL088833. 1072 Atwood, A.R., Sachs, J.P., 2014. Separating ITCZ- and ENSO-related rainfall changes in the Galápagos over the last 1073 3 kyr using D/H ratios of multiple lipid biomarkers. Earth Planet. Sci. Lett. 404, 408-419. 1074 Atwood, A.R., Wu, E., Frierson, D.M.W., Battisti, D.S., Sachs, J.P., 2016. Quantifying Climate Forcings and 1075 Feedbacks over the Last Millennium in the CMIP5–PMIP3 Models. J. Clim. 29, 1161-1178. 1076 Bartos, E.A., Rapp, A.D., Wodzicki, K.R., 2018. Increasing Frequency of Midtropospheric Dry Layers in the Pacific 1077 Intertropical Convergence Zone. Geophys. Res. Lett. 45, 13,523-513,529. 1078 Bhattacharya, T., Chiang, J.C.H., Cheng, W., 2017. Ocean-atmosphere dynamics linked to 800-1050 CE drying in 1079 mesoamerica. Quat. Sci. Rev. 169, 263-277. 1080 Biasutti, M., Voigt, A., Boos, W.R., Braconnot, P., Hargreaves, J.C., Harrison, S.P., Kang, S.M., Mapes, B.E., 1081 Scheff, J., Schumacher, C., Sobel, A.H., Xie, S.-P., 2018. Global energetics and local physics as drivers of past, 1082 present and future monsoons. Nat. Geosci. 11, 392-400. 1083 Bird, B.W., Abbott, M.B., Vuille, M., Rodbell, D.T., Stansell, N.D., Rosenmeier, M.F., 2011. A 2,300-year-long 1084 annually resolved record of the South American summer monsoon from the Peruvian Andes. Proc. Nat. Acad. Sci. 1085 U.S.A. 108, 8583-8588. 1086 Braconnot, P., Harrison, S.P., Kageyama, M., Bartlein, P.J., Masson-Delmotte, V., Abe-Ouchi, A., Otto-Bliesner, 1087 B., Zhao, Y., 2012. Evaluation of climate models using palaeoclimatic data. Nature Climate Change 2, 417-424. 1088 Broecker, W.S., Putnam, A.E., 2013. Hydrologic impacts of past shifts of Earth’s thermal equator offer insight into 1089 those to be produced by fossil fuel CO2. Proceedings of the National Academy of Sciences 110, 16710-16715. 1090 Buckley, B.M., Anchukaitis, K.J., Penny, D., Fletcher, R., Cook, E.R., Sano, M., Le, C.N., Wichienkeeo, A., Ton, 1091 T.M., Truong, M.H., 2010. Climate as a contributing factor in the demise of Angkor, Cambodia. Proc. Nat. Acad. 1092 Sci. U.S.A. 107, 6748-6752. 1093 Byrne, M.P., Pendergrass, A.G., Rapp, A.D., Wodzicki, K.R., 2018. Response of the Intertropical Convergence 1094 Zone to Climate Change: Location, Width, and Strength. Current Climate Change Reports 4, 355-370. 1095 Carolin, S.A., Cobb, K.M., Adkins, J.F., Clark, B., Conroy, J.L., Lejau, S., Malang, J., Tuen, A.A., 2013. Varied 1096 Response of Western Pacific Hydrology to Climate Forcings over the Last Glacial Period. Science 340, 1564. 1097 Chiang, J.C.H., 2009. The Tropics in Paleoclimate. Annual Review of Earth and Planetary Sciences 37, 263-297. 1098 Chiang, J.C.H., Bitz, C.M., 2005. Influence of high latitude ice cover on the marine Intertropical Convergence Zone. 1099 Clim. Dyn. 25, 477-496. 1100 Chiang, J.C.H., Kushnir, Y., Giannini, A., 2002. Deconstructing Atlantic Intertropical Convergence Zone 1101 variability: Influence of the local cross-equatorial sea surface temperature gradient and remote forcing from the 1102 eastern equatorial Pacific. Journal of Geophysical Research: Atmospheres 107, ACL 3-1-ACL 3-19. 1103 Collins, M., Tett, S.F.B., Cooper, C., 2001. The internal climate variability of HadCM3, a version of the Hadley 1104 Centre coupled model withou flux adjustments. Clim. Dyn. 17, 61-81. 1105 Colose, C.M., LeGrande, A.N., Vuille, M., 2016. Hemispherically asymmetric volcanic forcing of tropical 1106 hydroclimate during the last millennium. Earth System Dynamics 7, 681-696. 1107 Conroy, J.L., Overpeck, J.T., Cole, J.E., Shanahan, T.M., Steinitz-Kannan, M., 2008. Holocene changes in eastern 1108 tropical Pacific climate inferred from a Galapagos lake sediment record. Quat. Sci. Rev. 27, 1166-1180. 1109 Consortium, P.k., 2013. Continental-scale temperature variability during the past two millennia. Nat. Geosci. 6, 339- 1110 346. 1111 Consortium, P.k., 2019. Consistent multidecadal variability in global temperature reconstructions and simulations 1112 over the Common Era. Nat. Geosci. 12, 643-649. 1113 Crucifix, M., 2006. Does the Last Glacial Maximum constrain climate sensitivity? Geophys. Res. Lett. 33, L18701. 1114 Cruz, F.W., Vuille, M., Burns, S.J., Wang, X., Cheng, H., Werner, M., Lawrence Edwards, R., Karmann, I., Auler, 1115 A.S., Nguyen, H., 2009. Orbitally driven east–west antiphasing of South American precipitation. Nat. Geosci. 2, 1116 210-214. 1117 Cunningham, L.K., Austin, W.E.N., Knudsen, K.L., Eiriksson, J., Scourse, J.D., Wanamaker, A.D., Butler, P.G., 1118 Cage, A.G., Richter, T., Husum, K., Hald, M., Andersson, C., Zorita, E., Linderholm, H.W., Gunnarson, B.E., Sicre,
manuscript submitted to Paleoceanography and Paleoclimatology
