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Late Pleistocene changes in winter moisture source in the coastal southwest United States

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Late Pleistocene changes in winter moisture source in the coastal southwest United States

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Content Late Pleistocene changes in winter moisture source in the coastal southwest United States Mong Sin Wu Department of Earth Sciences, The University of Southern California Master Thesis December 2014 Supervisor: Dr. Sarah Feakins TABLE OF CONTENTS ABSTRACT ------------------------------------------------------------------------------------------- INTRODUCTION ------------------------------------------------------------------------------------ Leaf wax hydrogen isotopic composition as a proxy of hydroclimate --------------- STUDY REGION ------------------------------------------------------------------------------------- Lake Elsinore --------------------------------------------------------------------------------- Regional climatology ------------------------------------------------------------------------ Regional vegetation -------------------------------------------------------------------------- MATERIALS AND METHODS ------------------------------------------------------------------- Core collection ------------------------------------------------------------------------------- Age model ------------------------------------------------------------------------------------ Sampling of river water and suspended sediment --------------------------------------- Water isotopic analysis ---------------------------------------------------------------------- Lipid extraction and isotopic analysis ----------------------------------------------------- Bulk sedimentary and geochemical analysis --------------------------------------------- Pollen analysis -------------------------------------------------------------------------------- LOVECLIM climate model simulations -------------------------------------------------- RESULTS AND DISCUSSIONS ------------------------------------------------------------------- Modern-day leaf wax and river water δD and implications for ε wax/w ---------------- Corrections for δD wax reconstructions ----------------------------------------------------- Accounting for vegetation changes ----------------------------------------------- Accounting for ice volume changes ---------------------------------------------- Lake Elsinore past δD wax , δD precip and comparison with modern-day hydroclimate Lake Elsinore paleoclimate record and implications to atmospheric circulations -- Regional comparisons ----------------------------------------------------------------------- Climate model comparisons ---------------------------------------------------------------- CONCLUSIONS -------------------------------------------------------------------------------------- REFERENCE CITED -------------------------------------------------------------------------------- FIGURES AND TABLES --------------------------------------------------------------------------- SUPPLEMENTARY MATERIALS ---------------------------------------------------------------- Leaf wax δ 13 C constraints on sources of leaf wax biomarkers ------------------------ High sand unit -------------------------------------------------------------------------------- OSL dating methods ---------------------------------------------------------------- Supplementary references ------------------------------------------------------------------ Supplementary figures and tables ---------------------------------------------------------- 1 2 4 7 7 8 8 9 9 10 10 11 11 12 13 13 14 14 16 16 17 18 19 23 26 27 29 42 51 51 52 54 55 58 PREAMBLE This paper was a collaborative work. I was the main author and wrote the paper, conducted chemical extractions and isotopic analysis of leaf wax, generated the age model using BACON software, and carried out the data analysis. In addition, several people contributed to this paper. Sarah Feakins provided the main supervision to the research, provided most of the leaf wax hydrogen isotope data during the deglacial (9 – 19 cal. kyr BP), aided in the interpretation of the data, and was the main editor of the paper. Matthew Kirby collected the sediment core and provided the core samples, provided the radiocarbon dates and guided the use of BACON software towards constructing the age model, provided data of grain size, total carbonate, and total organic matter, wrote the section of the interpretation of high sand unit in the supplemental, aided in the interpretation of the data, and provided general guidance. Laurie Menviel carried out the LOVECLIM time-transient model simulation, provided the model data, and aided in the proxy-model comparison. Linda Heusser provided the graminoid pollen data necessary for the graminoid-correction of leaf wax hydrogen isotope data. Camilo Ponton assisted in the field sampling and filtration of river water and suspended sediments, and carried out the isotopic analysis of river water samples. Edward Rhodes provided the OSL dates and wrote the method section of OSL dating in the supplemental. This study was supported by U.S. National Science Foundation Grants 1203549 to Sarah Feakins as well as 031511, 0731843, and 1203549 to Matthew Kirby. Coring was supported by the American Chemical Society Petroleum Research Grant #4187-B8 to Matthew Kirby. I thank Mr. Pat Kilroy (lake manager), the City of Lake Elsinore, John Gregg and Gregg Drilling and Testing, Inc., Joe Holbrook, M.F.A. and the CSUF School of Theatre and Dance for recovering and opening the cores. I thank Alex Sessions for access to the Los Gatos water isotope analyzer at the California Institute of Technology; Miguel Rincon for laboratory assistance; Jessica Tierney for her guidance in ice volume correction; Josh West for guidance and assistance with river sampling and POM filtration; and Axel Timmermann for discussing model simulations and proxy-model comparisons as well as Steven Sherwood for his inputs about precipitation patterns over California. LOVECLIM model experiments were performed on a computational cluster owned by the Faculty of Science of the University of New South Wales, Australia. This paper was submitted in October 2014 to Quaternary Science Reviews. 1 Late Pleistocene changes in winter moisture source in the coastal southwest United States ABSTRACT During the Late Pleistocene, millennial-scale abrupt transitions in the North Atlantic may have influenced climate in remote regions including the western United States (US). Several speleothem reconstructions from the Great Basin and interior southwest US capture oxygen isotopic evidence for shifts in Pacific and North American Monsoon derived storm tracks, revealing the power of isotopes in precipitation to track hydroclimatic regime change. Here, we present a high-resolution lake sediment leaf wax hydrogen isotope record spanning 9 – 32 cal. kyr BP from Lake Elsinore, California, in the coastal southwest US. This site provides a valuable archive within an exclusively Pacific-influenced precipitation zone. We use leaf wax hydrogen isotopic composition ( δD wax ) as a proxy for δD of precipitation ( δD precip ), a tracer of moisture sources and storm trajectories. Our interpretation is further based on grain size and graminoid pollen data from the same core, as well as time-transient experiments performed with the Earth System model LOVECLIM. We find large-magnitude (~70‰) millennial-scale fluctuations of δD wax during the late glacial, and a positive shift (>80‰) during the deglaciation, suggesting fundamental alterations in Pacific moisture delivery and storm trajectories. The largest transition (>80‰) is observed starting at the end of H1 and with no marked perturbation during the Younger Dryas. During the glacial, generally negative δD wax values reflect moisture delivery by more D-depleted northerly storms likely resulting from a southward-displacement of the polar jet stream. Based on model comparisons these conditions are likely associated with cooler conditions in the North Atlantic Ocean and a deepened Aleutian low pressure cell in the North 2 Pacific. Positive anomalies in δD wax values during glacial millennial-scale intervals as well as the larger magnitude deglacial positive shift indicate a transition to a mode of atmospheric circulation delivering moisture from the tropical and subtropical North Pacific. Late glacial δD wax fluctuations generally correlate in timing with North Atlantic abrupt climate transitions, indicating an Atlantic-Pacific teleconnection, which is further highlighted by climate model simulations. INTRODUCTION The Late Pleistocene experienced abrupt, large magnitude climatic variability, first discovered in the North Atlantic region, but increasingly recognized to be a highly dynamic time period globally. In particular, during the last glacial, a series of layers of ice rafted debris (IRD) in North Atlantic marine sediments marked massive discharges of icebergs from the Laurentide Ice Sheet (LIS) (Bond et al., 1992), commonly known as ‘Heinrich events’. Pa/Th ratios from Bermuda Rise sediments reveal a significant slowdown of the Atlantic Meridional Overturning Circulation (AMOC) during Heinrich Events 1 and 2 (hereafter H1 and H2), and perhaps partial disruption during H3 as a result of substantial freshening of the North Atlantic (McManus et al., 2004; Lippold et al., 2009). A weakened AMOC has been linked to climatic changes around the North Atlantic (Martrat et al., 2007) as well as in many other regions around the world (e.g. Peterson et al., 2000; Wang et al., 2001; Deplazes et al., 2014; Dubois et al., 2014). In response to disruption of the AMOC, climate models have also simulated changes in Pacific atmospheric and oceanic circulation including deepening of the Aleutian low pressure cell and enhanced formation of North Pacific Deep Water (NPDW; Okumura et al., 2009, Okazaki et al., 2010; Menviel et al., 2011). These changes in Pacific ocean-atmosphere conditions are expected to influence climate in western North America. During the instrumental 3 period, North American climate has been linked to North Pacific conditions through observational data and mechanistically tested through model experiments (Dettinger et al., 1998; Herweijer et al., 2006). During the last glacial and Holocene, proxy reconstructions of terrestrial hydroclimate in the western United States (US) have also been linked to Pacific Ocean conditions (Cook et al., 2004, MacDonald et al., 2008, Kirby et al. 2010; McCabe-Glynn et al., 2013; Dingemans et al., 2014; Kirby et al., 2014). Many climate records in the western US document a highly dynamic climate during the last glacial and deglaciation in terms of fluctuating glacier extent (Clark and Bartlein, 1995; Benson et al., 1996; 1998), precipitation (Maher et al., 2014, Ibarra et al., in review), and pluvial lake levels (Oviatt, 1997, Adams & Wesnousky, 1999; Lyle et al., 2012). Speleothem δ 18 O records offer high resolution archives of past hydroclimate and capture millennial-scale abrupt climatic variability (Oster et al., 2009; Asmerom et al., 2010; Wagner et al., 2010; Lachniet et al., 2014). The timing of several of these hydroclimatic shifts recorded in the western US suggests coincidence with North Atlantic abrupt climate changes during the last glacial and deglaciation. However, inland regions receive precipitation from both the winter westerlies and the late- summer to early fall North American Monsoon (NAM) (Asmerom et al., 2010; Wagner et al., 2010; Lachniet et al., 2014). Furthermore, variations in antecedent rainout also influence isotopic composition and thus modify moisture source signals as storms move inland (Lee and Fung, 2008). To investigate how North