2012PA002422text01

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Auxiliary text
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Ecological associations of planktonic foraminiferal species
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The total foraminiferal flux peaks in mid-summer when surface waters are warm
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and well-stratified, though total fluxes are also high after upwelling in response to diatom
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blooms [Kincaid et al., 2000]. G. bulloides is abundant throughout the year regardless of
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upwelling, though elsewhere (including Oregon [Ortiz et al., 1995] and nearby San Pedro
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Basin [Sautter and Thunell, 1991]) it peaks during upwelling. T. quinqueloba also peaks
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during upwelling in SBB [Kincaid et al., 2000]. N. pachyderma (d) is most abundant in
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SBB when surface waters are warm and stratified in summer and autumn [Kincaid et al.,
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2000]. N. dutertrei and G. ruber have the same seasonal distribution as N. pachyderma
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(d), but their fluxes are an order of magnitude lower because they are nearer to the limits
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of their temperature tolerances [Kincaid et al., 2000]. O. universa is present throughout
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the year in SBB except during upwelling [Kincaid et al., 2000]. Although N. pachyderma
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(s) is not found in the modern SBB [Kincaid et al., 2000], it was abundant during MIS 3
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stadials [Hendy and Kennett, 2000], and is abundant today in North Pacific upwelling
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zones and in the CC [Kahn and Williams, 1981; Ortiz and Mix, 1992; Ortiz et al., 1996].
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It favors surface waters with little or no thermocline at 6-8°C with high nutrient levels,
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whereas N. pachyderma (d) usually inhabits the thermocline in moderately stratified
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waters at 8-14°C [Ortiz et al., 1995; Reynolds and Thunell, 1986].
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In summary, we use G. bulloides and T. quinqueloba as upwelling indicators
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[Kincaid et al., 2000; Ortiz et al., 1995; Sautter and Thunell, 1991] and N. pachyderma
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(s) as an indicator of increased transport of subarctic waters or upwelling [Kahn and
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Williams, 1981; Ortiz and Mix, 1992; Ortiz et al., 1996]. We also analyze abundances of
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N. pachyderma (d), which blooms when surface waters are stratified and <8-10°C
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[Kincaid et al., 2000; Ortiz et al., 1995; Reynolds and Thunell, 1986], and “warm forms”
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(G. inflata, G. ruber, O. universa, and N. dutertrei [Kennett and Venz, 1992]), which
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indicate temperatures >10.5-15°C [Bijma et al., 1990; Ortiz et al., 1995; Tolderlund and
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Be, 1971] and/or increased advection of southern-sourced waters [Kincaid et al., 2000;
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Sautter and Thunell, 1991]. Additionally, we plot total number of planktonic
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foraminifera, which responds to upwelling and summer warmth/stratification [Kincaid et
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al., 2000].
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Late Pleistocene evolution of N. pachyderma (s)
The MIS 18 and MIS 12 records have much lower % N. pachyderma dextral and
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much higher % N. pachyderma (s) during interstadials than the MIS 3 record [Hendy and
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Kennett, 1999; 2000], even though temperatures were similar. A possible explanation is
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that this species had not yet evolved to occupy its modern ecological niche. [Kucera and
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Kennett, 2000] found that N. pachyderma (s) did not adopt its modern morphology and
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ecological niche until ~1 Ma. It is possible that the ecology of this species continued to
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evolve during the Late Pleistocene, which may explain the difference between the clear
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climate response of % N. pachyderma dextral in the youngest core (~293 ka) versus its
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muted response in the older cores (~450 ka and ~735 ka).
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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
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planktonic specimens are plotted). This process yielded a near continuous (all gaps ≤10
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cm) record for cores MV0508-16JPC and MV0508-20JPC, with the exception of a gap at
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245-270 centimeters below sea floor (cmbsf) in MV0508-16JPC. The planktonic
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assemblage record for MV0508-11JPC is semi-continuous, with gaps of ≤25 cm, plus
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gaps at 55-95 and 105-155 cmbsf.
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Details of chronology
The ages of all cores in this study were estimated by an integrated and sequential
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process of: 1) Estimating ages of seismic stratigraphic reference horizons identified in
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ODP Site 893 (near the basin center) that lie between identified, dated horizons
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[Marshall, 2012; Nicholson et al., 2006]. Sedimentation rates were shown to be relatively
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constant near the basin center at time scales from decades to 1 Myr [Behl et al., 2007].
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We then correlated seismic reference horizons to seafloor outcrop [Marshall, 2012;
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Nicholson et al., 2006]. Ages of individual cores were estimated based on their
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stratigraphic position (Figure A1b) at or between seismic reference horizons, and
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adjusted based upon available biostratigraphic and tephrochonologic constraints (as
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described below).