1119 M.A., Sejrup, H.P., Jiang, H., Wilson, R.J.S., 2013. Reconstructions of surface ocean conditions from the northeast 1120 Atlantic and Nordic seas during the last millennium. Holocene 23, 921-935. 1121 Dansgaard, W., 1964. Stable isotopes in precipitation. Tellus 16, 436-468. 1122 Davey, M.K., Brookshaw, A., Ineson, S., 2014. The probability of the impact of ENSO on precipitation and near- 1123 surface temperature. Climate Risk Management 1, 5-24. 1124 Denniston, R.F., Ummenhofer, C.C., Wanamaker, A.D., Lachniet, M.S., Villarini, G., Asmerom, Y., Polyak, V.J., 1125 Passaro, K.J., Cugley, J., Woods, D., Humphreys, W.F., 2016. Expansion and Contraction of the Indo-Pacific 1126 Tropical Rain Belt over the Last Three Millennia. Scientific Reports 6, 34485. 1127 Deplazes, G., Lückge, A., Peterson, L.C., Timmermann, A., Hamann, Y., Hughen, K.A., Röhl, U., Laj, C., Cane, 1128 M.A., Sigman, D.M., Haug, G.H., 2013. Links between tropical rainfall and North Atlantic climate during the last 1129 glacial period. Nat. Geosci. 6, 213-217. 1130 Donohoe, A., Marshall, J., Ferreira, D., McGee, D., 2013. The Relationship between ITCZ Location and Cross- 1131 Equatorial Atmospheric Heat Transport: From the Seasonal Cycle to the Last Glacial Maximum. J. Clim. 26, 3597- 1132 3618. 1133 Eisen, M.B., Spellman, P.T., Brown, P.O., Botstein, D., 1998. Cluster analysis and display of genome-wide 1134 expression patterns. Proceedings of the National Academy of Sciences 95, 14863. 1135 Emile-Geay, J., McKay, N.P., Kaufman, D.S., von Gunten, L., Wang, J., Anchukaitis, K.J., Abram, N.J., Addison, 1136 J.A., Curran, M.A.J., Evans, M.N., Henley, B.J., Hao, Z., Martrat, B., McGregor, H.V., Neukom, R., Pederson, G.T., 1137 Stenni, B., Thirumalai, K., Werner, J.P., Xu, C., Divine, D.V., Dixon, B.C., Gergis, J., Mundo, I.A., Nakatsuka, T., 1138 Phipps, S.J., Routson, C.C., Steig, E.J., Tierney, J.E., Tyler, J.J., Allen, K.J., Bertler, N.A.N., Björklund, J., Chase, 1139 B.M., Chen, M.-T., Cook, E., de Jong, R., DeLong, K.L., Dixon, D.A., Ekaykin, A.A., Ersek, V., Filipsson, H.L., 1140 Francus, P., Freund, M.B., Frezzotti, M., Gaire, N.P., Gajewski, K., Ge, Q., Goosse, H., Gornostaeva, A., Grosjean, 1141 M., Horiuchi, K., Hormes, A., Husum, K., Isaksson, E., Kandasamy, S., Kawamura, K., Kilbourne, K.H., Koç, N., 1142 Leduc, G., Linderholm, H.W., Lorrey, A.M., Mikhalenko, V., Mortyn, P.G., Motoyama, H., Moy, A.D., Mulvaney, 1143 R., Munz, P.M., Nash, D.J., Oerter, H., Opel, T., Orsi, A.J., Ovchinnikov, D.V., Porter, T.J., Roop, H.A., Saenger, 1144 C., Sano, M., Sauchyn, D., Saunders, K.M., Seidenkrantz, M.-S., Severi, M., Shao, X., Sicre, M.-A., Sigl, M., 1145 Sinclair, K., St. George, S., St. Jacques, J.-M., Thamban, M., Kuwar Thapa, U., Thomas, E.R., Turney, C., Uemura, 1146 R., Viau, A.E., Vladimirova, D.O., Wahl, E.R., White, J.W.C., Yu, Z., Zinke, J., Consortium, P.A.k., 2017. A global 1147 multiproxy database for temperature reconstructions of the Common Era. Scientific Data 4, 170088. 1148 Falster, G., Konecky, B.L., McKay, N., Atwood, A.R., Conroy, J.L., Fischer, M., Jones, M., Jonkers, L., Stevenson, 1149 S., Thompson, D.M., Tyler, J.J., 2019. New insights into spatial and temporal dynamics of the global water cycle 1150 from the Iso2k database, AGU Fall Meeting. 1151 Fernandez-Donado, L., Gonzalez-Rouco, J.F., Raible, C.C., Ammann, C.M., Barriopedro, D., Garcia-Bustamante, 1152 E., Jungclaus, J.H., Lorenz, S.J., Luterbacher, J., Phipps, S.J., Servonnat, J., Swingedouw, D., Tett, S.F.B., Wagner, 1153 S., Yiou, P., Zorita, E., 2013. Large-scale temperature response to external forcing in simulations and 1154 reconstructions of the last millennium. Climate of the Past 9, 393-421. 1155 Flato, G., J. Marotzke, B. Abiodun, P. Braconnot, S.C. Chou, W. Collins, P. Cox, F. Driouech, S. Emori, V. Eyring, 1156 C. Forest, P. Gleckler, E. Guilyardi, C. Jakob, V. Kattsov, C. Reason and M. Rummukainen, 2013. Evaluation of 1157 Climate Models. In: Climate Change 2013: The Physical Science Basis. Contribution of Working Group I to the 1158 Fifth Assess- ment Report of the Intergovernmental Panel on Climate Change [Stocker, T.F., D. Qin, G.-K. Plattner, 1159 M. Tignor, S.K. Allen, J. Boschung, A. Nauels, Y. Xia, V. Bex and P.M. Midgley (eds.)]. Cambridge University 1160 Press, Cambridge, United Kingdom and New York, NY, USA. 1161 Gent, P.R., Danabasoglu, G., Donner, L.J., Holland, M.M., Hunke, E.C., Jayne, S.R., Lawrence, D.M., Neale, R.B., 1162 Rasch, P.J., Vertenstein, M., Worley, P.H., Yang, Z.-L., Zhang, M., 2011. The Community Climate System Model 1163 Version 4. J. Clim. 24, 4973-4991. 1164 Griffiths, M.L., Kimbrough, A.K., Gagan, M.K., Drysdale, R.N., Cole, J.E., Johnson, K.R., Zhao, J.-X., Cook, B.I., 1165 Hellstrom, J.C., Hantoro, W.S., 2016. Western Pacific hydroclimate linked to global climate variability over the past 1166 two millennia. Nature Communications 7, 11719. 1167 group, P.k.-P., 2015. Continental-scale temperature variability in PMIP3 simulations and PAGES 2k regional 1168 temperature reconstructions over the past millennium. Clim. Past 11, 1673-1699. 1169 Haug, G.H., Hughen, K.A., Sigman, D.M., Peterson, L.C., Rohl, U., 2001. Southward migration of the intertropical 1170 convergence zone through the Holocene. Science 293, 1304-1308. 1171 Hegerl, G.C., Crowley, T.J., Allen, M., Hyde, W.T., Pollack, H.N., Smerdon, J., Zorita, E., 2007. Detection of 1172 human influence on a new, validated 1500-year temperature reconstruction. J. Clim. 20, 650-666. 1173 Higley, M.C., Conroy, J.L., Schmitt, S., 2018. Last millennium meridional shifts in hydroclimate in the central 1174 tropical Pacific. Paleoceanography and Paleoclimatology 33, 354-366.
