1 Auxiliary text 2 Ecological associations of planktonic foraminiferal species 3 The total foraminiferal flux peaks in mid-summer when surface waters are warm 4 and well-stratified, though total fluxes are also high after upwelling in response to diatom 5 blooms [Kincaid et al., 2000]. G. bulloides is abundant throughout the year regardless of 6 upwelling, though elsewhere (including Oregon [Ortiz et al., 1995] and nearby San Pedro 7 Basin [Sautter and Thunell, 1991]) it peaks during upwelling. T. quinqueloba also peaks 8 during upwelling in SBB [Kincaid et al., 2000]. N. pachyderma (d) is most abundant in 9 SBB when surface waters are warm and stratified in summer and autumn [Kincaid et al., 10 2000]. N. dutertrei and G. ruber have the same seasonal distribution as N. pachyderma 11 (d), but their fluxes are an order of magnitude lower because they are nearer to the limits 12 of their temperature tolerances [Kincaid et al., 2000]. O. universa is present throughout 13 the year in SBB except during upwelling [Kincaid et al., 2000]. Although N. pachyderma 14 (s) is not found in the modern SBB [Kincaid et al., 2000], it was abundant during MIS 3 15 stadials [Hendy and Kennett, 2000], and is abundant today in North Pacific upwelling 16 zones and in the CC [Kahn and Williams, 1981; Ortiz and Mix, 1992; Ortiz et al., 1996]. 17 It favors surface waters with little or no thermocline at 6-8°C with high nutrient levels, 18 whereas N. pachyderma (d) usually inhabits the thermocline in moderately stratified 19 waters at 8-14°C [Ortiz et al., 1995; Reynolds and Thunell, 1986]. 20 In summary, we use G. bulloides and T. quinqueloba as upwelling indicators 21 [Kincaid et al., 2000; Ortiz et al., 1995; Sautter and Thunell, 1991] and N. pachyderma 22 (s) as an indicator of increased transport of subarctic waters or upwelling [Kahn and 23 Williams, 1981; Ortiz and Mix, 1992; Ortiz et al., 1996]. We also analyze abundances of 24 N. pachyderma (d), which blooms when surface waters are stratified and <8-10°C 25 [Kincaid et al., 2000; Ortiz et al., 1995; Reynolds and Thunell, 1986], and “warm forms” 26 (G. inflata, G. ruber, O. universa, and N. dutertrei [Kennett and Venz, 1992]), which 27 indicate temperatures >10.5-15°C [Bijma et al., 1990; Ortiz et al., 1995; Tolderlund and 28 Be, 1971] and/or increased advection of southern-sourced waters [Kincaid et al., 2000; 29 Sautter and Thunell, 1991]. Additionally, we plot total number of planktonic 30 foraminifera, which responds to upwelling and summer warmth/stratification [Kincaid et 31 al., 2000]. 32 33 34 Late Pleistocene evolution of N. pachyderma (s) The MIS 18 and MIS 12 records have much lower % N. pachyderma dextral and 35 much higher % N. pachyderma (s) during interstadials than the MIS 3 record [Hendy and 36 Kennett, 1999; 2000], even though temperatures were similar. A possible explanation is 37 that this species had not yet evolved to occupy its modern ecological niche. [Kucera and 38 Kennett, 2000] found that N. pachyderma (s) did not adopt its modern morphology and 39 ecological niche until ~1 Ma. It is possible that the ecology of this species continued to 40 evolve during the Late Pleistocene, which may explain the difference between the clear 41 climate response of % N. pachyderma dextral in the youngest core (~293 ka) versus its 42 muted response in the older cores (~450 ka and ~735 ka). 43 44 45 46 Continuity of record The faunal assemblage record was based 2 cm samples spaced every 5 cm and containing >300 planktonic specimens, where possible (only samples with >100 47 planktonic specimens are plotted). This process yielded a near continuous (all gaps ≤10 48 cm) record for cores MV0508-16JPC and MV0508-20JPC, with the exception of a gap at 49 245-270 centimeters below sea floor (cmbsf) in MV0508-16JPC. The planktonic 50 assemblage record for MV0508-11JPC is semi-continuous, with gaps of ≤25 cm, plus 51 gaps at 55-95 and 105-155 cmbsf. 52 53 54 Details of chronology The ages of all cores in this study were estimated by an integrated and sequential 55 process of: 1) Estimating ages of seismic stratigraphic reference horizons identified in 56 ODP Site 893 (near the basin center) that lie between identified, dated horizons 57 [Marshall, 2012; Nicholson et al., 2006]. Sedimentation rates were shown to be relatively 58 constant near the basin center at time scales from decades to 1 Myr [Behl et al., 2007]. 