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2) Further tuning the age of individual cores by comparing their oxygen isotopic records
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(range, average, and behavior) with the LR04 global benthic δ18O stack [Lisiecki and
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Raymo, 2005]. In this step, cores are shifted to the closest appropriate interval with the
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same character, e.g., full glacial, full interglacial, intermediate, warming or cooling trend,
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etc. In doing this, we recognize that the extremely high sedimentation rate planktonic
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records from SBB can contain higher variability or larger magnitude shifts than are
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recognized in the global stacked records.
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For example, the age of MV0508-11JPC was further constrained by its inclusion
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of a biostratigraphic datum (the last occurrence (LO) of the diatom Proboscia curvirostis
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during early MIS 8 [Koc et al., 2001]) and the presence of a large negative shift in both
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benthic and planktonic δ18O that only persists a few kyr, which places this record at the
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MIS 8.6-8.5 boundary. The MIS 8.5 sea-level highstand is recorded in coral terraces and
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dated at 293 ka ±5 kyr [Stirling et al., 2001]. Due to sediment winnowing in the upper
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portion of MV0508-11JPC (as evidenced by the paucity of foraminifera and the
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abundance of glauconite and sponge spicules), it is possible that this interval represents
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more time than the sedimentation rate implies.
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The age of MV0508-16JPC was further constrained by its inclusion of the LO of
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the nannofossil Pseudoemiliania lacunosa [Beaufort, pers. comm., 2010] at 450.61-
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452.20 ka [Sato et al., 2009]. Because this biostratigraphic datum was originally dated in
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marine sediments from the North Atlantic, we enlarge the age uncertainty of the
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biostratigraphic correlation to our core to ±5 kyr, to account for potential interbasin
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differences. This yields a core age of 450 ka ±5 kyr. This age is also consistent with the
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cool intermediate climatic state indicated by the oxygen isotopic record.
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The age of MV0508-20JPC was determined by its stratigraphic distance from the
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Lava Creek B ash (639 ka ±2 kyr), which was identified in an overlying core in the suite
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and dated at the USGS Tephra Laboratory, as well as its inclusion of the foraminifer
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Neogloboquadrina inglei, whose LO was at ~712 ka [Kennett et al., 2000; Kucera and
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Kennett, 2000]. The age was further narrowed by comparing the glacial-to-intermediate
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character of this core’s oxygen isotopic record with the LR04 stack oxygen isotope
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stratigraphy [Lisiecki and Raymo, 2005]. Elimination of inconsistent glacial or
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interglacial stages leaves MIS 18.3 [Prell et al., 1986] as the only interval consistent with
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seismic stratigraphic position, relative position with other dated cores, biostratigraphy,
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and oxygen isotopic character. Together, these constraints yield a core age of 735 ka ±5
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kyr, although this uncertainty does not reflect the low probability that MV0508-20JPC
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may represent the intermediate, Younger Dryas-like pause in the MIS 18 termination at
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~710 ka, or the intermediate pause in cooling at ~760 ka.
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Uncertainties in core ages were estimated based on the combined uncertainty in
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the nearest dated reference datum, the estimated uncertainty in the core's stratigraphic
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position and/or interpolated distance to the datum, and/or the confidence to which the
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isotopic record in the core could be matched to a distinctive dated excursion in the known
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global isotopic curve. We also note the possibility that observed LO horizons in the cores
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may predate the true LOs if they lie in the unsampled interval between cores; this
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uncertainty is not reflected in the reported age errors.
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Estimation of sedimentation rates
Sedimentation rates for the cores in our study were determined based on laminae
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counts and stratigraphic thicknesses of laminated intervals in each core [Escobedo, 2009].
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Average rates were 83 cm/kyr for MV0508-11JPC, 122 cm/kyr for MV0508-16JPC, and
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82 cm/kyr for MV0508-20JPC. However, to account for possible variation in
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sedimentation rate between laminated and bioturbated intervals, we use a general rate of
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100 ±20 cm/kyr to estimate the duration of various events and transitions in this study.
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Interlaboratory collaboration & comparison
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Analyses performed at UC Santa Barbara (UCSB) include the following: N.
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pachyderma (s) for core MV0508-11JPC and U. peregrina for cores MV0508-16JPC and
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MV0508-20JPC, plus G. bulloides for all cores. Analyses performed at the UC Davis
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Stable Isotope Laboratory (UCD) include the following: N. pachyderma (s) for cores
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MV0508-16JPC and MV0508-20JPC and U. peregrina for core MV0508-11JPC.
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An ANOVA on 101 duplicates showed that the 18O datasets from UCD and
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UCSB are statistically indistinguishable, but the 13C data are not, exhibiting a -0.18‰
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offset. This laboratory offset is corrected by shifting UCSB 13C data by +0.18‰; a
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second ANOVA showed the corrected UCSB and raw UCD data to be statistically
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indistinguishable. Thus, all plots show raw 18O data from both labs, raw UCD13C data,
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and “corrected” UCSB13C data.
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Calculating δ18O of seawater
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To calculate δ18Ow during each study interval, we referred to the sea level record
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of [Bintanja et al., 2005], which shows that sea level was ~60 and ~90 meters below sea
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level (mbsl) during the intervals examined for MIS 18 and 12, respectively, and at the
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MIS 8.6-8.5 boundary, sea level rose from ~65 to 40 mbsl. Each estimate has an error of
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±12 mbsl. To calculate the sea level effect on δ18Ow, we assume that a 10 m sea level
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change yields a 0.1‰ change in δ18Ow [Shackleton and Opdyke, 1973]. For the MIS 8.6-
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8.5 boundary, we imposed a linear sea level rise at 420-320 cmbsf (to match the negative
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shift in our δ18O data). To calculate MIS 1-3 paleotemperatures in a manner comparable
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to that used for our other records, we imposed the following linear sea level changes,
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based on the sea level records of [Siddall et al., 2003] and [Bard et al., 1996]: from 60 ka
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to 28 ka, sea level was held constant at 80 mbsl, then from 28 ka to 18.1 ka (the time of
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most positive benthic δ18O values in SBB), sea level dropped to 120 mbsl. From 18.1 ka
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to 6 ka, sea level rose from 120 mbsl to 0 mbsl, and since 6 ka, sea level was constant.
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For both the G. bulloides and U. peregrina paleotemperature calculations, we
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approximated the effect of salinity on δ18Ow by applying regionally calibrated
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δ18Ow:salinity relationships to modern salinity values from CALCOFI station 80.52 [Lynn
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et al., 1982]. For the G. bulloides calculation, we applied the Southern California Bight
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relationship of [Bemis et al., 2002]:
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δ18Ow = 0.39*salinity – 13.23
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(3)
For U. peregrina, we used the relationship of [Zahn and Mix, 1991]:
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δ18Ow = 0.405*salinity – 14.014
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to reflect the influence of NPIW on SBB deep waters.
(4)
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Auxiliary figure caption
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Figure A1. a) High-resolution multibeam bathymetry [MBARI, 2000] over the Mid-
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Channel trend in Santa Barbara Basin with locations of cores (small black dots) and an
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example chirp line (Line 3) acquired in 2005. Locations of specific cores discussed in the
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text are shown by labels and large yellow dots. b) High-resolution USGS-Melville chirp
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line 3 over western culmination of the Mid-Channel Trend showing dipping outcrop
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strata of late-Quaternary age, mapped stratigraphic reference horizons (green, pink,
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violet, orange) and core locations (including along line with approximate core penetration
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depths converted to two-way travel time (seconds)). Core MV0508-11JPC, discussed in
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text, is labeled in red; other cores discussed in text are on separate chirp lines. Modified
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from [Nicholson et al., 2006].
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Figure A2. MIS 1-3, 0-60 ka δ18O records [Hendy and Kennett, 1999; 2003] and
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calculated paleotemperatures. a) Paleotemperature records converted from the G.
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bulloides δ18O record of [Hendy and Kennett, 1999] using the calibration of [Mulitza et
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al., 2003] (brown line), and from the U. peregrina δ18O record of [Hendy and Kennett,
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2003] using the equation of [Shackleton and Opdyke, 1973] (gray line). Salinity and sea
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level corrections are described in text. Red bars show 1σ error envelope for
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paleotemperature calculations (±0.8°C for U. peregrina, ±1.25°C for G. bulloides). Four
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unrealistically cold (<2°C) benthic temperature points are shown but are not connected to
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the line. b) δ18O of G. bulloides (navy blue line). Numbers next to δ18O data denote
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correlations with Greenland interstadial (IS) events [Hendy and Kennett, 1999].
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Figure A3. MIS 8.6-8.5, 293 ka ±5 kyr δ18O records versus a synthetic Greenland δ18O
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reconstruction of millennial-scale variability [Barker et al., 2011]. a) G. bulloides δ18O
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(navy blue diamonds) and 5-point running average (navy blue curve), plotted against age
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by assuming a constant sedimentation rate of 100 cm/kyr and pinning the most negative
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δ18O value in the record (at 309 cmbsf) to the sea-level highstand of MIS 8.5, dated to
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293 ka ±5 kyr [Stirling et al., 2001] . b) U. peregrina δ18O (turquoise circles), plotted
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against age in the same manner as b). c) Synthetic Greenland δ18O reconstruction of
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millennial-scale variability using the SpeloAge scale, plotted as ‰ deviations from the
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7000-year smoothed record [Barker et al., 2011 “GLT_syn_hi”]. The synthetic
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Greenland δ18O reconstruction was modeled from the EPICA Dome C δD record
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assuming a bipolar seesaw mechanism, then millennial-scale variability was isolated by
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subtracting a 7000-year smoothed signal from the total signal [Barker et al., 2011].
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