manuscript submitted to Paleoceanography and Paleoclimatology
1175 Hodell, D.A., Brenner, M., Curtis, J.H., 2005. Terminal Classic drought in the northern Maya lowlands inferred 1176 from multiple sediment cores in Lake Chichancanab (Mexico). Quat. Sci. Rev. 24, 1413-1427. 1177 Hwang, Y.T., Frierson, D.M.W., Kang, S.M., 2013. Anthropogenic sulfate aerosol and the southward shift of 1178 tropical precipitation in the late 20th century. Geophys. Res. Lett. 40, 1-6. 1179 Kang, S.M., Held, I.M., Frierson, D.M.W., Zhao, M., 2008. The response of the ITCZ to extratropical thermal 1180 forcing: Idealized slab-ocean experiments with a GCM. J. Clim. 21, 3521-3532. 1181 Kanner, L.C., Burns, S.J., Cheng, H., Edwards, R.L., 2012. High-Latitude Forcing of the South American Summer 1182 Monsoon During the Last Glacial. Science 335, 570-573. 1183 Kaufman, D.S., Axford, Y.L., Henderson, A.C.G., McKay, N.P., Oswald, W.W., Saenger, C., Anderson, R.S., 1184 Bailey, H.L., Clegg, B., Gajewski, K., Hu, F.S., Jones, M.C., Massa, C., Routson, C.C., Werner, A., Wooller, M.J., 1185 Yu, Z., 2016. Holocene climate changes in eastern Beringia (NW North America) – A systematic review of multi- 1186 proxy evidence. Quat. Sci. Rev. 147, 312-339. 1187 Kaufman, D.S., Schneider, D.P., McKay, N.P., Ammann, C.M., Bradley, R.S., Briffa, K.R., Miller, G.H., Otto- 1188 Bliesner, B.L., Overpeck, J.T., Vinther, B.M., Arctic Lakes 2k Project Members, 2009. Recent Warming Reverses 1189 Long-Term Arctic Cooling. Science 325, 1236-1239. 1190 Kobashi, T., Kawamura, K., Severinghaus, J.P., Barnola, J.-M., Nakaegawa, T., Vinther, B.M., Johnsen, S.J., Box, 1191 J.E., 2011. High variability of Greenland surface temperature over the past 4000 years estimated from trapped air in 1192 an ice core. Geophys. Res. Lett. 38. 1193 Konecky, B.L., McKay, N.P., Churakova, O.V., Comas-Bru, L., Dassié, E.P., DeLong, K.L., Falster, G.M., Fischer, 1194 M.J., Jones, M.D., Jonkers, L., Kaufman, D.S., Leduc, G., Managave, S.R., Martrat, B., Opel, T., Orsi, A.J., Partin, 1195 J.W., Sayani, H.R., Thomas, E.K., Thompson, D.M., Tyler, J.J., Abram, N.J., Atwood, A.R., Conroy, J.L., Kern, Z., 1196 Porter, T.J., Stevenson, S.L., von Gunten, L., 2020. The Iso2k Database: A global compilation of paleo-δ18O and 1197 δ²H records to aid understanding of Common Era climate. Earth Syst. Sci. Data Discuss. 2020, 1-49. 1198 Landrum, L., Otto-Bliesner, B.L., Wahl, E.R., Conley, A., Lawrence, P.J., Rosenbloom, N., Teng, H., 2013. Last 1199 Millennium Climate and Its Variability in CCSM4. J. Clim. 26, 1085-1111. 1200 Lechleitner, F.A., Breitenbach, S.F.M., Rehfeld, K., Ridley, H.E., Asmerom, Y., Prufer, K.M., Marwan, N., 1201 Goswami, B., Kennett, D.J., Aquino, V.V., Polyak, V., Haug, G.H., Eglinton, T.I., Baldini, J.U.L., 2017. Tropical 1202 rainfall over the last two millennia: evidence for a low-latitude hydrologic seesaw. Scientific Reports 7. 1203 Lehner, F., Born, A., Raible, C.C., Stocker, T.F., 2013. Amplified Inception of European Little Ice Age by Sea Ice- 1204 Ocean-Atmosphere Feedbacks. J. Clim. 26, 7586-7602. 1205 Lenters, J.D., Cook, K.H., 1997. On the Origin of the Bolivian High and Related Circulation Features of the South 1206 American Climate. J. Atmos. Sci. 54, 656-677. 1207 Mann, M.E., Zhang, Z.H., Rutherford, S., Bradley, R.S., Hughes, M.K., Shindell, D., Ammann, C., Faluvegi, G., Ni, 1208 F.B., 2009. Global Signatures and Dynamical Origins of the Little Ice Age and Medieval Climate Anomaly. Science 1209 326, 1256-1260. 1210 Marcott, S.A., Shakun, J.D., Clark, P.U., Mix, A.C., 2013. A Reconstruction of Regional and Global Temperature 1211 for the Past 11,300 Years. Science 339, 1198-1201. 1212 Marsland, S.J., Haak, H., Jungclaus, J.H., Latif, M., Röske, F., 2003. The Max-Planck-Institute global ocean/sea ice 1213 model with orthogonal curvilinear coordinates. Ocean Modelling 5, 91-127. 1214 Martin, N.J., Conroy, J.L., Noone, D., Cobb, K.M., Konecky, B.L., Rea, S., 2018. Seasonal and ENSO Influences on 1215 the Stable Isotopic Composition of Galápagos Precipitation. Journal of Geophysical Research: Atmospheres 123, 1216 261-275. 1217 Masse, G., Rowland, S.J., Sicre, M.A., Jacob, J., Jansen, E., Belt, S.T., 2008. Abrupt climate changes for Iceland 1218 during the last millennium: Evidence from high resolution sea ice reconstructions. Earth Planet. Sci. Lett. 269, 564- 1219 568. 