Atlantic abrupt climate changes might have affected atmospheric circulation in the eastern North Pacific and hence moisture delivery to the coastal southwest US, we develop a well-dated, high-resolution paleoclimate record from a sediment core from Lake Elsinore, California (US) spanning 9 – 32 cal. kyr BP. This study extends the 4 deglacial record from 9 – 19 cal. kyr BP previously published from the same core (Kirby et al., 2013). The time period covered by our record includes several abrupt climate transitions including HE 1 – 3, Bølling-Allerød (BA), and the Younger Dryas (YD), providing repeated opportunities to test the linkage between Atlantic and Pacific climatic events. Our record is located in the coastal southwest US where precipitation is almost exclusively derived from the Pacific Ocean (Cayan and Peterson, 1989). We compare our coastal southwest US record with speleothem records further inland in the western US to assess spatial differences and to deconvolve the different origins (e.g. Pacific only vs. Pacific and NAM precipitation) of hydroclimatic variability across the western US during late glacial and deglaciation. Leaf wax hydrogen isotopic composition as a proxy of hydroclimate Our approach is centered on the use of the hydrogen isotopic composition of plant leaf waxes ( δD wax ) to reconstruct the hydrogen isotopic composition of past precipitation ( δD precip ), which is assumed to be the same as the source of water to the plant ( δD w ). In addition we also measure weight percentage of sand, total carbonate and total organic matter to infer runoff and lake sedimentary conditions. δD wax primarily records δD precip subject to large biosynthetic fractionations (Sachse et al., 2012), and has been used as a proxy for δD precip in many lake sediments (Konecky et al., 2011; Tierney et al., 2011), including those spanning the last glacial (Tierney et al., 2008; Niedermeyer et al., 2010) and within the coastal southwest US (Kirby et al., 2013, 2014; Feakins et al., 2014). Some factors may affect hydroclimatic reconstructions using δD wax and we correct for these effects in our analysis. Firstly, changes in ice volume would change δD of sea water ( δD sw ) and hence δD precip . We account for the varying δD sw through our record, following the approach 5 of Tierney et al. (2011). Secondly, we consider the possible role of shifting vegetation in changing the net or apparent fractionation between δD w and δD wax, which is defined as ε wax/w = [( δD wax + 1) / ( δD w + 1) – 1] (eq. 1) The net fractionation can differ among plant types as suggested by plant based surveys (Smith and Freeman, 2006; Hou et al., 2007; Sachse et al., 2012). The largest difference is observed between non-graminoids and graminoids (grasses and sedges). The latter experiences a known larger ε wax/w due to growth from the basal meristem where more D-depleted leaf water is utilized in wax synthesis (Liu et al., 2006; Smith and Freeman, 2006; Hou et al., 2007; McInerney et al., 2011; Sachse et al., 2012). Feakins (2013) presented a correction for multiple plant types towards reconstructions of past hydrological changes based on isotopic offsets identified in Sachse et al. (2012). Since the largest changes, and those which are physiological rather than climatic in nature, are expected to derive from the proportion of graminoids, we simplify the vegetation-correction approach here to consider the role of graminoids only and apply a graminoid correction to the δD wax record to extract a record of changing δD w and thus δD precip . δD precip integrates the history of a water parcel in the atmosphere, including the source of moisture, upstream transport history, condensation height and temperature, as well as rainout processes (Dansgaard, 1964; Frankenberg et al., 2009; LeGrande and Schmidt, 2009; Buenning et al., 2012). In the western US, δD precip has been shown to be dictated dominantly by moisture source region and storm trajectories, with storms sourced from tropical and subtropical North Pacific delivering more D-enriched precipitation, whereas those sourced from the northern North Pacific deliver more D-depleted precipitation (Friedman et al., 1992, 2002; Berkelhammer et al., 2012; McCabe-Glynn et al., 2013). δD precip measurements and trajectory analysis of individual 6 storms arriving to the southern Sierra Nevada (2001 – 2005 A.D.) shows that δD precip varies from -23 to -68‰ for tropical Pacific storms to -98 to -158‰ for northern Pacific storms (Berkelhammer et al., 2012; McCabe-Glynn et al., 2013). Changes in major storm tracks are likely related to changes in large scale atmospheric circulation, including the mean position and strength of the polar jet stream, and the location and intensity of the characteristic North Pacific high and Aleutian low pressure cells (Berkelhammer et al., 2012), which are likely linked to changes in sea surface temperature patterns. Changes in sea surface temperature in the Kuroshio Current extension region over the past century have been correlated with fluctuations of δ 18 O recorded in a speleothem in the southern Sierra Nevada (McCabe-Glynn et al., 2013). Thus, over longer timescales, we expect that changes in oceanic and atmospheric circulation might likewise control the location of storm tracks and the isotopic composition of moisture received. Overall we interpret more D-depleted precipitation at Lake Elsinore as reflecting a northern North Pacific moisture source and northwesterly storm tracks, and D-enriched precipitation as reflecting tropical and subtropical moisture sources with southwesterly storm tracks. Although moisture source dominantly controls δD precip in the region, we acknowledge that minor influences may come from other isotopic effects including those associated with the temperature and elevation of condensation, the amount of antecedent precipitation, re- evaporation from raindrops during decent, as well as temporal variability within storms (Dansgaard, 1964; Lee and Fung, 2008; Coplen et al., 2008; Buenning et al., 2012;). Of note, the ‘amount effect’ which is not observed in individual storm measurements, is somewhat shown by temporal time series in southern California (Feakins et al., 2014; McCabe-Glynn et al., 2013, c.f. Berkelhammer et al., 2012). As a result the isotopic composition of the end members from the 7 different storm tracks may not remain constant across the glacial, but relative differences are expected to persist. STUDY REGION Lake Elsinore Lake Elsinore (33.66 °N, 117.35 °W, 380 masl, Fig. 1) is the largest natural freshwater lake (25 km 2 ) in the coastal southwest US (36 km inland). This is a rare lake in a semi-arid region with few natural lakes, particularly large lakes, making this a site of unique potential for a late Pleistocene hydroclimate reconstruction from the coastal southwest US. Several shorter paleoclimate reconstructions have already been generated from this site (Kirby et al., 2005, 2007; 2010; 2013) and here we develop a record that extends to 32 cal. ka BP. Lake Elsinore is fed by inflow from the ephemeral San Jacinto River which originates in the San Jacinto Mountains (maximum elevation 3302 m) and drains a catchment of about 1870 km 2 into Lake Elsinore (Fig. 1). Some additional discharge into the lake comes from surface runoff from the north slope of Santa Ana Mountains, proximal to the lake. However, this only represents a small area of the catchment that receives low precipitation amounts and its contribution is not substantial. Today, the annual mean discharge rate (from 1916 – 2012 A.D.) of San Jacinto River is 0.46 m 3 /s, with >80% of the total discharge happening between January and March (U.S. Geological Survey, 2014). Most runoff occurs during and following storm events a few days a year and during snowmelt from high elevations when flows can reach as high as 453 m 3 /s as recorded in Feb 27, 1927 A.D. (U.S. Geological Survey, 2014). Highly episodic lake inflows drive fluctuations in lake level and surface area in this (generally) closed basin lake (Kirby et al., 2010, 2004).Modern lake level averages ~8 m; episodic overflow has occurred when lake levels have exceeded 13 m depth (Lynch, 1931; Anderson, 2001). 8 Regional climatology The coastal southwest US experiences a winter-dominated precipitation seasonality with almost all precipitation falling between the months of October to May derived from westerly storm tracks from the Pacific (Cayan and Peterson, 1989). Mean annual precipitation (MAP) in the Lake Elsinore catchment ranges from 300 mm/yr at lowland elevations (~400 – 700 masl, comprising >80% of the catchment area), to orographically enhanced precipitation (rain and snow) of 800 mm/yr at elevations >1500 masl on Mt. San Jacinto (PRISM Climate Group, Oregon State University, 2013). At lake level mean annual temperature (MAT) averages 18°C, with summer average temperatures of 25°C, resulting in high potential evaporation rates and lake water loss of >1.4 m/yr (Anderson, 2001). The orographic effect of Mt San Jacinto allows for sufficient precipitation amounts to sustain a lake in this hot and semi-arid region. Regional vegetation In the Lake Elsinore catchment, vegetation is characterized by low-lying sage scrub, chamise (Adenostoma) and chaparral (dominated by Ceanothus) in the lowlands transitioning into scrub oak (Quercus) chaparral at mid elevations. Above ~1300 masl, the chaparral shrublands merge with full-stature oak (Quercus including Q. chrysolepis; canyon live oak) and Pine (Pinus, including P . lambertiana, sugar pine) woodlands transitioning into open mixed conifer forests on the peaks of the San Jacinto Mountains above 2500 masl (Hanawalt and Whittaker, 1977; Thorne, 1977). Precipitation and temperature gradients with elevation create a moisture availability gradient which explains the altitudinal distribution of vegetation in the catchment with full-stature trees only above the 1300 masl chaparral-forest transition (Fig. 1c). The major plant communities have persisted since the Tertiary; however, latitudinal and altitudinal migration in response to past and present climate change has produced the rich floral 9 diversity that typifies California vegetation (Barbour et al., 2007; Crimmins et al., 2011). While these vegetation assemblages have been largely persistent over our study period based on the study of pollen (Heusser, unpublished data), the elevation of the chaparral-forest transition may have shifted within the catchment and the relative predominance of different vegetation types may have varied in shifted association with climatic changes. MATERIALS AND METHODS Core collection Sediment core LEDC10-1 was collected in June 2010 at the depocenter of Lake Elsinore, as determined by seismic reflection data (Pyke et al., 2009). The core was extracted using a hollow stemmed auger push core system with a lined hole aboard a stabilized coring barge, with sediment recovery rate better than 90% of the total core length. Coring began at 9 m and ended at 30 m below the modern lake bed. The start at 9 m was intended to overlap with and extend Holocene cores collected previously from the lake (Kirby et al., 2010). The core was split and stored at 4°C and sampled at the California State University Fullerton (CSUF) Paleoclimatology and Paleotsunami Laboratory. The core stratigraphy was described, documented, and digitally photographed. The sediment core is mainly composed of dark-color silt mixed with clay in most depths, with a layer of light-color sandy material observed between 2334 and 2500 cm. The layer between 2517 and 2539 cm is considered as ‘reworked’ by unconsolidated sand material falling from above when the coring method shifted from the pound technique, which was used briefly to acquire the stiff sands between 2334 and 2500, back to the push technique. Hence data between 2517 and 2539 cm are not considered in the following discussion. 