59 We then correlated seismic reference horizons to seafloor outcrop [Marshall, 2012; 60 Nicholson et al., 2006]. Ages of individual cores were estimated based on their 61 stratigraphic position (Figure A1b) at or between seismic reference horizons, and 62 adjusted based upon available biostratigraphic and tephrochonologic constraints (as 63 described below). 64 2) Further tuning the age of individual cores by comparing their oxygen isotopic records 65 (range, average, and behavior) with the LR04 global benthic δ18O stack [Lisiecki and 66 Raymo, 2005]. In this step, cores are shifted to the closest appropriate interval with the 67 same character, e.g., full glacial, full interglacial, intermediate, warming or cooling trend, 68 etc. In doing this, we recognize that the extremely high sedimentation rate planktonic 69 records from SBB can contain higher variability or larger magnitude shifts than are 70 recognized in the global stacked records. 71 For example, the age of MV0508-11JPC was further constrained by its inclusion 72 of a biostratigraphic datum (the last occurrence (LO) of the diatom Proboscia curvirostis 73 during early MIS 8 [Koc et al., 2001]) and the presence of a large negative shift in both 74 benthic and planktonic δ18O that only persists a few kyr, which places this record at the 75 MIS 8.6-8.5 boundary. The MIS 8.5 sea-level highstand is recorded in coral terraces and 76 dated at 293 ka ±5 kyr [Stirling et al., 2001]. Due to sediment winnowing in the upper 77 portion of MV0508-11JPC (as evidenced by the paucity of foraminifera and the 78 abundance of glauconite and sponge spicules), it is possible that this interval represents 79 more time than the sedimentation rate implies. 80 The age of MV0508-16JPC was further constrained by its inclusion of the LO of 81 the nannofossil Pseudoemiliania lacunosa [Beaufort, pers. comm., 2010] at 450.61- 82 452.20 ka [Sato et al., 2009]. Because this biostratigraphic datum was originally dated in 83 marine sediments from the North Atlantic, we enlarge the age uncertainty of the 84 biostratigraphic correlation to our core to ±5 kyr, to account for potential interbasin 85 differences. This yields a core age of 450 ka ±5 kyr. This age is also consistent with the 86 cool intermediate climatic state indicated by the oxygen isotopic record. 87 The age of MV0508-20JPC was determined by its stratigraphic distance from the 88 Lava Creek B ash (639 ka ±2 kyr), which was identified in an overlying core in the suite 89 and dated at the USGS Tephra Laboratory, as well as its inclusion of the foraminifer 90 Neogloboquadrina inglei, whose LO was at ~712 ka [Kennett et al., 2000; Kucera and 91 Kennett, 2000]. The age was further narrowed by comparing the glacial-to-intermediate 92 character of this core’s oxygen isotopic record with the LR04 stack oxygen isotope 93 stratigraphy [Lisiecki and Raymo, 2005]. Elimination of inconsistent glacial or 94 interglacial stages leaves MIS 18.3 [Prell et al., 1986] as the only interval consistent with 95 seismic stratigraphic position, relative position with other dated cores, biostratigraphy, 96 and oxygen isotopic character. Together, these constraints yield a core age of 735 ka ±5 97 kyr, although this uncertainty does not reflect the low probability that MV0508-20JPC 98 may represent the intermediate, Younger Dryas-like pause in the MIS 18 termination at 99 ~710 ka, or the intermediate pause in cooling at ~760 ka. 100 Uncertainties in core ages were estimated based on the combined uncertainty in 101 the nearest dated reference datum, the estimated uncertainty in the core's stratigraphic 102 position and/or interpolated distance to the datum, and/or the confidence to which the 103 isotopic record in the core could be matched to a distinctive dated excursion in the known 104 global isotopic curve. We also note the possibility that observed LO horizons in the cores 105 may predate the true LOs if they lie in the unsampled interval between cores; this 106 uncertainty is not reflected in the reported age errors. 