1220 Masson-Delmotte, V., M. Schulz, A. Abe-Ouchi, J. Beer, A. Ganopolski, J.F. González Rouco, E. Jansen, K. 1221 Lambeck, J. Luterbacher, T. Naish, T. Osborn, B. Otto-Bliesner, T. Quinn, R. Ramesh, M. Rojas, X. Shao and A. 1222 Timmermann, 2013. Information from Paleoclimate Archives. In: Climate Change 2013: The Physical Science 1223 Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate 1224 Change [Stocker, T.F., D. Qin, G.-K. Plattner, M. Tignor, S.K. Allen, J. Boschung, A. Nauels, Y. Xia, V. Bex and 1225 P.M. Midgley (eds.)] Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA. 1226 Moberg, A., Sonechkin, D.M., Holmgren, K., Datsenko, N.M., Karlen, W., 2005. Highly variable Northern 1227 Hemisphere temperatures reconstructed from low- and high-resolution proxy data. Nature 433, 613-617. 1228 Nelson, D.B., Sachs, J.P., 2016. Galápagos hydroclimate of the Common Era from paired microalgal and mangrove 1229 biomarker 2H/1H values. Proceedings of the National Academy of Sciences 113, 3476–3481.
manuscript submitted to Paleoceanography and Paleoclimatology
1230 Neukom, R., Gergis, J., Karoly, D.J., Wanner, H., Curran, M., Elbert, J., Gonzalez-Rouco, F., Linsley, B.K., Moy, 1231 A.D., Mundo, I., Raible, C.C., Steig, E.J., van Ommen, T., Vance, T., Villalba, R., Zinke, J., Frank, D., 2014. Inter- 1232 hemispheric temperature variability over the past millennium. Nature Climate Change 4, 362-367. 1233 Novello, V.F., Cruz, F.W., Karmann, I., Burns, S.J., Stirkis, N.M., Vuille, M., Cheng, H., Edwards, R.L., Santos, 1234 R.V., Frigo, E., Barreto, E.A.S., 2012. Multidecadal climate variability in Brazil's Nordeste during the last 3000 1235 years based on speleothem isotope records. Geophys. Res. Lett. 39. 1236 Orrison, R., Vuille, M., 2020. Dynamical Drivers of South American Monsoon Variability over the Last 1237 Millennium, AGU 2020 Fall Meeting. 1238 Otto-Bliesner, B.L., Brady, E.C., Fasullo, J., 2016. Climate Variability and Change since 850 CE: An Ensemble 1239 Approach with the Community Earth System Model Bull. Am. Meteorol. Soc. 97, 735-754. 1240 PAGES Hydro2k Consortium, 2017. Comparing proxy and model estimates of hydroclimate variability and change 1241 over the Common Era. Climate of the Past 13, 1851-1900. 1242 Phipps, S.J., Rotstayn, L.D., Gordon, H.B., Roberts, J.L., Hirst, A.C., Budd, W.F., 2012. The CSIRO Mk3L climate 1243 system model version 1.0-Part 2: Response to external forcings. Geoscientific Model Development 5, 649-682. 1244 Pope, V.D., Gallani, M.L., Rowntree, P.R., Stratton, R.A., 2000. The impact of new physical parametrizations in the 1245 Hadley Centre climate model: HadAM3. Clim. Dyn. 16, 123-146. 1246 Raddatz, T.J., Reick, C.H., Knorr, W., Kattge, J., Roeckner, E., Schnur, R., Schnitzler, K.-G., Wetzel, P., Jungclaus, 1247 J., 2007. Will the tropical land biosphere dominate the climate–carbon cycle feedback during the twenty-first 1248 century? Clim. Dyn. 29, 565-574. 1249 Richey, J.N., Sachs, J.P., 2016. Precipitation changes in the western tropical Pacific over the past millennium. 1250 Geology 44, 671-674. 1251 Risi, C., Bony, S., Vimeux, F., 2008. Influence of convective processes on the isotopic composition (delta O-18 and 1252 delta D) of precipitation and water vapor in the tropics: 2. Physical interpretation of the amount effect. Journal of 1253 Geophysical Research-Atmospheres 113. 1254 Risi, C., Bony, S., Vimeux, F., Jouzel, J., 2010. Water-stable isotopes in the LMDZ4 general circulation model: 1255 Model evaluation for present-day and past climates and applications to climatic interpretations of tropical isotopic 1256 records. Journal of Geophysical Research: Atmospheres 115. 1257 Roberts, W.H.G., Valdes, P.J., Singarayer, J.S., 2017. Can energy fluxes be used to interpret glacial/interglacial 1258 precipitation changes in the tropics? Geophys. Res. Lett. 44, 6373-6382. 1259 Rotstayn, L.D., Jeffrey, S.J., Collier, M.A., Dravitzki, S.M., Hirst, A.C., Syktus, J.I., Wong, K.K., 2012. Aerosol- 1260 and greenhouse gas-induced changes in summer rainfall and circulation in the Australasian region: a study using 1261 single-forcing climate simulations. Atmos. Chem. Phys. 12, 6377-6404. 1262 Sachs, J.P., Sachse, D., Smittenberg, R.H., Zhang, Z.H., Battisti, D.S., Golubic, S., 2009. Southward movement of 1263 the Pacific intertropical convergence zone AD 1400-1850. Nat. Geosci. 2, 519-525. 1264 Schmidt, G.A., Jungclaus, J.H., Ammann, C.M., Bard, E., Braconnot, P., Crowley, T.J., Delaygue, G., Joos, F., 1265 Krivova, N.A., Muscheler, R., Otto-Bliesner, B.L., Pongratz, J., Shindell, D.T., Solanki, S.K., Steinhilber, F., Vieira, 1266 L.E.A., 2011. Climate forcing reconstructions for use in PMIP simulations of the last millennium (v1.0). 