10 Age model The age model of the sediment core is based on twenty four samples of organic macrofossils (mostly charcoal) and seven bulk samples, which were collected, washed with acid to remove any carbonate, and sent to Lawrence Livermore National Laboratory’s Center for Accelerator Mass Spectrometry or to the University of California Irvine W.M. Keck Carbon Cycle Accelerator Mass Spectrometry Laboratory for radiocarbon dating (Table 1). Three bulk samples were collected at depths where macrofossils were also found in order to calculate the reservoir effect difference between bulk and macrofossil radiocarbon ages. The radiocarbon ages of the other four bulk samples, collected at depths where macrofossils were not found, were then corrected with this reservoir age by subtracting the reservoir age from the bulk age. The age model is based on Bayesian techniques implemented using the Bacon 2.2 program (Fig. 2). As an independent assessment, we also dated two sandy horizons towards the base using Infra-Red Stimulated Luminescence (IRSL) signals from single grains of K-feldspar (Fig. 2; details in Supplementary Text S2). Sampling of river water and suspended sediment In order to compare the paleoclimate record with the modern hydroclimate, and to assess the variation in leaf wax inputs across topography, river water samples were collected from ephemeral flow in the San Jacinto River, and its tributary Strawberry Creek, at three elevations (615, 630, and 1630 masl) including above and below the chaparral-forest transition at 1300 masl (Fig. 1b and c). Sampling was performed in the wet season (27 Jan 2013) during a flow event in San Jacinto River, which flows only for a few days following a storm. Sampling was achieved from the banks by immersing 20 L vessels into the stream. Water was filtered taking care to transfer 100% of the particulates by re-suspending and rinsing out sediments from the vessel. 11 Suspended sediment >0.2 µm was collected on polyethylsulfone (PES) filters (90 and 147 mm diameter) within pressurized filter units operated by bicycle pump (similar to the methods of Galy et al., 2011 and Ponton et al., 2014). The particulates were rinsed off the filter with milliQ water and subsequently freeze-dried using a Virtis 2k unit. River water was collected into tightly capped glass vials for water isotopic analysis. Water isotopic analysis δD and δ 18 O values of river water were measured on 8 replicate injections of 1 µL water samples using a Los Gatos Research DLT-1000 liquid water isotope analyzer at the California Institute of Technology. Replicate measurements yielded a mean precision (1σ) of 0.3‰ (n = 24) and were calibrated using 3 working standards (Maui Water, δD = −10.6‰, δ 18 O = −3.3‰; Caltech internal standard, δ D = −73.4‰, δ 18 O = −9.7‰; and LGR Water # 2, δ D = −117.0‰, δ 18 O = −15.5‰) to the VSMOW−SLAP isotopic scale with accuracy determined to better than 0.2‰ and 0.1‰ respectively. Lipid extraction and isotopic analysis The core was sampled for compound specific isotopic analysis at an average 26 cm interval with each sample integrating 2 cm. Additional samples were selected across key transitions identified in the first round of compound specific isotopic analyses. Sediment samples (typically ~5 g) were freeze-dried, ground to homogenize, then solvent-extracted with 9:1 v/v dichloromethane (DCM) to methanol (MeOH) at 100°C and 1500 psi for 30 minutes using an Accelerated Solvent Extraction system (ASE 350 ® , Dionex). Total lipid extracts were separated by column chromatography through LC-NH 2 gel into a neutral fraction and acid fraction, using 2:1 DCM:isopropanol and 4% formic acid in diethyl ether respectively. The acid fraction was methylated using MeOH of known isotopic composition in hydrochloric acid (19:1 v/v) at 70°C 12 for 12 h. The product was then diluted with milliQ water and partitioned into hexane. The hexane fraction was then further separated over column chromatography through silica gel (5% water- deactivated, 5 cm column) using hexane and DCM respectively. The DCM fraction contained the saturated fatty acid methyl esters (FAMEs), and was blown dry under nitrogen gas and dissolved in hexane for isotopic measurements. δD values of long chain n-alkanoic acid methyl esters (identified first by GC-MS) were measured by GC-IRMS (Thermo Scientific Trace gas chromatograph connected to a Delta V Plus mass spectrometer via an Isolink pyrolysis furnace at 1400°C) at the University of Southern California. δD values were normalized to the Vienna Standard Mean Ocean Water/Standard Light Antarctic Precipitation (VSMOW/SLAP) hydrogen isotopic scale by comparing with an external standard containing 15 n-alkane compounds (C 16 to C 30 ) obtained from A. Schimmelmann, Indiana University, Bloomington. The δD values of the external standard mixture span -9 to -254‰. The RMS error of replicate measurements of the standard was 3.2‰. An internal standard (C 34 n-alkane) was co-injected with each sample to check for stability throughout the sequence. The standard deviation of the internal standard was 3.6‰ across all analyses. δD values were then corrected for the δD of the added methyl group by mass balance ( δD of methanol = -198.3±3.9‰, based on derivatization of phthalic acid of known isotopic composition, and back calculating the δD of methanol). Bulk sedimentary and geochemical analysis Sediment grain size distribution and bulk geochemistry were analyzed at a 2 cm and 1 cm intervals respectively, with multi-decadal time resolution. After organic matter, carbonates, and biogenic silica were removed from the sediment with 30 mL 30% H 2 O 2 , 10 mL 1N HCl, and 10mL 1N NaOH respectively, grain size distribution was measured with a laser diffraction grain 13 size analyzer (Malvern Mastersizer 2000) coupled to a Hydro 2000G. A tuff standard (TS2) with a known distribution between 1 and 16 µm was measured in parallel to the samples to verify analytical stability throughout the run. Sand is defined by particles with a grain size distribution between 62.5 µm – 2 mm and reported as percentage by volume (sand %). Percent total carbonate (TC) was measured by loss on ignition procedure at 950 °C (Dean, 1974). Pollen analysis Samples for graminoid (Poaceae and Cyperaceae) pollen counts were collected mostly at every 2 to 10 cm. Samples were processed with standard methods, involving disaggregation of ~5 g of dry sample with 200 mL of 1% potassium hydroxide, sieved to retain the 7–150 μm size fraction, and successively treated with hydrochloric acid and hydrofluoric acid, before washing and acetolysis (9:1 v/v acetic anhydride:sulfuric acid) preceded and followed by acetic anhydride (Heusser and Stock, 1984). Samples were stained with Safranin O and mounted with glycerine gelatin onto slides for counting. At least 300 pollen grains were counted in each sample. LOVECLIM climate model simulations We compare our proxy record with rainfall changes simulated by a transient Dansgaard- Oeschger hindcast experiment (Menviel et al. 2011, Menviel et al. 2014) that was conducted with the Earth System Model of Intermediate Complexity: LOVECLIM (Goosse et al. 2010). The ocean component of LOVECLIM consists of a free-surface primitive equation model with a horizontal resolution of 3° longitude, 3° latitude and 20 depth layers. The atmospheric component is a spectral T21, three-level model based on quasi-geostrophic equations of motion. LOVECLIM also includes a dynamic-thermodynamic sea ice model, a land surface scheme, a dynamic global vegetation model (VECODE) and a marine carbon cycle model (LOCH). This 14 transient simulation uses time-varying orbital, ice sheet and atmospheric CO 2 forcings. As LOVECLIM does not include an interactive ice sheet, freshwater withholding from the ocean during phases of ice sheet growth and freshwater release into the ocean as a result of ice sheet calving and ablation are not explicitly captured. To mimic the time-evolution of these terms and their effect on the oceanic circulation, we apply a freshwater forcing to the North Atlantic region (55°W – 10°W, 50°N –65°N) . The freshwater forcing time series is obtained through an iterative procedure, which optimizes the anomalous freshwater flux such that the simulated temperature anomalies in the eastern subtropical North Atlantic best match the target alkenone-based SST anomalies reconstructed from the Iberian margin core MD01-2443 (Martrat et al. 2007). RESULTS AND DISCUSSIONS Modern-day leaf wax and river water δD and implications for ε wax/w We have constraints on the nature of the δD wax recorder of δD precip from two lines of evidence, stream transported particulate organic matter (POM; this study) and modern plant studies (Feakins and Sessions, 2010; Feakins et al., 2014). Of three stream water samples collected here, two POM samples at 615 and 630 masl yield sufficient C 28 n-alkanoic acid for measurement. The one sample taken at 1630 masl draining conifer forests contains high concentrations of short-chain (<C 22 ) n-alkanoic acid but very low abundance of long-chain (>C 26 ) n-alkanoic acids. This is likely due to the dominance of Pinus species, which produce little n- alkanoic acids >C 24 (Feakins et al., 2014). Thus the highlands probably are not significant contributors of C 28 n-alkanoic acid to Lake Elsinore, making the lake core sedimentary record dominantly a record of mid and low elevations, which is also the largest region of the catchment by area (Fig. 1). Values from lowland elevations (615 and 630 masl) average -151±4.8‰ for δD wax in POM and -71.9±1.1‰ for δD of stream water ( δD w ) (Table 2). Although just two 15 samples from a single runoff event, this provides our best estimate of a representative values for modern day lake inputs. From the measured δD wax and δD w values, we obtain an ε wax/w of - 86±5.3‰ (compound 1σ uncertainty). Today atop Mt. San Jacinto (2540 masl), δD precip ranges from -91‰ in winter months to - 68‰ in summer months, and averages at -84‰ annually (measured across 6 years from 1982 – 1989 A.D.; Friedman et al., 1992). Within the coniferous forest at 1620 masl, δD precip for the winter of 2006 – 2007 A.D. was reported to be -46.7‰ (Feakins et al., 2014), a value that is likely to be enriched above normal given the anomalously dry winter. Within a single runoff event we measure the δD of stream water to be -73.7±0.4‰ at 1630 masl in 27 Jan 2013 (Table 2). However, δD wax in Lake Elsinore probably primarily records δD precip at lower elevations where vegetation contributes more C 28 n-alkanoic acids as discussed. No precipitation measurements are available from lower elevations but the Online Isotopes in Precipitation Calculator (OIPC; Bowen and Wilkinson, 2002) estimates annual δD precip ranging from -78‰ at 1300 (chaparral-forest transition) to -65‰ at 380 masl (lake level) based on expected elevation isotope gradients, close to the estimated altitudinal gradient of -10‰/km for southwest California (Williams and Rodoni, 1997). Our river water samples (-71.9±1.1‰) at ~630 masl integrates δD signal of precipitation received at higher elevations, probably at ~900 masl on average, according to OIPC-derived δD precip gradient. An earlier study on five species of modern plants at 1630 masl in the Lake Elsinore catchment has reported an average ε wax/w of -91±23‰ ( δD wax of -152±22‰; δD w of -67‰) for the C 29 n-alkane (Feakins and Sessions, 2010). In wider region across the coastal southwest US they found an average ε wax/w of -94±21‰ based on 30 species (Feakins and Sessions, 2010). Further north near Santa Barbara, a similar study on C 28 n-alkanoic acids found ε wax/w 16 averages -94±22‰ (based on Quercus agrifolia, the dominant producer of C 28 n-alkanoic acids at the site) (Feakins et al., 2014). Overall, our calculated ε wax/w of -86±5.3‰ representing a catchment-scale average is similar to results obtained from previous studies of modern plants. We therefore consider this measured fractionation when evaluating sedimentary δD wax as a proxy for past δD precip . Corrections for δD wax reconstructions Accounting for vegetation changes To account for the possible influence of changing vegetation types on ε wax/w , we corrected δD wax values using the proportion of graminoids inferred from pollen counts (Fig. 3b). This is based upon the assumption that graminoid pollen proportions relate to graminoid biomass and leaf wax production, which is a simplification, but the best available control. We took a simple two end-member mixing model approach with C 3 graminoid and non-graminoid forms representing two main categories that have the most distinctive ε wax/w (C 3 graminoids: -149±28‰, 1 c.