107 108 109 Estimation of sedimentation rates Sedimentation rates for the cores in our study were determined based on laminae 110 counts and stratigraphic thicknesses of laminated intervals in each core [Escobedo, 2009]. 111 Average rates were 83 cm/kyr for MV0508-11JPC, 122 cm/kyr for MV0508-16JPC, and 112 82 cm/kyr for MV0508-20JPC. However, to account for possible variation in 113 sedimentation rate between laminated and bioturbated intervals, we use a general rate of 114 100 ±20 cm/kyr to estimate the duration of various events and transitions in this study. 115 116 Interlaboratory collaboration & comparison 117 Analyses performed at UC Santa Barbara (UCSB) include the following: N. 118 pachyderma (s) for core MV0508-11JPC and U. peregrina for cores MV0508-16JPC and 119 MV0508-20JPC, plus G. bulloides for all cores. Analyses performed at the UC Davis 120 Stable Isotope Laboratory (UCD) include the following: N. pachyderma (s) for cores 121 MV0508-16JPC and MV0508-20JPC and U. peregrina for core MV0508-11JPC. 122 An ANOVA on 101 duplicates showed that the 18O datasets from UCD and 123 UCSB are statistically indistinguishable, but the 13C data are not, exhibiting a -0.18‰ 124 offset. This laboratory offset is corrected by shifting UCSB 13C data by +0.18‰; a 125 second ANOVA showed the corrected UCSB and raw UCD data to be statistically 126 indistinguishable. Thus, all plots show raw 18O data from both labs, raw UCD13C data, 127 and “corrected” UCSB13C data. 128 129 Calculating δ18O of seawater 130 To calculate δ18Ow during each study interval, we referred to the sea level record 131 of [Bintanja et al., 2005], which shows that sea level was ~60 and ~90 meters below sea 132 level (mbsl) during the intervals examined for MIS 18 and 12, respectively, and at the 133 MIS 8.6-8.5 boundary, sea level rose from ~65 to 40 mbsl. Each estimate has an error of 134 ±12 mbsl. To calculate the sea level effect on δ18Ow, we assume that a 10 m sea level 135 change yields a 0.1‰ change in δ18Ow [Shackleton and Opdyke, 1973]. For the MIS 8.6- 136 8.5 boundary, we imposed a linear sea level rise at 420-320 cmbsf (to match the negative 137 shift in our δ18O data). To calculate MIS 1-3 paleotemperatures in a manner comparable 138 to that used for our other records, we imposed the following linear sea level changes, 139 based on the sea level records of [Siddall et al., 2003] and [Bard et al., 1996]: from 60 ka 140 to 28 ka, sea level was held constant at 80 mbsl, then from 28 ka to 18.1 ka (the time of 141 most positive benthic δ18O values in SBB), sea level dropped to 120 mbsl. From 18.1 ka 142 to 6 ka, sea level rose from 120 mbsl to 0 mbsl, and since 6 ka, sea level was constant. 143 For both the G. bulloides and U. peregrina paleotemperature calculations, we 144 approximated the effect of salinity on δ18Ow by applying regionally calibrated 145 δ18Ow:salinity relationships to modern salinity values from CALCOFI station 80.52 [Lynn 146 et al., 1982]. For the G. bulloides calculation, we applied the Southern California Bight 147 relationship of [Bemis et al., 2002]: 148 δ18Ow = 0.39*salinity – 13.23 149 (3) For U. peregrina, we used the relationship of [Zahn and Mix, 1991]: 150 δ18Ow = 0.405*salinity – 14.014 151 to reflect the influence of NPIW on SBB deep waters. (4) 152 153 Auxiliary figure caption 154 Figure A1. a) High-resolution multibeam bathymetry [MBARI, 2000] over the Mid- 155 Channel trend in Santa Barbara Basin with locations of cores (small black dots) and an 156 example chirp line (Line 3) acquired in 2005. Locations of specific cores discussed in the 157 text are shown by labels and large yellow dots. b) High-resolution USGS-Melville chirp 158 line 3 over western culmination of the Mid-Channel Trend showing dipping outcrop 159 strata of late-Quaternary age, mapped stratigraphic reference horizons (green, pink, 160 violet, orange) and core locations (including along line with approximate core penetration 161 depths converted to two-way travel time (seconds)). Core MV0508-11JPC, discussed in 162 text, is labeled in red; other cores discussed in text are on separate chirp lines. Modified 163 from [Nicholson et al., 2006]. 