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 UCD13C data,
127
and “corrected” UCSB13C 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
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