σ, n = 47; non-graminoids: -113±31‰, 1 c.σ, n = 168; compiled by Sachse et al., 2012 from references therein) with a mean difference (ε diff ) of 36‰. Although the compilation is based on C 29 n-alkanes, the very similar ε wax/w values between C 28 n-alkanoic acids (-94‰±22‰; Feakins et al., 2014) and C 29 n-alkanes (-94±21‰; Feakins and Sessions, 2010) in the coastal southwest US lend some confidence to assuming a similar ε diff between the two compounds. We do not consider C 4 graminoids to be significant as they are rare in the modern day coastal southwest US (Still et al., 2003), and despite the drop in pCO 2 , the colder and wetter climate during the glacial period (Kirby et al., 2013) would be unlikely to favor C 4 graminoids (Ehleringer et al., 1997). Further support for this conclusion comes from the δ 13 C values of C 28 n-alkanoic acid together with graminoid pollen from a subset of samples (see Supplementary Text S1). We propose that 17 correction for graminoid pollen accounts for the possible influence of changing vegetation type on the overall fractionation in the catchment, but we acknowledge that other changes in vegetation could also influence the magnitude of the fractionation over time. We calculated a graminoid-corrected δD wax using the formula: Graminoid-corrected δD wax = δD wax + [f graminoid × ε diff ] (eq. 2) The graminoid-corrected δD wax values represent the scenario where leaf waxes were contributed solely from non-graminoid plants. This allows a better estimation of variations of δD precip without influence from graminoid contributions. Uncertainties of this approach come from ε wax/w and f graminoid , given the small number of calibration studies on n-alkanoic acids and the likely disproportional production and dispersal of pollen compared to leaf waxes. Accounting for ice volume changes During the Last Glacial Maximum (LGM), the δ 18 O of global surface ocean water was 1‰ heavier due to ice volume changes (Schrag et al., 1996). We account for the consequent changes in δD of sea water (Δ δD sw ) using the LR04 benthic δ 18 O stack (Lisiecki and Raymo, 2005) following Tierney et al. (2011). First scaling the record to obtain an 8‰ Δ δD sw between LGM and the present and then translating the corresponding changes in δD along the global meteoric water line (Craig, 1961; Fig. 3c). Ice volume-corrected δD wax were then calculated using this equation: Ice volume-corrected δD wax = [( δD wax + 1) / (Δ δD sw + 1)] – 1 (eq. 3) Combined corrections for graminoid and ice volume changes amount to adjustments of 3‰ on average and no greater than 14‰ (Fig. 3d). These corrections are minimal in magnitude and pattern relative to the large fluctuations in the δD wax record, demonstrating that this record is a 18 robust indicator of hydrological changes. Hereafter discussions of δD wax refer to the ice volume- and graminoid-corrected δD wax values. Lake Elsinore past δD wax , δD precip and comparison with modern-day hydroclimate Lake Elsinore δD wax varies between -211‰ and -90‰ from 32 – 9 cal. kyr BP (Fig. 4a). The late glacial (32 – 15 cal. kyr BP) exhibits generally lower δD wax values (although with ~70‰ fluctuations) transitioning to higher δD wax values in early Holocene following the global trend of deglaciation. Using ε wax/w of -86±5.3‰ we estimate glacial δD precip may have oscillated between approximately -125‰ and -55‰. Comparison with modern storm-track δD precip implies that during most of the glacial, coastal southwest US received a greater proportion of northerly moisture than today (or precipitation at colder temperatures or with greater antecedent rain-out). Times with δD precip values resembling that of today, occurred in briefer intervals and may indicate a similar to modern moisture source, or more subtropical and tropical moisture sources but at a slightly cooler temperature. We note that δD precip values during the B-A and YD appear to be similar to today, whereas during the early Holocene, D-enriched precipitation implies more tropical moisture sources. Some of these early Holocene values imply a precipitation of 0‰, which would only be realistic if conditions were very arid resulting in evaporation of raindrops during descent (Williams and Rodoni, 1997) and leaf water leading to smaller ε wax/w (Kahmen et al., 2013a,b). This interpretation is consistent with low sand % values at this time (Fig. 4), evidence for low runoff under an arid climate in the early Holocene. The magnitude of positive δD precip shift during the deglaciation (~85‰) is the most pronounced of the entire record. Warming during deglaciation may account for part of that isotopic shift by increasing the temperature of condensation. Since quantification of deglacial 19 warming in coastal southwest US from proxy reconstructions is lacking, we turn to model simulation (Menviel et al., 2011), which estimates a ~2°C deglacial warming of air temperature (Fig. 7h). Assuming the mean condensation temperature varies with mean surface temperature, a 2°C warming would account for about +12‰ shift of δD precip (Dansgaard, 1964). The remaining +73‰ isotopic shift is likely independent of temperature and relates to storm track shifting from more northerly to more tropical North Pacific sources, possibly added to by lower precipitation totals (the amount effect i.e., less antecedent precipitation). The 70‰ magnitude of δD precip shifts during the glacial and deglaciation in the coastal southwest US is similar to modern synoptic scale variations in moisture sources: 80‰ is the approximate difference between northern and tropical North Pacific-sourced storms in southern Sierra Nevada (McCabe-Glynn et al., 2013). If the modern synoptic range of precipitation isotopes may be taken as analogue for those during the glacial then this would suggest that millennial modes stuck near the extremes of modern synoptic storms. D-depleted precipitation modes may reflect persistent influx of northerly North Pacific moistures sources, perhaps compounded with isotope effects from greater precipitation amounts and colder condensation temperatures. D-enriched precipitation modes occurred in short intervals during glacial times, with precipitation isotopes similar to modern long term averages. Given the sampling resolution, sedimentation rates and multi-decadal averaging we cannot comment on the synoptic range of variability in precipitation isotopic compositions over the coastal southwest US during past times. Lake Elsinore paleoclimate record and implications to atmospheric circulations Lake Elsinore δD wax values show a significant positive shift from the late glacial towards early Holocene, with significant fluctuations occurring within the late glacial. Grain size analysis 20 reveals sand ranges between 0 – 15% for most of the core, which we interpret as a proxy for more voluminous, persistent and/or more intense runoff, following Kirby et al. (2013). A notable exception is between 25.6 – 27.6 cal. kyr BP (23.3 – 25.5 m) where sand % averages at ~30% and exceeds 70% at ca. 26.3 cal. ka BP (Fig. 4b, yellow bar). We interpret this high sand unit as representing a progradation of the sandy littoral zone during a prolonged dry period. Otherwise with expanded lake dimensions in wet times it would be difficult to produce high sand concentrations so far from the inlet (see Supplementary Text S2). Total carbonate (TC%) stays low through the glacial and increases abruptly after 15 cal. ka BP reaching >15% during the early Holocene, suggesting chemical precipitation of carbonate favored by lower lake level and/or warmer epilimnion temperatures (Kirby et al., 2013). The +80‰ shift in δD wax values and drop of 10% sand concentration during the deglaciation indicates a fundamental change in hydroclimate. This can be explained by a southward displacement of the polar jet stream, as it steered around the topographic obstruction of the LIS, resulting in what has long been referred to as the ‘dipping westerlies’ (Antevs, 1948). This mode of circulation is expected to dominate during the glacial, and we find isotopic evidence that a northern North Pacific moisture source dominates the δD wax record from Lake Elsinore across the glacial. During deglaciation, the retreat of the LIS would cause a northward migration of the polar jet stream allowing more tropical and subtropical moisture to reach coastal southwest US, accompanied by progressively drier climate. We do not see a significant response of Lake Elsinore δD wax during the YD (ca. 12.8 – 11.8 cal. kyr BP). However, at the start of the YD, sand % drops to low levels, indicating a drying that persists throughout the Holocene, with no evidence for a perturbation at the end of the YD. 21 Although the late glacial (32.2 – 15 cal. kyr BP) is generally more D-depleted compared to early Holocene, large-scale abrupt fluctuations of δD wax varying between -211‰ and -134‰ was experienced. More negative δD wax occurred during ca. 32.2 – 30.7, 27.4 – 24.4, 22.6 – 20.6, and 19.5 – 15.3 cal. kyr BP, indicating more delivery of northern Pacific moisture to the coastal southwest US. Among these four low δD wax intervals, 32.2 – 30.7 cal. kyr BP and 27.4 – 24.4 kyr BP coincides with the time of reduced AMOC intensity, increased IRD and cooling of the North Atlantic associated with H3 and H2 respectively (Martrat et al., 2007; Lippold et al., 2009, Hodell et al., 2010) (Fig. 5). From 22.6 – 20.6 cal. kyr BP, no apparent AMOC slowdown or peak IRD deposition is evident (Lippold et al., 2009; Hodell et al., 2010); however, we note a slight drop in Greenland NGRIP ice core δ 18 O over this interval (Rasmussen et al., 2008) as well as a blip of cooler North Atlantic SST at ca. 22 cal. ka BP (Martrat et al., 2007) (Fig. 5). The low δD wax from 19.5 – 15.3 cal. kyr BP was the longest low δD wax excursion among all. The onset of this interval precedes that of H1 (18 cal. ka BP), and the ending of this interval coincides with the termination of H1 (15 cal. ka BP) (McManus et al., 2004; Hodell et al., 2010). However, we do note that the onset time of 19.5 cal. ka BP coincides with a minor slowdown of AMOC and a gradual drop in North Atlantic SST (McManus et al., 2004; Lippold et al., 2009) (Fig. 5). Overall, all four low δD wax intervals recorded in Lake Elsinore are associated with cooler high-latitude North Atlantic, with the first two clearly correlated with H3 and H2. In the modern southwest US, precipitation associated with southward displacement of the polar jet stream during winter dominates the region’s hydroclimate (Cayan and Peterson, 1989). A colder high-latitude North Atlantic as well as extended Arctic sea ice would strengthen the winter season Aleutian low resulting in lower North Pacific SSTs as well as southward displaced and strengthened polar jet stream (Negrini, 2002; Francis et al., 2009). A similar atmospheric 22 teleconnection between the Atlantic and Pacific is also observed in climate model experiments which apply freshwater forcing to the North Atlantic