164 165 Figure A2. MIS 1-3, 0-60 ka δ18O records [Hendy and Kennett, 1999; 2003] and 166 calculated paleotemperatures. a) Paleotemperature records converted from the G. 167 bulloides δ18O record of [Hendy and Kennett, 1999] using the calibration of [Mulitza et 168 al., 2003] (brown line), and from the U. peregrina δ18O record of [Hendy and Kennett, 169 2003] using the equation of [Shackleton and Opdyke, 1973] (gray line). Salinity and sea 170 level corrections are described in text. Red bars show 1σ error envelope for 171 paleotemperature calculations (±0.8°C for U. peregrina, ±1.25°C for G. bulloides). Four 172 unrealistically cold (<2°C) benthic temperature points are shown but are not connected to 173 the line. b) δ18O of G. bulloides (navy blue line). Numbers next to δ18O data denote 174 correlations with Greenland interstadial (IS) events [Hendy and Kennett, 1999]. 175 176 Figure A3. MIS 8.6-8.5, 293 ka ±5 kyr δ18O records versus a synthetic Greenland δ18O 177 reconstruction of millennial-scale variability [Barker et al., 2011]. a) G. bulloides δ18O 178 (navy blue diamonds) and 5-point running average (navy blue curve), plotted against age 179 by assuming a constant sedimentation rate of 100 cm/kyr and pinning the most negative 180 δ18O value in the record (at 309 cmbsf) to the sea-level highstand of MIS 8.5, dated to 181 293 ka ±5 kyr [Stirling et al., 2001] . b) U. peregrina δ18O (turquoise circles), plotted 182 against age in the same manner as b). c) Synthetic Greenland δ18O reconstruction of 183 millennial-scale variability using the SpeloAge scale, plotted as ‰ deviations from the 184 7000-year smoothed record [Barker et al., 2011 “GLT_syn_hi”]. The synthetic 185 Greenland δ18O reconstruction was modeled from the EPICA Dome C δD record 186 assuming a bipolar seesaw mechanism, then millennial-scale variability was isolated by 187 subtracting a 7000-year smoothed signal from the total signal [Barker et al., 2011]. 188 189 References 190 191 192 193 194 195 196 197 198 199 200 201 202 203 204 205 206 207 208 209 210 211 212 213 214 215 216 217 218 219 220 221 222 223 Bard, E., B. Hamelin, M. Arnold, L. Montaggioni, G. Cabioch, G. Faure, and F. Rougerie (1996), Deglacial sea-level record from Tahiti corals and the timing of global meltwater discharge, Nature, 382(6588), 241-244. Barker, S., G. Knorr, R. L. Edwards, F. Parrenin, A. E. Putnam, L. C. Skinner, E. Wolff, and M. Ziegler (2011), 800,000 Years of Abrupt Climate Variability, Science, 334(6054), 347-351. Bemis, B. E., H. J. Spero, and R. C. Thunell (2002), Using species-specific paleotemperature equations with foraminifera: a case study in the Southern California Bight, Marine Micropaleontology, 46(3-4), 405-430. Bintanja, R., R. S. W. van de Wal, and J. Oerlemans (2005), Modelled atmospheric temperatures and global sea levels over the past million years, Nature, 437(7055), 125-128. Hendy, I. L., and J. P. Kennett (1999), Latest Quaternary North Pacific surface-water responses imply atmosphere-driven climate instability, Geology, 27(4), 291-294. Hendy, I. L., and J. P. Kennett (2000), Dansgaard-Oeschger Cycles and the California Current System: Planktonic Foraminiferal Response to Rapid Climate Change in Santa Barbara Basin, Ocean Drilling Program Hole 893A, Paleoceanography, 15(1), 30-42. Hendy, I. L., and J. P. Kennett (2003), Tropical forcing of North Pacific intermediate water distribution during Late Quaternary rapid climate change?, Quaternary Science Reviews, 22(5-7), 673-689. Kahn, M. I., and D. F. Williams (1981), Oxygen and carbon isotopic composition of living planktonic foraminifera from the northeast Pacific Ocean, Palaeogeography, Palaeoclimatology, Palaeoecology, 33(1-3), 47-69. Kennett, J. P., G. A. Rozo-Vera, and M. L. M. Castillo (2000), Latest Neogene planktonic foraminiferal biostratigraphy of the California margin, Proceedings of the Ocean Drilling Program, Scientific Results, 167, 41-62. Kincaid, E., R. C. Thunell, J. Le, C. B. Lange, A. L. Weinheimer, and F. M. H. Reid (2000), Planktonic foraminiferal fluxes in the Santa Barbara Basin: response to seasonal and interannual hydrographic changes, Deep Sea Research Part II: Topical Studies in Oceanography, 47(5-6), 1157-1176. Koc, N., L. Labeyrie, S. Manthe, B. P. Flower, D. A. Hodell, and A. Aksu (2001), The last occurrence of Proboscia curvirostris in the North Atlantic marine isotope stages 9-8, Marine Micropaleontology, 41(1), 9-23. 