1267 Geoscientific Model Development 4, 33-45. 1268 Schmidt, G.A., Ruedy, R., Hansen, J.E., Aleinov, I., Bell, N., Bauer, M., Bauer, S., Cairns, B., Canuto, V., Cheng, 1269 Y., Del Genio, A., Faluvegi, G., Friend, A.D., Hall, T.M., Hu, Y., Kelley, M., Kiang, N.Y., Koch, D., Lacis, A.A., 1270 Lerner, J., Lo, K.K., Miller, R.L., Nazarenko, L., Oinas, V., Perlwitz, J., Perlwitz, J., Rind, D., Romanou, A., 1271 Russell, G.L., Sato, M., Shindell, D.T., Stone, P.H., Sun, S., Tausnev, N., Thresher, D., Yao, M.-S., 2006. Present- 1272 day atmospheric simulations using GISS ModelE: comparison to in situ, satellite, and reanalysis data. J. Clim. 19, 1273 153-192. 1274 Schurer, A.P., Hegerl, G.C., Mann, M.E., Tett, S.F.B., Phipps, S.J., 2013. Separating Forced from Chaotic Climate 1275 Variability over the Past Millennium. J. Clim. 26, 6954-6973. 1276 Schurer, A.P., Tett, S.F.B., Hegerl, G.C., 2014. Small influence of solar variability on climate over the past 1277 millennium. Nat. Geosci. 7, 104-108. 1278 Shuman, B.N., Routson, C., McKay, N., Fritz, S., Kaufman, D., Kirby, M.E., Nolan, C., Pederson, G.T., St-Jacques, 1279 J.M., 2018. Placing the Common Era in a Holocene context: millennial to centennial patterns and trends in the 1280 hydroclimate of North America over the past 2000 years. Clim. Past 14, 665-686. 1281 Sigl, M., McConnell, J.R., Toohey, M., Curran, M., Das, S.B., Edwards, R., Isaksson, E., Kawamura, K., Kipfstuhl, 1282 S., Krueger, K., Layman, L., Maselli, O.J., Motizuki, Y., Motoyama, H., Pasteris, D.R., Severi, M., 2014. Insights 1283 from Antarctica on volcanic forcing during the Common Era. Nature Climate Change 4, 693-697. 1284 Sigl, M., Winstrup, M., McConnell, J.R., Welten, K.C., Plunkett, G., Ludlow, F., Büntgen, U., Caffee, M., 1285 Chellman, N., Dahl-Jensen, D., Fischer, H., Kipfstuhl, S., Kostick, C., Maselli, O.J., Mekhaldi, F., Mulvaney, R.,
manuscript submitted to Paleoceanography and Paleoclimatology
1286 Muscheler, R., Pasteris, D.R., Pilcher, J.R., Salzer, M., Schüpbach, S., Steffensen, J.P., Vinther, B.M., Woodruff, 1287 T.E., 2015. Timing and climate forcing of volcanic eruptions for the past 2,500 years. Nature 523, 543-549. 1288 Singarayer, J.S., Valdes, P.J., Roberts, W.H.G., 2017. Ocean dominated expansion and contraction of the late 1289 Quaternary tropical rainbelt. Scientific Reports 7, 9382. 1290 Sinha, A., Stott, L., Berkelhammer, M., Cheng, H., Edwards, R.L., Buckley, B., Aldenderfer, M., Mudelsee, M., 1291 2011. A global context for megadroughts in monsoon Asia during the past millennium. Quat. Sci. Rev. 30, 47-62. 1292 Sletten, H.R., Railsback, L.B., Liang, F., Brook, G.A., Marais, E., Hardt, B.F., Cheng, H., Edwards, R.L., 2013. A 1293 petrographic and geochemical record of climate change over the last 4600years from a northern Namibia stalagmite, 1294 with evidence of abruptly wetter climate at the beginning of southern Africa's Iron Age. Palaeogeography, 1295 Palaeoclimatology, Palaeoecology 376, 149-162. 1296 Stansell, N.D., Steinman, B.A., Abbott, M.B., Rubinov, M., Roman-Lacayo, M., 2012. Lacustrine stable isotope 1297 record of precipitation changes in Nicaragua during the Little Ice Age and Medieval Climate Anomaly. Geology 41, 1298 151-154. 1299 Swingedouw, D., Terray, L., Cassou, C., Voldoire, A., Salas-Mélia, D., Servonnat, J., 2011. Natural forcing of 1300 climate during the last millennium: fingerprint of solar variability. Clim. Dyn. 36, 1349-1364. 1301 Taylor, K.E., Crucifix, M., Braconnot, P., Hewitt, C.D., Doutriaux, C., Broccoli, A.J., Mitchell, J.F.B., Webb, M.J., 1302 2007. Estimating shortwave radiative forcing and response in climate models. J. Clim. 20, 2530-2543. 1303 Tierney, J.E., Abram, N.J., Anchukaitis, K.J., Evans, M.N., Giry, C., Kilbourne, K.H., Saenger, C.P., Wu, H.C., 1304 Zinke, J., 2015. Tropical sea surface temperatures for the past four centuries reconstructed from coral archives. 1305 Paleoceanography 30, 226-252. 1306 Tierney, J.E., Oppo, D.W., Rosenthal, Y., Russell, J.M., Linsley, B.K., 2010. Coordinated hydrological regimes in 1307 the Indo-Pacific region during the past two millennia. Paleoceanography 25. 1308 Toohey, M., Sigl, M., 2017. Volcanic stratospheric sulfur injections and aerosol optical depth from 500 BCE to 1309 1900 CE. Earth Syst. Sci. Data 9, 809-831. 