leading to an AMOC weakening and a cooling of the North Atlantic (Okumura et al., 2009). Therefore, we infer that lower δD wax values in Lake Elsinore originate from cooler conditions in the North Atlantic, which strengthens the winter season Aleutian low and displace the jet stream southward. This in turn favors delivery of northern Pacific moisture to the coastal southwest US. We interpret higher δD wax values as derived from enhanced transport of tropical and subtropical North Pacific moisture, perhaps including delivery in the form of so-called Atmospheric Rivers (ARs) that extend northeastward from near Hawaii bringing tropical moisture and heavy precipitation to the California coast (Neiman et al., 2008; Dettinger, 2011). AR events bring more D-enriched precipitation compared to northerly Pacific storms (Spry et al., 2014), although within the storm very D-depleted values have been recorded associated with changing condensation height and temperature (Coplen et al., 2008). Today, ARs happen in winter when a high pressure cell in the Gulf of Alaska blocks the polar jet and splits it into two. The southern branch then brings a low latitude moisture plume to California. During times of weaker Aleutian low (as opposed to strengthened Aleutian low when high-latitude North Atlantic is colder), there could be more frequent blocking high over Alaska favoring the formation of ARs. Between 32.2 and 20 cal. kyr BP, times of lower δD wax are associated with less persistent and/or intense runoff (lower sand % and the anomalous high sand unit). Conversely, there is a good positive correlation between higher δD wax and inferred runoff, suggesting that more tropical and subtropical moisture brought by southwesterly storms results in more intense and/or persistent precipitation during the late glacial. This implies that ARs, which today bring most of the heavy rainfall to California, and produce most of the major floods (Neiman et al., 2008; 23 Dettinger, 2011), also had an important role during late glacial. In the Holocene, changes in the frequency of ARs have been proposed also to be important to establishment and persistence of pluvials in the coastal southwest US (Kirby et al., 2012). Therefore, we interpret intervals of higher δD wax during the late glacial as indicating periods of more frequent ARs and a wetter glacial climate in the coastal southwest US, sometimes lasting several millennia. When the AMOC slows down and/or the northern North Atlantic cools, less frequent ARs (lower δD wax ) result in intervals of relatively drier climate in the otherwise wetter-than-Holocene late glacial. Starting from 20 cal. ka BP, sand % increases to ~10% until 15 cal. ka BP. The higher sand % from 20 – 15 cal. kyr BP is in contrast to the lower δD wax . However, when we compare the more detailed high-frequency fluctuations between δD wax and sand %, we find a positive correlation between the two variables – higher δD wax , higher sand % – as observed elsewhere in the record. We infer that the low δD wax and high sand content observed between 20 – 15 cal. Kyr BP likely reflects a wetter background climate that could be due to a warming of the Northeastern Pacific Ocean (Menviel et al., 2011). The fact that the magnitudes of δD wax oscillations are remarkably similar across the late glacial, despite their different nature, duration and scale of the associated North Atlantic cold events, is intriguing. This suggests either that there are only two possible modes of atmospheric circulation during the late glacial, or that some independent oscillation of regional climate is at play. In either case, the scale and abruptness of the isotopic swings depict a highly sensitive switching of hydroclimate regime in the region. Regional comparisons Offshore coastal southwest US in the Santa Barbara Basin (SBB), Mg/Ca SST reconstructions (Pak et al., 2012) show lower (~3°C) SSTs owing to strengthened California 24 Current during times of more negative δD wax in Lake Elsinore, except for during 28 – 25 cal. kyr BP where a hiatus in the SBB record hinders comparison (Fig. 6b). If the majority of the moisture delivered to coastal southwest US came from the coastal ocean, a lower temperature of evaporation would result in a more depleted δD of vapor, and hence lower δD precip . However, δD precip in west coast of US is probably dominantly sourced from and controlled by conditions in the open North Pacific Ocean. The lack of correlation between coastal SSTs and speleothem δ 18 O at Crystal Cave in southern Sierra Nevada over the past century supports this argument (McCabe-Glynn et al., 2013). Therefore, we preclude coastal SSTs as a factor controlling δD wax in Lake Elsinore. Instead, it is more likely that the concurrent shifts in moisture sources at Lake Elsinore and Santa Barbara Basin SST (strength of California Current) reflect large scale changes in North Pacific climate. To put Lake Elsinore δD wax record under a regional hydroclimatic context, we compare our record with published speleothem records in a broader region of the western US (Fig. 6). In coastal southwest US, Moaning Cave on the western slope of central Sierra Nevada shows a transition from more depleted δ 18 O during H1 to more enriched δ 18 O during B-A consistent with Lake Elsinore δD wax , although the record only starts at ca. 16 cal. ka BP (Fig. 6c; Oster et al., 2009). Unlike at Lake Elsinore, there is a return to glacial-like conditions during the YD at Moaning Cave (Oster et al., 2009), perhaps because of a greater temperature drop at the high elevation Sierran site, or because of its more northerly position, suggesting only a modest southward shift of the westerly storm track during the YD. The Leviathan Chronology speleothem record (Fig. 6d) in the Great Basin reflects shifting moisture sources as well as temperature changes during the last glacial (Lachniet et al., 2014). More negative δ 18 O values are recorded from 20.1 – 15.6 cal. kyr BP with no apparent major transition at the onset of H1 25 (Lachniet et al., 2011). The persistent D-depleted isotopes during this interval from Lake Elsinore are entirely consistent with this record and scenario, suggesting that the precipitation pattern was common to the Great Basin and coastal southwest US. A speleothem δ 18 O record from Fort Stanton, New Mexico provides an opportunity for coastal-to-interior southwest US comparison from 9 – 32 cal. kyr BP without hiatus (Fig. 6e). During times of negative δD wax excursions at Lake Elsinore, Fort Stanton documents more depleted δ 18 O values (Asmerom et al., 2010) analogous to Lake Elsinore, although with minor temporal offsets probably due to age uncertainties, a notable shorter duration for H2, and no evidence for a prior-to-H1 negative shift. In contrast, Cave of Bells in Arizona shows rather invariant δ 18 O values between 32 – 15 cal. kyr BP (Fig. 6f; Wagner et al., 2010). During the YD, both Fort Stanton and Cave of Bells display a negative isotopic shift not evident in Lake Elsinore. The differences between Lake Elsinore and Fort Stanton have been attributed to the penetration of the North American Monsoon (NAM) into New Mexico and Arizona but not California (Kirby et al., 2013). This leads us to propose that the LIS retreat and global warmth over-rode the AMOC influences on Pacific storm tracks leading to a weaker influence of the YD at the coastal southwest US compared to interior southwest US where the Atlantic disruption is felt through the NAM. Extension of the Lake Elsinore record here allows for a longer comparison of the Pacific- only and Pacific/NAM influenced records. Unfortunately the comparisons are not conclusive as the two Pacific/NAM influenced records do not agree (Fort Stanton shows much larger oscillations than Cave of Bells) such that none of the records appear to be very similar across this interval. It is notable that the Lake Elsinore record displays the largest shifts of any of these records both within glacial abrupt transitions and deglaciation documenting the sensitivity of the coastal southwest hydroclimate. 26 Climate model comparisons A transient simulation of the period 35 – 10 cal. kyr BP, including millennial-scale variability, is performed with the Earth System model LOVECLIM. In order to present the millennial-scale cooling events recorded in Greenland ice core (Dansgaard et al., 1993, Huber et al., 2006) and Iberian margin marine sediment cores (Martrat et al., 2007), the AMOC was weakened through the addition of freshwater into the North Atlantic. In general, the simulated SST off the Iberian margin compare well with the reconstructed SST variations from marine sediment cores (Martrat et al., 2007; Menviel et al., 2014), although a significant drop in modeled SST at ca. 27.8 cal. ka BP is an overestimation of AMOC slowdown (Fig. 7d). In LOVECLIM, an AMOC weakening leads to a deepening and southward displacement of the Aleutian low, consistent with our interpretation of more negative δD wax at Lake Elsinore reflecting northern North Pacific moisture. A shutdown of the AMOC leads to stronger precipitation over the west coast of North America, with a simulated precipitation increase over coastal southwest US of +8% during H3 and H2 (as well as 27.8 cal. ka BP, see comment above), and a larger 30% increase during H1 (Fig. 7c,d). While we do not find any evidence for wetter conditions in the sand record across H2 and H3, wetter conditions are very much apparent for the larger magnitude event, H1, with strong increase in the sand size present in the core. We do find that the isotopic shift occurs earlier in the core than the AMOC shutdown, and thus we find support for the idea that the growth phase of the LIS as it approached it’s maximum extent was also associated with wetter conditions across the western US (Lachniet et al., 2011). In the model simulation, the large precipitation increase at H1 is due to a very weak AMOC (perhaps overestimated) as well as formation of NPDW, (Fig 7g). NPDW formation during that period is also documented by young radiocarbon ages of marine sediment in the 27 northwest (Okazaki et al., 2010) and northeast (Rae et al., 2014) Pacific Ocean. Model simulations suggest that NPDW formation leads to a strengthening and northward shift of the Kuroshio Current, hence warming the northeastern Pacific (Menviel et al., 2011; Okazaki et al., 2010), in agreement with some proxy records (de Vernal and Pedersen, 1997; Sarnthein et al., 2006 ; Okazaki et al., 2010). A warmer northeastern Pacific could promote evaporation and hence the moisture flux carried by westerlies towards western US, resulting in wetter climate, consistent with higher sand % in Lake Elsinore as well as expansion of the Great Basin pluvial lakes (Oviatt, 1997, Adams & Wesnousky, 1998; Benson et al., 2011; Lyle et al., 2012). Our proxy-model comparison therefore provides strong evidence for wet conditions during H1 and a clear mechanism for delivery of north Pacific moisture in the model dynamics, further corroborated by our leaf wax isotopic record from Lake Elsinore. CONCLUSIONS We develop a leaf wax δD record spanning 9 – 32 cal. kyr BP from Lake Elsinore in the coastal southwest US. After accounting for influences of changing ice volume and vegetation type on δD wax , we find abrupt, large magnitude fluctuations both within the late glacial (>65‰) and during deglaciation (>80‰), which we infer reflects variations in δD precip associated with changing Pacific storm trajectories. More negative (<−175‰) δD wax values are recorded on average during the glacial (relatively to the Holocene) and persistently during intervals that span 32.2 – 30.7, 27.4 – 24.4, 22.6 – 20.6, and 19.5 – 15.3 cal. kyr BP. Each of these intervals is interpreted as receiving dominant storm tracks from the northern Pacific. These negative δD wax excursions appear to coincide with