224 225 226 227 228 229 230 231 232 233 234 235 236 237 238 239 240 241 242 243 244 245 246 247 248 249 250 251 252 253 254 255 256 257 258 259 260 261 262 263 264 265 266 267 268 Kucera, M., and J. Kennett (2000), Biochronology and evolutionary implications of late Neogene California margin planktonic foraminiferal events, Marine Micropaleontology, 40(1), 67-82. Lisiecki, L. E., and M. E. Raymo (2005), A Pliocene-Pleistocene stack of 57 globally distributed benthic δ18O records, Paleoceanography, 20, PA1003. Lynn, R. J., K. A. Bliss, L. E. Eber, and C. M. R. Committee (1982), Vertical and horizontal distributions of seasonal mean temperature, salinity, sigma-t, stability, dynamic height, oxygen, and oxygen saturation in the California Current, 19501978, Marine Research Committee. Marshall, C. J. (2012), Sedimentation in an active fold and thrust belt, from 1 Ma to present, Santa Barbara Channel, California, California State University, Long Beach, Long Beach. Monterey Bay Aquarium Research Institute (2000), Santa Barbara multibeam survey, Digital Data Ser., 4, Moss Landing, Calif. Mulitza, S., D. Boltovskoy, B. Donner, H. Meggers, A. Paul, and G. Wefer (2003), Temperature: d18O relationships of planktonic foraminifera collected from surface waters, Palaeogeography, Palaeoclimatology, Palaeoecology, 202(1), 143-152. Nicholson, C., et al. (2006), Santa Barbara Basin study extends global climate record, Eos, Transactions, American Geophysical Union, 87(21), 205-208. Ortiz, J. D., and A. C. Mix (1992), The spatial distribution and seasonal succession of planktonic foraminifera in the California Current off Oregon, September 1987– September 1988, Geological Society, London, Special Publications, 64(1), 197. Ortiz, J. D., A. C. Mix, and R. W. Collier (1995), Environmental control of living symbiotic and asymbiotic foraminifera of the California Current, Paleoceanography, 10(6), 987-1009. Ortiz, J. D., A. C. Mix, W. Rugh, J. M. Watkins, and R. W. Collier (1996), Deepdwelling planktonic foraminifera of the northeastern Pacific Ocean reveal environmental control of oxygen and carbon isotopic disequilibria, Geochimica et Cosmochimica Acta, 60(22), 4509-4523. Prell, W. L., J. Imbrie, D. G. Martinson, J. J. Morley, N. G. Pisias, N. J. Shackleton, and H. F. Streeter (1986), Graphic correlation of oxygen isotope stratigraphy application to the late Quaternary, Paleoceanography, 1(2), 137-162. Reynolds, L. A., and R. C. Thunell (1986), Seasonal Production and Morphologic Variation of Neogloboquadrina pachyderma (Ehrenberg) in the Northeast Pacific, Micropaleontology, 32(1), 1-18. Sato, T., S. Chiyonobu, and D. A. Hodell (2009), Data report: Quaternary calcareous nannofossil datums and biochronology in the North Atlantic Ocean, IODP Site U1308, Channell, JET, Kanamatsu, T., Sato, T., Stein, R., Alvarez Zarikian, CA, Malone, MJ, the Expedition, 303, 306. Sautter, L. R., and R. C. Thunell (1991), Planktonic foraminiferal response to upwelling and seasonal hydrographic conditions - sediment trap results from San Pedro Basin, southern California bight, Journal of Foraminiferal Research, 21(4), 347363. Shackleton, N. J., and N. D. Opdyke (1973), Oxygen isotope and palaeomagnetic stratigraphy of Equatorial Pacific core V28-238: Oxygen isotope temperatures 269 270 271 272 273 274 275 276 277 278 279 280 281 and ice volumes on a 105 year and 106 year scale, Quaternary Research, 3(1), 3955. Siddall, M., E. J. Rohling, A. Almogi-Labin, C. Hemleben, D. Meischner, I. Schmelzer, and D. A. Smeed (2003), Sea-level fluctuations during the last glacial cycle, Nature, 423(6942), 853-858. Stirling, C. H., T. M. Esat, K. Lambeck, M. T. McCulloch, S. G. Blake, D.-C. Lee, and A. N. Halliday (2001), Orbital Forcing of the Marine Isotope Stage 9 Interglacial, Science, 291(5502), 290-293. Zahn, R., and A. C. Mix (1991), Benthic foraminiferal d18O in the ocean's temperaturesalinity-density field: Constraints on Ice Age thermohaline circulation, Paleoceanography, 6(1), 1-20.