1310 Vuille, M., Burns, S.J., Taylor, B.L., Cruz, F.W., Bird, B.W., Abbott, M.B., Kanner, L.C., Cheng, H., Novello, V.F., 1311 2012. A review of the South American monsoon history as recorded in stable isotopic proxies over the past two 1312 millennia. Climate of the Past 8, 1309-1321. 1313 Vuille, M., Werner, M., 2005. Stable isotopes in precipitation recording South American summer monsoon and 1314 ENSO variability: observations and model results. Clim. Dyn. 25, 401-413. 1315 Wallace, J.M., Rasmusson, E.M., Mitchell, T.P., Kousky, V.E., Sarachik, E.S., von Storch, H., 1998. On the 1316 structure and evolution of ENSO-related climate variability in the tropical Pacific: Lessons from TOGA. J. Geophys. 1317 Res. 103, 14,241-214,259. 1318 Wang, Y.J., Cheng, H., Edwards, R.L., He, Y.Q., Kong, X.G., An, Z.S., Wu, J.Y., Kelly, M.J., Dykoski, C.A., Li, 1319 X.D., 2005. The Holocene Asian monsoon: Links to solar changes and North Atlantic climate. Science 308, 854- 1320 857. 1321 Wodzicki, K.R., Rapp, A.D., 2016. Long-term characterization of the Pacific ITCZ using TRMM, GPCP, and ERA- 1322 Interim. Journal of Geophyscial Research: Atmospheres 121, 3153-3170. 1323 Yan, H., Sun, L., Wang, Y., Huang, W., Qiu, S., Yang, C., 2011. A record of the Southern Oscillation Index for the 1324 past 2,000 years from precipitation proxies. Nat. Geosci. 4, 611-614. 1325 Yan, H., Wei, W., Soon, W., An, Z., Zhou, W., Liu, Z., Wang, Y., Carter, R.M., 2015. Dynamics of the intertropical 1326 convergence zone over the western Pacific during the Little Ice Age. Nat. Geosci. 8, 315-320. 1327 Zhang, P.Z., Cheng, H., Edwards, R.L., Chen, F.H., Wang, Y.J., Yang, X.L., Liu, J., Tan, M., Wang, X.F., Liu, J.H., 1328 An, C.L., Dai, Z.B., Zhou, J., Zhang, D.Z., Jia, J.H., Jin, L.Y., Johnson, K.R., 2008. A Test of Climate, Sun, and 1329 Culture Relationships from an 1810-Year Chinese Cave Record. Science 322, 940-942. 1330
dĂďůĞϭ͘>ŝƐƚŽĨƚŚĞƉĂůĞŽĐůŝŵĂƚĞĚĂƚĂƵƐĞĚŝŶƚŚŝƐƐƚƵĚLJ
ZĞŐŝŽŶ ^ŝƚĞEĂŵĞ ů:ƵŶĐŽ>ĂŬĞ͕^ĂŶƌŝƐƚŽďĂů/ƐůĂŶĚ͕'ĂůĄƉĂŐŽƐ ĂƐƚĞƌŶWĂĐŝĨŝĐ ů:ƵŶĐŽ>ĂŬĞ͕^ĂŶƌŝƐƚŽďĂů/ƐůĂŶĚ͕'ĂůĄƉĂŐŽƐ WŽnjĂĚĞůĂƐŝĂďůĂƐ͕/ƐĂďĞůĂ/ƐůĂŶĚ͕'ĂůĄƉĂŐŽƐ ĞŶƚƌĂůWĂĐŝĨŝĐ tĂƐŚŝŶŐƚŽŶ/ƐůĂŶĚ͕<ŝƌŝďĂƚŝ
^ŽƵƚŚWĂĐŝĨŝĐ ,ĂƌĂŝ>ĂŬĞηϭ͕ZĞŶĚŽǀĂ/ƐůĂŶĚ͕^ŽůŽŵŽŶ/ƐůĂŶĚƐ;^,ZϭͿ ƉƵĂLJ͕dĂŚĂĂ͕&ƌĞŶĐŚWŽůLJŶĞƐŝĂ DϵϴͲϮϭϴϭ ƵŬŝƚƐƐĂŵĂǀĞ͕DĂůĂLJƐŝĂŶŽƌŶĞŽ ^ŶĂŝů^ŚĞůůĂǀĞ͕DĂůĂLJƐŝĂŶŽƌŶĞŽ;^^ϬϭͿ DϵϴͲϮϭϳϲ DĂŬĂƐƐĂƌ^ƚƌĂŝƚ͕/ŶĚŽŶĞƐŝĂ;ϯϭD͕ϯϰ''Ϳ DĂŬĂƐƐĂƌ^ƚƌĂŝƚ͕/ŶĚŽŶĞƐŝĂ;ĐŽŵƉŽƐŝƚĞ͗ϯϭD͕DϲϬ͕ϯϰ''͕ϯϮ''Ϳ tĞƐƚĞƌŶWĂĐŝĨŝĐͬ DĂƌŝƚŝŵĞŽŶƚŝŶĞŶƚ DĂŬĂƐƐĂƌ^ƚƌĂŝƚ͕/ŶĚŽŶĞƐŝĂ;DϵϴϮϭͲϲϬͿ ZĂŶƵ>ĂŵŽŶŐĂŶ͕ ĂƐƚ:ĂǀĂ͕/ŶĚŽŶĞƐŝĂ >ĂŬĞ>ĂĚŝŶŐ͕^t/ŶĚŽŶĞƐŝĂ >ĂŬĞ>ŽŐƵŶŐ͕ĞĂƐƚĞƌŶ:ĂǀĂ͕/ŶĚŽŶĞƐŝĂ >ŝĂŶŐ>ƵĂƌĂǀĞ͕&ůŽƌĞƐ͕/ŶĚŽŶĞƐŝĂ;ƐƉůŝĐĞ͗>ZϬϲͲϭ͕>ZϬϲͲϯͿ DϵϴͲϮϭϳϬ ƵƐƚƌĂůŝĂ ĂǀĞ
hŶƵƐĞĚƌĞĐŽƌĚƐ;>ĂĐŚŶŝĞƚ͕ϮϬϭϱƐƵŐŐĞƐƚƐƚŚĞƐĞƐƚĂůĂŐŵŝƚĞƐŽƵƚŽĨŝƐŽƚŽƉŝĐĞƋƵŝůŝďƌŝƵŵǁŝƚŚĐ DĞƐŽĂŵĞƌŝĐĂ zŽŬĂůƵŵĂǀĞ͕ĞůŝnjĞ ^ŽƵƚŚͬĞĂƐƚƐŝĂ >ŝĂŶŚƵĂĂǀĞ͕ĞĂƐƚĞƌŶŚŝŶĂ ^ŽƵƚŚŵĞƌŝĐĂ >ĂŐŽhŵĂLJŽ͕WĞƌƵ /ŶĚŝĂŶKĐĞĂŶ ŶũŽŚŝďĞĂǀĞ͕DĂĚĂŐĂƐĐĂĂƌ >Ăƚ >ŽŶ ůĞǀĂƚŝŽŶ ;ĚĞŐEͿ ;ĚĞŐ Ϳ ;ŵͿ WƌŽdžLJĂƌĐŚŝǀĞ DĞĂƐƵƌĞŵĞŶƚ ͲϬ͘ϵ Ͳϴϵ͘ϱ ϲϳϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĚŝŶŽƐƚĞƌŽů , ͲϬ͘ϵ Ͳϴϵ͘ϱ ϲϲϬ ůĂŬĞƐĞĚŝŵĞŶƚ ƐĞĚŝŵĞŶƚŐƌĂŝŶƐŝnjĞ;йƐĂŶĚͿ Ͳϭ͘Ϭ Ͳϵϭ͘Ϭ ůĂŬĞƐĞĚŝŵĞŶƚ ,ůŝƉŝĚďŝŽŵĂƌŬĞƌƐ ϰ͘ϳ ͲϭϲϬ͘ϰ ůĂŬĞƐĞĚŝŵĞŶƚ ,ŽĨƚŽƚĂůůŝƉŝĚĞdžƚƌĂĐƚ Ͳϴ͘ϲ ϭϱϳ͘ϰ ůĂŬĞƐĞĚŝŵĞŶƚ ĚŝŶŽƐƚĞƌŽů , Ͳϭϲ͘ϳ Ͳϭϱϭ͘ϱ Ϭ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ dŝͬĂ ϲ͘ϯ ϭϮϱ͘ϴ ͲϮϭϭϰ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ KĂŶĚDŐͬĂ;ĨŽƌĂŵƐͿ ϰ͘ϯ ϭϭϱ͘Ϭ ϮϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϰ͘Ϯ ϭϭϰ͘ϵ ϭϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϬ͘Ϯ ϭϯϯ͘ϰ ͲϮϯϴϮ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ KĂŶĚDŐͬĂ;ĨŽƌĂŵƐͿ Ͳϯ͘ϵ ϭϭϵ͘ϱ ͲϰϱϵƚŽͲϱϬϯ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ,ϯϬĨĂƚƚLJĂĐŝĚ Ͳϯ͘ϵ ϭϭϵ͘ϱ ͲϰϱϰƚŽͲϱϬϯ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ KĂŶĚDŐͬĂ;ĨŽƌĂŵƐͿ Ͳϱ͘Ϯ ϭϭϳ͘ϱ Ͳϭϭϴϱ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ KĂŶĚDŐͬĂ;ĨŽƌĂŵƐͿ Ͳϴ͘Ϭ ϭϭϯ͘ϰ ϮϰϬ ůĂŬĞƐĞĚŝŵĞŶƚ ŵŽůйDŐŝŶĐĂůĐŝƚĞ Ͳϴ͘Ϭ ϭϭϯ͘ϯ ϯϮϰ ůĂŬĞƐĞĚŝŵĞŶƚ ,ŶͲĂůŬĂŶŽŝĐĂĐŝĚ Ͳϴ͘Ϭ ϭϭϯ͘ϯ Ϯϭϱ ůĂŬĞƐĞĚŝŵĞŶƚ ŽĨŽƌŐĂŶŝĐŵĂƚĞƌŝĂů Ͳϴ͘ϱ ϭϮϬ͘ϰ ϱϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϭϬ͘ϲ ϭϮϱ͘ϰ ͲϴϯϮ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ KĂŶĚDŐͬĂ;ĨŽƌĂŵƐͿ Ͳϭϱ͘Ϯ ϭϮϴ͘ϰ ϭϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϯ͘ϲ ϭϬϱ͘ϭ ϭϲϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϯ͘ϲ ϭϬϵ͘ϭ ϭϰϵϱ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϯ͘ϲ ϭϬϵ͘ϭ ϭϰϵϱ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϯ͘ϯ ϭϬϱ͘Ϭ ϭϮϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϭ͘ϳ ϭϭϬ͘ϰ ϭϵϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϭ͘ϳ ϭϭϬ͘ϰ ϭϵϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϯϬ͘ϱ ϭϭϬ͘ϰ Ϯϵϰ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ϯϱ͘ϯ ϭϬϴ͘ϭ ϲϴϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ϯϱ͘ϯ ϭϬϴ͘ϭ ϲϴϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ϯϱ͘ϯ ϭϬϴ͘ϭ ϲϴϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ϯϰ͘ϰ ϭϬϮ͘ϴ ϭϳϮϭ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĂƵƚŚŝŐĞŶŝĐͿ ϮϬ͘ϭ ϭϭϳ͘ϰ ͲϭϳϮϳ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĚŝĂƚŽŵƐͿ ϭϵ͘Ϭ ϴϮ͘Ϭ ϰϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϭϴ͘ϵ ϴϭ͘ϵ ϲϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϭϲ͘ϳ ϭϭϮ͘ϳ ůĂŬĞƐĞĚŝŵĞŶƚ ƐĞĚŝŵĞŶƚŐƌĂŝŶƐŝnjĞ ϭϲ͘ϳ ϭϭϮ͘ϳ ůĂŬĞƐĞĚŝŵĞŶƚ ƐĞĚŝŵĞŶƚŐƌĂŝŶƐŝnjĞ ϭϲ͘ϳ ϭϭϮ͘ϳ ůĂŬĞƐĞĚŝŵĞŶƚ ƐĞĚŝŵĞŶƚŐƌĂŝŶƐŝnjĞ ϲ͘ϱ Ͳϭ͘ϰ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĂƵƚŚŝŐĞŶŝĐͿ ͲϬ͘ϰ Ϯϵ͘ϴ ϵϭϳ ůĂŬĞƐĞĚŝŵĞŶƚ йDŐŝŶĐĂůĐŝƚĞ Ͳϭϵ͘ϰ ϭϳ͘ϵ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϮϳ͘ϯ ϯϮ͘ϲ ϮϬ ůĂŬĞƐĞĚŝŵĞŶƚ йƉůĂŶŬƚŽŶŝĐĚŝĂƚŽŵƐ ϮϬ͘ϳ Ͳϴϵ͘ϱ ϮϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϮϬ͘ϲ Ͳϴϳ͘ϱ ϭϰ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;LJƚŚĞƌŝĚĞůůĂŝůŽƐǀĂLJŝͿ ϮϬ͘ϲ Ͳϴϳ͘ϱ ϭϰ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;WLJƌŐŽƉŚŽƌƵƐĐŽƌŽŶĂƚƵƐͿ ϮϬ͘ϲ ͲϭϬϰ͘ϳ ϮϬϬϬ ůĂŬĞƐĞĚŝŵĞŶƚ dŝĐŽŶĐĞŶƚƌĂƚŝŽŶ ϭϵ͘ϵ Ͳϴϴ͘ϴ ůĂŬĞƐĞĚŝŵĞŶƚ ůĂŬĞƐĞĚŝŵĞŶƚĚĞŶƐŝƚLJ ϮϬ͘ϲ Ͳϴϵ͘ϳ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;WLJƌŐŽƉŚŽƌƵƐĐŽƌŽŶĂƚƵƐͿ ϭϵ͘ϭ Ͳϵϳ͘ϱ Ϯϯϳϲ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĂƵƚŚŝŐĞŶŝĐͿ ϭϳ͘ϰ Ͳϵϵ͘Ϯ ϵϯϰ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ϭϳ͘Ϭ Ͳϴϵ͘ϴ ϭϭϬ ůĂŬĞƐĞĚŝŵĞŶƚ ŵĂŐŶĞƚŝĐƐƵƐĐĞƉƚŝďŝůŝƚLJ ϭϲ͘ϵ Ͳϴϵ͘ϴ ϭϭϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ŽĐŚůŝŽƉŝŶĂƐƉ͘Ϳ ϭϲ͘ϵ Ͳϴϵ͘ϴ ϭϭϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;LJƚŚĞƌŝĚĞůůĂŝůŽƐǀĂLJŝͿ ϭϲ͘ϵ Ͳϴϵ͘ϴ ϭϭϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;WLJƌŐŽƉŚŽƌƵƐƐƉ͘Ϳ ϭϭ͘ϵ Ͳϴϱ͘ϵ ϰϰ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ŽƐƚƌŽĐŽĚͿ ϭϬ͘ϳ Ͳϲϱ͘Ϯ Ͳϴϵϯ ŵĂƌŝŶĞƐĞĚŝŵĞŶƚ dŝĐŽŶĐĞŶƚƌĂƚŝŽŶ ϰ͘ϱ Ͳϳϯ͘ϵ ϮϬϳϬ ůĂŬĞƐĞĚŝŵĞŶƚ йůŝƚŚŝĐƐ ͲϮ͘ϴ Ͳϳϵ͘Ϯ ϰϮϬϬ ůĂŬĞƐĞĚŝŵĞŶƚ ƌĞĚĐŽůŽƌŝŶƚĞŶƐŝƚLJ ͲϮ͘ϴ Ͳϳϵ͘Ϯ ϰϮϬϬ ůĂŬĞƐĞĚŝŵĞŶƚ ŐƌĞLJƐĐĂůĞ Ͳϱ͘ϵ Ͳϳϳ͘ϰ ϴϳϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ͳϱ͘ϵ Ͳϳϳ͘ϯ ϴϲϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ͳϱ͘ϵ Ͳϳϳ͘ϯ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ͳϲ͘ϭ Ͳϳϳ͘Ϯ ϵϬϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϭϬ͘ϳ Ͳϳϲ͘ϭ ϰϯϬϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĂƵƚŚŝŐĞŶŝĐͿ Ͳϭϭ͘ϯ Ͳϳϱ͘ϴ ϯϴϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϭϮ͘ϰ Ͳϰϭ͘ϲ ϲϴϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K Ͳϭϯ͘ϵ ͲϳϬ͘ϴ ϱϲϳϬ ŐůĂĐŝĂůŝĐĞ ŝĐĞ K Ͳϭϱ͘Ϯ Ͳϱϲ͘ϴ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ͲϮϰ͘ϲ Ͳϰϴ͘ϲ ϭϯϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K
ĐĂǀĞĚƌŝƉǁĂƚĞƌƐͿ ϭϲ͘Ϯ Ͳϴϵ͘ϭ ϯϲϲ ƐƉĞůĞŽƚŚĞŵ ĚϭϴKƐƉĞůĞŽ Ϯϵ͘ϱ ϭϬϵ͘ϱ ϰϱϱ ƐƉĞůĞŽƚŚĞŵ ĚϭϴKƐƉĞůĞŽ Ͳϭϱ͘ϰ Ͳϵϳ͘ϱ ϯϴϴϬ ůĂŬĞƐĞĚŝŵĞŶƚ ĐĂƌďŽŶĂƚĞ K;ĚŝĂƚŽŵƐͿ Ͳϭϱ͘ϱ ϰϲ͘ϵ ϱϬ ƐƉĞůĞŽƚŚĞŵ ĐĂƌďŽŶĂƚĞ K ůƵƐƚĞƌη͕ ǀŐƌĞƐŽůƵƚŝŽŶ /ŶĚĞdžη ;LJĞĂƌƐͿ ZĞĨĞƌĞŶĐĞ ĂƚĂƐĞƚhZ> ϭ͕ϳ ϲϳ ƚǁŽŽĚĞƚĂů͕͘ϮϬϭϰ ŚƚƚƉƐ͗ͬͬĚŽŝ͘ŽƌŐͬϭϬ͘ϭϱϵϰͬWE' ͘ϴϯϰ ϭ͕ϭϯ ϭϭ ŽŶƌŽLJĞƚĂů͕͘ϮϬϬϴ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕Ϯϯ ϰϱ EĞůƐŽŶĂŶĚ^ĂĐŚƐ͕ϮϬϭϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ƉŶĂƐ͘ŽƌŐͬĐŽŶƚĞŶƚͬƐƵƉƉůͬ ϰϮ ^ĂĐŚƐĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĂƚƵƌĞ͘ĐŽŵͬĂƌƚŝĐůĞƐͬŶŐĞŽ ϭϴϱ DĂůŽŶĞLJ͕ϮϬϭϴ ϭ͕ϯϯ ϭ dŽŽŵĞLJĞƚĂů͕͘ϮϬϭϲ ŚƚƚƉƐ͗ͬͬĂŐƵƉƵďƐ͘ŽŶůŝŶĞůŝďƌĂƌLJ͘ǁŝůĞLJ͘ĐŽ ϭ͕ϰϬ ϯϰ ^ƚŽƚƚĞƚĂů͕͘ϮϬϬϰ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϲ ŚĞŶĞƚĂů͕͘ϮϬϭϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϭϮ ϱϯ ĂƌŽůŝŶĞƚĂů͕͘ϮϬϭϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϰϭ ϰϮ ^ƚŽƚƚĞƚĂů͕͘ϮϬϬϰ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϯϮ ϯϯ dŝĞƌŶĞLJĞƚĂů͕͘ϮϬϭϬ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕Ϯϲ ϭϬ KƉƉŽĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϵ EĞǁƚŽŶĞƚĂů͕͘ϮϬϬϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ ƌĂƵƐďĂLJĞƚĂů͕͘ϮϬϬϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϵ <ŽŶĞĐŬĞLJĞƚĂů͕͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕Ϯϴ ϭϳ ZŽĚLJƐŝůůĞƚĂů͕͘ϮϬϭϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϭϲ ϭϮ 'ƌŝĨĨŝƚŚƐĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭϲϯ ^ƚŽƚƚĞƚĂů͕͘ϮϬϬϰ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ ĞŶŶŝƐƚŽŶĞƚĂů͕͘ϮϬϭϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϯϬ ϱ dĂŶĞƚĂů͕͘ϮϬϭϬ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϭϭ ϰ ĂŝĞƚĂů͕͘ϮϬϭϬĂ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϭϭ Ϯϳ ĂŝĞƚĂů͕͘ϮϬϭϬď ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϯϳ ϯ ŚĂŶŐĞƚĂů͕͘ϮϬϬϴ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯϲ ŽŶŐĞƚĂů͕͘ϮϬϭϬĂ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯϲ ŽŶŐĞƚĂů͕͘ϮϬϭϬď ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϭϵ ϯ ,ƵĞƚĂů͕͘ϮϬϬϴ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϭϰ ϴ LJŬŽƐŬŝĞƚĂů͕͘ϮϬϬϱ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϯϱ ϰ tĂŶŐĞƚĂů͕͘ϮϬϬϱ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ ŚĂŽĞƚĂů͕͘ϮϬϭϱ ĨƚƉ͗ͬͬĨƚƉ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉƵďͬĚĂƚĂͬƉĂůĞ ϯ͕ϭϴ ϯ ,ŝůůŵĂŶĞƚĂů͕͘ϮϬϭϰ ŚƚƚƉ͗ͬͬŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵĚLJͬϭϳϱ Ϯ͕ϯϲ ϭϳ tĂŶŐĞƚĂů͕͘ϭϵϵϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ ĞƌŬĞůŚĂŵŵĞƌĞƚĂů͕͘ϮϬϭϬ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ ^ŝŶŚĂĞƚĂů͕͘ϮϬϭϭ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭϰ zĂŶĞƚĂů͕͘ϮϬϭϭĂ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯϭ zĂŶĞƚĂů͕͘ϮϬϭϭď ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϲ zĂŶĞƚĂů͕͘ϮϬϭϭĐ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕Ϯϵ ϱ ^ŚĂŶĂŚĂŶĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϰ ZƵƐƐĞůůĞƚĂů͕͘ϮϬϬϳ ŚƚƚƉƐ͗ͬͬĨŝŐƐŚĂƌĞ͘ĐŽŵͬĂƌƚŝĐůĞƐͬWƌŽdžLJͺĚ ϭϭ ^ůĞƚƚĞŶĞƚĂů͕͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϯϵ ϭϴ ^ƚĂŐĞƌĞƚĂů͕͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϰ ϯ DĞĚŝŶĂͲ ůŝnjĂůĚĞĞƚĂů͕͘ϮϬϭϬ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϭ ϳ ƵƌƚŝƐĞƚĂů͕͘ϭϵϵϲĂ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϭ ϳ ƵƌƚŝƐĞƚĂů͕͘ϭϵϵϲď ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ DĞƚĐĂůĨĞĞƚĂů͕͘ϮϬϭϬ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϯ ϰ ,ŽĚĞůůĞƚĂů͕͘ϮϬϬϴ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕Ϯ ϰ ,ŽĚĞůůĞƚĂů͕͘ϮϬϬϱ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϴ ϭϴ ŚĂƚƚĂĐŚĂƌLJĂĞƚĂů͕͘ϮϬϭϱ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕Ϯϭ Ϯ >ĂĐŚŶŝĞƚĞƚĂů͕͘ϮϬϭϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϭϱ ϱ ƐĐŽďĂƌĞƚĂů͕͘ϮϬϭϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯϱ ƵƌƚŝƐĞƚĂů͕͘ϭϵϵϴĂ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭϰ ƵƌƚŝƐĞƚĂů͕͘ϭϵϵϴď ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϱ ϭϮ ƵƌƚŝƐĞƚĂů͕͘ϭϵϵϴĐ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϱ ^ƚĂŶƐĞůů͕ϮϬϭϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϭϳ ϱ ,ĂƵŐĞƚĂů͕͘ϮϬϬϭ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕ϭϬ ϭϬ ŝƌĚĞƚĂů͕͘ϮϬϭϳ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϮϮ Ϭ͘ϱ DŽLJĞƚĂů͕͘ϮϬϬϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕Ϯϳ Ϭ͘ϯ ZŽĚďĞůůĞƚĂů͕͘ϭϵϵϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϯϴ ϯ ƉĂĠƐƚĞŐƵŝĞƚĂů͕͘ϮϬϭϰ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϳϭ ŚĞŶŐĞƚĂů͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϲ Ϯϵ ǀĂŶƌĞƵŬĞůĞŶĞƚĂů͕͘ϮϬϬϴ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ ZĞƵƚĞƌĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϵ Ϯ ŝƌĚĞƚĂů͕͘ϮϬϭϭ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕ϮϬ ϱ <ĂŶŶĞƌĞƚĂů͕͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϯ͕Ϯϰ ϱ EŽǀĞůůŽĞƚĂů͕͘ϮϬϭϮ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϯϭ ϭϬ dŚŽŵƉƐŽŶĞƚĂů͕͘ϮϬϭϯ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ ϭ͕Ϯϱ ϭ EŽǀĞůůŽĞƚĂů͕͘ϮϬϭϲ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ŶŽĂĂ͘ŐŽǀͬƉĂůĞŽͬƐƚƵ Ϯ͕ϯϰ Ϯ sƵŝůůĞĞƚĂů͕͘ϮϬϭϮͬdĂLJůŽƌ͕ϮϬϭϬ
ϲϳ ϱϯ
ZĞĂƐŽŶ <ĞŶŶĞƚƚĞƚĂů͕͘ϮϬϭϮ ŽƐĨŽƌĚĞƚĂů͕͘ϮϬϬϵ Ϯ͕ϴ ϱϬ ĂŬĞƌĞƚĂů͕͘ϮϬϬϵ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ EŽĂŐĞĐŽŶƚƌŽůƉŽŝŶ ϭ͕ϰϬ ϯ sŽĂƌŝŶƚƐŽĂĞƚĂů͕͘ϮϬϭϳ ŚƚƚƉƐ͗ͬͬǁǁǁ͘ŶĐĚĐ͘ ĂƚĂĐŽǀĞƌĂŐĞŝƐфϱ EŽƚĞ͗Ϭ͘ϱͲϮĚĞŐŽĨĨƐĞƚĂĚĚĞĚƚŽůŽŶĨŽƌĚŝĨĨĞƌĞŶƚƐƚƵĚŝĞƐĂƚƐĂŵĞƐŝƚĞĨŽƌ
WƌŽdžLJŬĞLJ͗ϬсĐĂǀĞ͕ϭсůĂŬĞ͕ϮсŽĐĞĂŶ͕ϯсŝĐĞ ϰϭϬϯ ƵĚLJͬϲϭϬϵ ͬϮϬϭϲͬϬϯͬϬϵͬϭϱϭϲϮϳϭϭϭϯ͘^ƵƉƉůĞŵĞŶƚĂů Žϱϱϰη^ĞĐϭϭ
ŵͬĚŽŝͬĨƵůůͬϭϬ͘ϭϬϬϮͬϮϬϭϱWϬϬϮϴϳϬ ƵĚLJͬϲϮϯϲ ƵĚLJͬϭϵϴϴϱ ƵĚLJͬϭϵϴϬϲ ƵĚLJͬϲϮϯϲ ƵĚLJͬϭϬϰϯϴ ƵĚLJͬϴϲϵϵ ƵĚLJͬϱϱϯϰ ƵĚLJͬϮϮϰϬϵ ƵĚLJͬϭϰϭϮϵ ƵĚLJͬϭϯϭϳϳ ƵĚLJͬϴϲϯϮ ƵĚLJͬϲϮϯϲ ƵĚLJͬϮϬϱϯϬ ƵĚLJͬϭϭϭϳϳ ƵĚLJͬϵϳϰϮ ƵĚLJͬϵϳϰϮ ƵĚLJͬϴϲϮϵ ƵĚLJͬϭϬϰϰϱ ƵĚLJͬϭϬϰϰϱ ƵĚLJͬϲϬϵϱ ƵĚLJͬϱϰϰϭ ƵĚLJͬϱϰϯϵ ĞŽͬƐƉĞůĞŽƚŚĞŵͬƐŝĂͬŚŝŶĂͬŽŶŐŐĞϮϬϭϱ͘džůƐ ϱϭϱ ƵĚLJͬϮϱϵϴ ƵĚLJͬϴϲϯϵ ƵĚLJͬϭϭϵϯϳ ƵĚLJͬϭϭϭϵϳ ƵĚLJͬϭϭϭϵϳ ƵĚLJͬϭϭϭϵϳ ƵĚLJͬϴϲϱϳ ĚĂƚĂͬϰϳϴϳϱϳϱͬϭ ƵĚLJͬϭϰϮϲϴ ƵĚLJͬϭϰϲϮϵ ƵĚLJͬϵϳϵϭ ƵĚLJͬϱϰϴϰ ƵĚLJͬϱϰϴϰ ƵĚLJͬϭϵϳϲϴ ƵĚLJͬϲϭϬϴ ƵĚLJͬϲϭϬϳ ƵĚLJͬϭϳϳϯϱ ƵĚLJͬϭϮϵϳϮ ƵĚLJͬϭϯϲϳϱ ƵĚLJͬϱϰϴϮ ƵĚLJͬϱϰϴϮ ƵĚLJͬϱϰϴϮ ƵĚLJͬϭϯϭϵϱ ƵĚLJͬϮϱϲϬ ƵĚLJͬϮϮϮϳϱ ƵĚLJͬϱϰϴϴ ƵĚLJͬϱϰϴϳ ƵĚLJͬϭϵϯϬϭ ƵĚLJͬϭϯϴϯϲ ƵĚLJͬϵϳϵϬ ƵĚLJͬϴϲϯϬ ƵĚLJͬϭϭϭϳϰ ƵĚLJͬϭϲϰϬϱ ƵĚLJͬϭϯϲϳϬ ƵĚLJͬϭϰϭϳϰ ƵĚLJͬϮϬϬϰϯ ŶƚƐǁŝƚŚŝŶϯϱϬͲϭϵϱϬ ϱϬйŽĨůĂƐƚŵŝůůĞŶŶŝƵŵ ƌǀŝƐƵĂůĐůĂƌŝƚLJ