North Atlantic cold events. This inference is consistent with LOVECLIM 28 simulations, which find that an AMOC weakening deepens the Aleutian Low and leads to a southward displacement of the polar jet stream in the North Pacific. The negative δD wax excursion from 19.5 – 15.3 cal. kyr BP is coeval with H1, a time of week AMOC as well as enhanced NPDW. Formation of NPDW during H1 associated with higher levels of atmospheric CO 2 and stronger high northern latitude insolation leads to an SST increase in the northeastern Pacific, promoting evaporation and moisture influx to the western US. LOVECLIM not only predicts formation of NPDW but also a large and persistent increase in wetness throughout H1. This model result is consistent with the sand evidence for high runoff and isotopic evidence for atmospheric circulation being stuck in a mode that delivers abundant northern Pacific moisture to the coastal southwest US, probably via a southward shift of the dominant storm track. Late glacial millennial-scale events (29.9 – 27.7, 23.9 – 22.8, and 20.4 – 19.9 cal. kyr BP) are associated with higher δD wax values (> −165‰), which represent contributions from tropical and subtropical Pacific sources. Today subtropical moisture sources deliver brief but intense precipitation events (so-called Atmospheric Rivers) and this could also be the mode of delivery during these intervals. We reconstruct more positive δD wax values in the early Holocene suggesting warmer tropical or subtropical sources as well as arid conditions in the region. The largest δD wax shift (>80‰) starts from 15 cal. ka BP and coincides with the global warming and LIS retreat of the deglaciation, indicating fundamental alteration from glacial hydroclimate regime dominated by northerly sources, towards a more tropical and subtropical Pacific moisture influx during the Holocene. We propose that storm trajectories persisted for millennia in either of two distinct modes during the glacial, with abrupt transitions between modes, apparently triggered by AMOC or 29 sometimes NPDW transitions. 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(c) Profile of elevation and precipitation across a transect from Lake Elsinore to San Jacinto Peak, with distribution of vegetation types and elevations of river sampling (blue diamonds). 43 Figure 2. Age model for Lake Elsinore sediment core produced by BACON 2.2 program. Green markers show the calendar ages with probability distribution based on inputs of radiocarbon dates with 2σ range. Shading represents Monte Carlo iterations of age models. The grey dotted lines encompass 95 percentile of the iterations. The red curve represents the mean values of the age model, which is applied to this core. Two OSL age estimates (orange open circle with 2σ error bar) are taken within the high sand unit to provide additional information. This age model extends that of Kirby et al. (2013). 44 Figure 3. (a) Corrections of Lake Elsinore δD wax values for vegetation type and ice volume changes (black: uncorrected δD wax ; green: graminoid-corrected δD wax ; red: graminoid- and ice volume-corrected δD wax ). (b) Lake Elsinore graminoid pollen percentage used to correct for vegetation type. (c) Change in sea water δD (Δ δD sw ) due to changing ice volume inferred from LR04 benthic stack (Lisiecki and Raymo, 2005). The y-axis is scaled to x5 magnification to that of fig. 3a for easy comparison. (d) The change in δD wax after graminoid and ice volume correction. The y-axis is scaled to x5 magnification to that of fig. 3a for easy comparison. 45 Figure 4. Proxy reconstructions from Lake Elsinore. (a) δD wax with uncertainties (black). (b) Weight percentage of sand (sand %). Orange shows all sand % data plotted in log scale to accommodate the high sand unit. Black shows 20-point lowpass-filtered sand % in linear scale with high sand unit removed for easier comparison with δD wax . (c) Weight percentage of total carbonate (TC%). (d) Weight percentage of total organic matter (TOM%). The yellow bar highlights the high sand unit (see Supplementary Text S2). 46 Figure 5. Comparison of Lake Elsinore δD wax to North Atlantic proxy records. (a) Lake Elsinore δD wax record (this study). Blue bars: times of negative δD wax excursions recorded in Lake Elsinore. (b) Pa/Th record from Bermuda Rise indicating variability of AMOC intensity (pink: McManus et al., 2004; purple: Lippold et al., 2009). H3 – 1 and YD are labeled where AMOC slows down (Bermuda Rise Pa/Th ratio increases) and North Atlantic IRD increases. (c) Record of North Atlantic ice rafted debris (IRD) with age model synchronized with Greenland ice core (Hodell et al., 2010). (d) Iberian Margin SST reconstructions based on alkenone paleothermometry (Martrat et al., 2007). (e) Greenland NGRIP ice core δ 18 O record (NGRIP members, 2004) on GICC05 chronology (Rasmussen et al., 2008). Grey curve is the original data and black line shows the 20-point low-pass filtered data. Dansgaard-Oeschger interstadials (Dansgaard et al., 1993) are labeled in numbers. 47 Figure 6. Comparison of precipitation isotope proxy archives across the western United States. (a) Lake Elsinore δD wax record (this study). Blue bars: times of negative δD wax excursions recorded in Lake Elsinore. (b) Santa Barbara Basin sediment Mg/Ca SST record (Pak et al., 2012). (c) Moaning Cave speleothem δ 18 O record (Oster et al., 2009). (d) Great Basin composite speleothem δ 18 O record ‘Leviathan Chronology’ (Lachniet et al., 2014) comprising data from Leviathan Cave (light green) and Pinnacle Cave (dark green). (e) New Mexico Fort Stanton speleothem δ 18 O record (Asmerom et al., 2010). (f) Arizona Cave of Bells speleothem δ 18 O record (Wagner et al., 2010). 48 Figure 7. Proxy and model insights into forcings. (a) Lake Elsinore δD wax record. (b) Lake Elsinore sand %. (c) Modeled precipitation amount over Lake Elsinore. (d) Proxy-recorded (blue; McManus et al., 2004; Lippold et al., 2009) and model simulated (purple) variations in AMOC strength. (e) North Atlantic IRD record (Hodell et al., 2010). (f) Proxy-recorded (Martrat et al., 2007; dark blue) and model-simulated (light green) Iberian Margin SST. (g) Modeled NPDW formation. (h) Modeled air temperature over Lake Elsinore. 49 Table 1. Radiocarbon age control for Lake Elsinore core. Depth (cm) Laboratory ID# Material δ 13 C (‰) 14 C age (yr BP) 14 C age used in BACON σ Cal. yr BP (avg.) Cal. yr BP (2σ min) Cal. yr BP (2σ max) 1274.5 a N93630 Gastropods -25 8655 8655 35 9613 9541 9684 1274.5 a N93631 Gastropods -25 8710 8710 35 9664 9548 9780 1397 a N94679 d Bulk -25 10155 9450 46 11858 11685 12030 1509 a N94680 d Bulk -22.9 10950 10240 46 12806 12648 12964 1541 a N94681 d Bulk -25 11650 10940 46 13513 13334 13691 1619 a N94682 d Bulk -24.6 12200 11490 46 14045 13887 14202 1684 a N95444 Mixed discrete -25 12140 12140 280 14252 13437 15067 1710 a N95445 Mixed discrete -25 12460 12460 120 14591 14091 15090 1724 a 134836 Mixed discrete -25 12190 12190 290 14313 13470 15155 1738 a N95446 Mixed discrete -25 13420 13420 230 16161 15390 16932 1747.5 a N94683 c Bulk -26.4 13380 n.a. 40 n.a. n.a. n.a. 1747.5 a N94003 Charcoal -25 13260 13260 35 16139 15578 16699 1778.5 a N94004 Wood -25 13775 13775 35 16892 16734 17049 1823.5 a N94684 c Bulk -25 15250 n.a. 60 n.a. n.a. n.a. 1823.5 a N94005 Charcoal -25 14360 14360 30 17472 17154 17790 1823.5 a N94243 Charcoal -25 14310 14310 30 17394 17082 17706 1871 b 134837 Charcoal; Charred grass -25 14740 14740 200 17942 17460 18424 1998 a N94006 Seeds -25 16580 16580 40 19699 19461 19936 2019.5 a N94685 c Bulk -22.5 17490 n.a. 70 n.a. n.a. n.a. 2019.5 a N94007 Seeds -25 16880 16880 40 20077 19832 20321 2081.5 a N94008 Wood -25 17980 17980 180 21520 20911 22128 2198.5 a N94010 Seeds -25 19630 19630 40 23457 23134 23779 2281 b 134839 Small twig -23.4 20870 20870 90 25199 24881 25516 2292.5 b 134840 Charcoal -24.1 21370 21370 90 25700 25510 25899 2344.5 a N94245 Charcoal -25 21035 21035 40 25248 24940 25556 2344.5 a N94011 Charcoal -25 21120 21120 70 25164 24834 25494 2385 b 118908 Charcoal -25 22010 22010 80 26460 26078 26843 2426 b 134841 Small charcoal -23.9 21760 21760 210 26061 25650 26481 2831 b 118909 Charcoal -25 25940 25940 110 30720 30419 31016 2860.5 a 150331 Wood -25 26550 26550 90 31116 30970 31262 2860.5 a 150337 Wood -25 26270 26270 80 30969 30758 31179 a Identification number from measurements at Lawrence Livermore National Laboratory (LLNL) b Identification number from measurements at University of California Irvine (UCI) c Bulk samples used to calculate reservoir effect 14 C age difference d Bulk 14 C ages corrected for reservoir effect 50 Table 2. River suspended sediment samples from Lake Elsinore catchment. Sample Latitude Longitude Elevation (m) δD wax (‰) σ (‰) δD w (‰) σ (‰) δ 18 O w (‰) σ (‰) ε wax/w (‰) c. σ a (‰) SJR 1 33.736 -116.824 615 -152 2.3 -71.6 0.5 -10.2 0.2 SJR 2 33.736 -116.816 630 -151 4.2 -72.2 1 -10.2 0.2 Mean -151 4.8 -71.9 1.1 -10.2 0.3 -86 5.3 SJR 3 33.741 -116.718 1630 n.a. n.a. -73.7 0.4 -11.9 0.1 a Compound 1σ uncertainty 51 SUPPLEMENTARY MATERIALS Leaf wax δ 13 C constraints on sources of leaf wax biomarkers To test the sources of the leaf wax, we analyzed the carbon isotopic composition of C 28 n- alkanoic acid ( δ 13 C wax ). δ 13 C wax values were measured for 17 samples distributed across the record and concentrated around high and low δD wax values, to assess whether the δD wax record is robust to vegetation changes captured by δ 13 C wax . Compound-specific δ 13 C values of n-alkanoic acids as methyl esters were measured by same instrument and procedures similar to measuring δD wax , but via the GC Isolink combustion oven at 1000 °C. Measured δ 13 C values are normalized to the Vienna Pee Dee Belemnite (VPDB) isotopic scale using external standard containing a mixture of 8 fatty acid methyl esters ( δ 13 C values ranging from -30.9 to -23.2‰) and a mixture of 15 n-alkanes ( δ 13 C values ranging from -33.3 to -28.6‰) obtained from A. Schimmelmann, University of Indiana. RMS errors of replicate measurements of the standards are 0.34‰ and 0.14‰ respectively. δ 13 C values were then corrected for the added methyl group by mass balance ( δ 13 C of methanol = -25.45±0.37‰ based on offline determinations). The δ 13 C wax values in this subset of samples range from -28 to -32‰ (Fig. S1; Table S1). These values fall within the variations of terrestrial C 3 plants (Kristen et al., 2010; Garcin et al., 2014), suggesting that the catchment has been dominated by C 3 vegetation. However, given the measured fluctuations of up to 4‰, there is still the possibility of a 13 C-enrichment due to C 4 graminoid contributions. To test this possibility, we compare δ 13 C wax with graminoid pollen abundance (Fig. S1). The comparison shows that periods of more graminoid abundance do not coincide with relatively higher δ 13 C wax values, suggesting that the graminoids are primarily of the C 3 type. Therefore, for the purpose of graminoid-correction of δD wax discussed in the main text, 52 we find it safe to assume graminoids are C 3 . Another possibility that may cause relatively enriched δ 13 C wax values is increased contribution from aquatic macrophytes (Chikaraishi and Naraoka, 2003; Aichner et al., 2010), especially emergent plants which may produce some C 28 n- alkanoic acids (Ficken et al., 2000; Gao et al., 2011). To explore this possibility we look into evidence of increased lake productivity indicated by percentage of total organic matter (TOM%) (Fig. 4d). However, we find no significant correlation between δ 13 C wax and TOM% or aquatic macrophyte pollen (Heusser, unpublished data) suggesting that aquatic inputs were not the drivers of higher δ 13 C wax values. We therefore suggest that δ 13 C wax values more likely reflect changes within the terrestrial C 3 domain or perhaps appearance of C 4 woody shrubs. For example, C 3 trees and shrubs tend to have slightly higher δ 13 C (bulk and long-chain n-alkanes) on average relative to C 3 forbs and graminoids (Jessup et al., 2003; Garcin et al., 2014). We also note that other effects such as changes in water availability and water use efficiency may account for variability in δ 13 C of >4‰ (Liu et al., 2005; Hou et al., 2007). Although we cannot identify exactly what ecological changes were influencing δ 13 C wax values, we suggest that the δD wax record is robust to these changes and most likely reflect hydrological changes. This is further supported by the lack of correlation between δD wax and δ 13 C wax (Fig. S2). High sand unit The high sand unit (Fig. 4b, yellow bar) extends from 2330 – 2550 m depth and 25.6 – 27.6 cal. kyr BP. There are three possibilities for the unusual sand unit’s origin: mass wasting event following a catastrophic event such as an earthquake-induced landslide, extreme wetness and run-off, or the progradation of the northwest lake edge sandy littoral zone during a prolonged dry interval. 53 To test the first scenario, we examined radiocarbon ages above and within the upper part of the sand layer, and collected two samples within the sand layer for optically stimulated luminescence (OSL) dating (Fig. 2, open circle; details below). Radiocarbon ages show continuous sedimentation and no significant change in sedimentation rate. The two OSL ages, although slightly offset from radiocarbon ages, are in chronological order suggesting deposition over time. We also have examined the core’s physical attributes and found no evidence of erosive features (e.g. basal scouring), normally graded grain size distribution, or features associated with rapid sediment setting. Therefore, we reject the first scenario and conclude that the unusual sand unit represents deposition over time. Second, we reject the high sand – high runoff model by Kirby et al. (2013) for this interval. Unlike the rest of the core where sand averages 5 – 10 % and is characterized as disseminated very fine sand in a clayey silt matrix, the unusual sand unit averages 30 % - and up to 70 % - comprising a significant proportion of fine to coarse sand. The sand concentration and size classes are too high and too large to invoke greater background sand mobilization alone; hence it is unlikely simply caused by higher runoff particularly because the core site is far from the lake’s only major inlet. Furthermore, as higher runoff implies higher lake level, it is difficult to envision a scenario where sand (including coarse sand) contributing >30 % of total sediment is transported across a gentle sloping lake floor into the far western basin where core LEDC10-1 was extracted. As a third consideration, the unusual sand unit may represent the progradation of the northern sandy littoral zone during a prolonged dry interval. Supporting this contention is a ~2.4 m thick sand unit, located ~22.6 – 25 m below the lake floor, discovered from the northwestern edge of Lake Elsinore using a truck mounted cone penetration test (CPT). This CPT was taken 54 from the modern dry shore in 2009 adjacent to core LESS02-10 (Kirby et al., 2005). Since this site is well-located to record changes in the position and extent of the littoral zone, we infer that the sand unit in core LEDC10-1 (23.3 – 25.5 m) represents the leading edge of the sandy littoral zone during a prolonged dry interval. However, unlike today when prolonged drought causes Lake Elsinore to desiccate, lower evaporation (Maher et al., 2014), higher water tables, and snow melt during the glacial sustained the shallow lake without desiccation. Based on rejection to the first two scenarios, and the evidence supporting this scenario, we conclude that the high sand unit is most likely representing progradation of the littoral zone indicative of prolonged dry period within the otherwise wetter-than-Holocene glacial period. OSL dating methods Luminescence dating was undertaken using the IRSL (Infra-Red Stimulated Luminescence) signal in 175–200 µm K-feldspar grains carefully extracted from intact sandy zones towards the base of the core. Samples were extracted from the core and prepared under laboratory lighting conditions at UCLA. Samples were wet sieved, treated with dilute HCl, and grains denser than 2.58 gcm -3 removed using a lithium metatungstate solution. After rinsing, grains were treated in dilute hydrofluoric acid for around 10 minutes to remove their outer surfaces, rinsed, dried and sieved again to remove material <175 µm. Samples were mounted in Risø single grain holders, and measured within an automated Risø TL-DA-20D luminescence reader, incorporating an EMI 9235QB photomultiplier, fitted with a BG3 and NG39 filter combination, to detect blue luminescence emissions. A single grain post-IR IRSL single aliquot regenerative-dose (SAR) protocol was used, modified for single grain application from Buylaert et al. (2009). This protocol has been tested successfully using known age samples, many from Southern California, with results and further technical details presented by Rhodes (in revision). 55 This approach incorporates a preheat of 10s at 250°C, with IRSL measurements at 50 and 225°C using a filtered IR laser, and a hot IRSL bleach for 40s using IR diodes at 290°C at the end of each SAR cycle. Multiple cycles of different regenerative doses, a zero dose, and a repeat of the first laboratory dose were administered to construct a growth curve. Age estimation used calculated sediment dose rates from measured K (ICP-OES), U and Th (ICP-MS) concentrations, present overburden depth for cosmic dose rate, an assumed 12.5% internal K concentration, and saturated (25%) water content. Fading tests were conducted using conventional multiple grain aliquots over several months. The upper sample displayed no fading, while the lower sample showed a little fading with a g-value of around 3 %/decade. The fading-corrected value is a little older than for the upper sample, while the uncorrected value would be younger (24,400 years). The IRSL age estimates (one with no fading observed, one corrected) are a little higher than the radiocarbon age estimates on either side. Supplementary references Aichner, B., Herzschuh, U., Wilkes, H., 2010, Influence of aquatic macrophytes on the stable carbon isotopic signatures of sedimentary organic matter in lakes on the Tibetan Plateau: Organic Geochemistry, v. 41, p. 706–718. Buylaert, J.P., Murray, A.S., Thompsen, K.J., Jain, M., 2009, Testing the potential of an elevated temperature IRSL signal from K-feldspar: Radiation Measurements, v. 44, p. 560–565. Chikaraishi, Y., Naraoka, H., 2003, Compound-specific δD–δ13C analyses of n-alkanes extracted from terrestrial and aquatic plants: Phytochemistry, v. 63, p. 361–371. 56 Ficken, K.J., Li, B., Swain, D.L., Eglinton, G., 2000, An n-alkane proxy for the sedimentary input of submerged/floating freshwater aquatic macrophytes: Organic Geochemistry, v. 31, p. 745–749. Gao, L., Hou, J., Toney, J., MacDonald, D., Huang, Y., 2011, Mathematical modeling of the aquatic macrophyte inputs of mid-chain n-alkyl lipids to lake sediments: Implications for interpreting compound specific hydrogen isotopic records: Geochimica et Cosmochimica Acta, v. 75, p. 3781–3791. Garcin, Y., Schefuß, E., Schwab, V.F., Garreta, V., Gleixner, G., Vincens, A., Todou, G., Séné, O., Onana, J.-M., Achoundong, G., Sachse, D., 2014, Reconstructing C3 and C4 vegetation cover using n-alkane carbon isotope ratios in recent lake sediments from Cameroon, Western Central Africa: Geochimica et Cosmochimica Acta, v. 142, p. 482– 500. Hou, J., D’Andrea, W.J., MacDonald, D., Huang, Y., 2007, Evidence for water use efficiency as an important factor in determining the δD values of tree leaf waxes: Organic Geochemistry, v. 38, p. 1251 – 1255. Jessup, K.E., Barnes, P.W., Boutton, T.W., 2003, Vegetation dynamics in a Quercus-Juniperus savanna: An isotopic assessment: Journal of Vegetation Science, v. 14, p. 841–852. Kristen, I., Wilkes, H., Vieth, A., Zink, K.-G., Plessen, B., Thorpe, J., Partridge, T.C., Oberhänsli, H., 2010, Biomarker and stable carbon isotope analyses of sedimentary organic matter from Lake Tswaing: evidence for deglacial wetness and early Holocene drought from South Africa: Journal of Paleolimnology, v. 44, p. 143–160. 57 Liu, W., Feng, X., Ning, Y., Zhang, Q., Cao, Y., An, A., 2005, δ13C variation of C3 and C4 plants across an Asian monsoon rainfall gradient in arid northwestern China: Global Change Biology, v. 11, p. 1094–1100. Rhodes, E.J., Dating sediments using potassium feldspar single-grain IRSL: initial methodological considerations: Quaternary International, in revision, October 2014. 58 Supplementary figures and tables Figure S1. Comparison of δ 13 C wax values with graminoid pollen abundance. Error bars indicate 1σ uncertainties of replicate measurements. Figure S2. δ 13 C wax vs. δD wax . No significant correlation is shown between the two variables, implying δD wax variations are robust to vegetation changes captured by δ 13 C wax. 59 Table S1. Lake Elsinore C 28 n-alkanoic acid δD wax , δ 13 C wax , and graminoid pollen abundance. Depth (cm) Age (cal. yr BP) δD wax (‰) Graminoid- corrected δD wax (‰) Graminoid- and ice volume- corrected δD wax (‰) σ (‰) δ 13 C wax (‰) σ (‰) Graminoid pollen % 1201 8788 -129 -125 -125 2.5 12 1230 9126 -118 -115 -116 3.3 7.6 1240 9245 -115 -114 -115 0 -32.2 0.7 2.2 1270 9575 -144 -141 -142 0.7 6.7 1286 9743 -121 -115 -116 0.7 15 1324 10102 -112 -108 -109 3.4 11.7 1350 10349 -93 -89 -90 0.2 13.3 1381 10653 -106 -100 -101 1.3 17.1 1401 10843 -130 -125 -126 1.1 13.9 1457 11490 -145 -129 -131 0.9 45.8 1485 11827 -143 -137 -140 0.9 17.2 1506 12111 -154 -148 -151 5.8 15.3 1533 12547 -151 -146 -149 1.6 13.6 1563 13071 -151 -143 -147 0.2 23.3 1596 13262 -149 -145 -149 2.5 -31.0 0.1 8.8 1641 13568 -146 -144 -148 3 4.8 1677 14061 -153 -151 -155 4.3 4.6 1706 14527 -170 -168 -173 3.3 6 1735 15276 -184 -182 -187 2.6 6.2 1768 16321 -182 -181 -187 2.5 3.8 1811 17217 -169 -168 -175 0.1 1.9 1841 17708 -192 -190 -197 1.7 6.2 1865 17951 -184 -183 -190 7 3.4 1880 18103 -170 -168 -175 5.8 -28.8 0.9 6.9 1911 18592 -191 -189 -196 0.3 6.9 1952 19219 -186 -184 -191 4.5 -29.4 0.9 4.8 1971 19520 -190 -189 -195 0 4.6 1996 19936 -137 -135 -143 1.4 4.1 2022 20361 -160 -158 -165 0.7 3.9 2041 20674 -179 -177 -183 0.2 5.8 2074 21370 -200 -197 -203 0.1 -30.1 0.0 7.7 2094 21798 -187 -182 -189 2 12.7 2115 22252 -186 -183 -189 0.4 8.2 2122 22388 -198 -195 -202 0 8.4 2135 22580 -176 -174 -180 2.6 -28.1 0.0 6.9 2155 22879 -142 -141 -147 4.4 5.3 2175 23190 -132 -130 -136 0.8 -32.2 0.1 6.8 2195 23506 -146 -141 -148 1.6 14.1 60 2215 23901 -157 -152 -159 3.7 12.1 2217 23944 -136 -131 -138 1.1 -28.1 0.1 14.2 2242 24474 -186 -177 -183 1.8 -31.8 0.0 24.1 2252 24674 -168 -162 -168 0.6 -30.3 0.2 15.9 2273 25101 -181 -175 -182 1.9 -31.0 0.5 15 2275 25142 -155 -150 -156 1.7 14.2 2279 25223 -190 -184 -190 1 14.2 2291 25339 -208 -195 -200 1 36.2 2310 25507 -203 -193 -199 1.6 26.4 2325 25612 -201 -191 -197 3.4 26.6 2344 25674 -176 -171 -177 1.6 -32.0 0.6 14.3 2380 26027 -183 -179 -184 0.2 11 2421 26456 -176 -168 -174 0.3 22.2 2462 26764 -195 -192 -197 3.5 9.7 2480 26923 -187 -178 -184 3.9 25.9 2497 27096 -208 -206 -211 0.1 -29.6 0.7 7.5 2530 27429 -179 -175 -181 0.7 9.3 2555 27676 -154 -149 -156 6.5 13.5 2598 28130 -159 -155 -161 4 11.3 2643 28592 -134 -128 -134 0.2 -27.9 0.2 16.4 2680 28947 -157 -151 -157 1.4 16.8 2726 29400 -162 -157 -163 1.7 12.6 2780 29934 -162 -158 -164 4 -28.9 0.3 10.5 2810 30228 -192 -189 -194 n.a. 8.9 2825 30373 -169 -165 -170 0.5 12.1 2855 30694 -194 -188 -194 0.1 -32.3 0.1 16.3 2870 30872 -189 -184 -190 3.8 13.7 2881 31002 -200 -192 -198 1.6 21.3 2900 31217 -205 -198 -203 6.9 -30.8 0.5 19.9 2915 31382 -201 -197 -203 1.8 11.2 2961 31849 -188 -185 -191 0.8 6.6 3001 32256 -189 -185 -191 1.8 11

Abstract (if available)

Abstract During the Late Pleistocene, millennial-scale abrupt transitions in the North Atlantic may have influenced climate in remote regions including the western United States (US). Several speleothem reconstructions from the Great Basin and interior southwest US capture oxygen isotopic evidence for shifts in Pacific and North American Monsoon derived storm tracks, revealing the power of isotopes in precipitation to track hydroclimatic regime change. Here, we present a high-resolution lake sediment leaf wax hydrogen isotope record spanning 9 – 32 cal. kyr BP from Lake Elsinore, California, in the coastal southwest US. This site provides a valuable archive within an exclusively Pacific-influenced precipitation zone. We use leaf wax hydrogen isotopic composition (δDwax) as a proxy for δD of precipitation (δDprecip), a tracer of moisture sources and storm trajectories. Our interpretation is further based on grain size and graminoid pollen data from the same core, as well as time-transient experiments performed with the Earth System model LOVECLIM.We find large-magnitude (~70‰) millennial-scale fluctuations of δDwax during the late glacial, and a positive shift (>80‰) during the deglaciation, suggesting fundamental alterations in Pacific moisture delivery and storm trajectories. The largest transition (>80‰) is observed starting at the end of H1 and with no marked perturbation during the Younger Dryas. During the glacial, generally negative δDwax values reflect moisture delivery by more D-depleted northerly storms likely resulting from a southward-displacement of the polar jet stream. Based on model comparisons these conditions are likely associated with cooler conditions in the North Atlantic Ocean and a deepened Aleutian low pressure cell in the North Pacific. Positive anomalies in δDwax values during glacial millennial-scale intervals as well as the larger magnitude deglacial positive shift indicate a transition to a mode of atmospheric circulation delivering moisture from the tropical and subtropical North Pacific. Late glacial δDwax fluctuations generally correlate in timing with North Atlantic abrupt climate transitions, indicating an Atlantic-Pacific teleconnection, which is further highlighted by climate model simulations.

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Creator Wu, Mong Sin (author)

Core Title Late Pleistocene changes in winter moisture source in the coastal southwest United States

School College of Letters, Arts and Sciences

Degree Master of Science

Degree Program Geological Sciences

Publication Date 11/24/2014

Defense Date 10/27/2014

Publisher University of Southern California (original), University of Southern California. Libraries (digital)

Tag AMOC,Late Pleistocene,leaf wax hydrogen isotope,NPDW,OAI-PMH Harvest,Southwest United States

Format application/pdf (imt)

Language English

Contributor Electronically uploaded by the author (provenance)

Advisor Feakins, Sarah J. (committee chair), Berelson, William M. (committee member), Hammond, Douglas E. (committee member)

Creator Email mongsinw@usc.edu

Permanent Link (DOI) https://doi.org/10.25549/usctheses-c3-517796

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