Auxiliary material for
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Deep water formation in the North Pacific and deglacial CO2 rise
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James W.B. Rae1,2,3,4 *, Michael Sarnthein5, Gavin L. Foster6, Andy Ridgwell2, Pieter M.
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Grootes7, Tim Elliott1
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1 Bristol Isotope Group, Department of Earth Sciences, University of Bristol, Bristol BS8 1RJ,
UK
2 BRIDGE, School of Geographical Sciences, University of Bristol, Bristol BS8 1SS, UK
3 Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena
91125, USA
4 Department of Earth and Environmental Sciences, University of St Andrews, St Andrews,
KY16 9AL, UK
5 Institut für Geowissenschaften, University of Kiel, Kiel D-24098, Germany
6 Ocean and Earth Science, National Oceanography Centre Southampton, University of
Southampton, Southampton SO14 3ZH, UK
7 Institute of Ecosystem Research, University of Kiel, Kiel D-24098, Germany
Paleoceanography, 2014
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Auxiliary materials include:
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1. text01 (Rae_2014_Paleo_SUPPLEMENT.doc): Supplementary Information, which contains
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background on the boron isotope proxy for paleo-pH, and five supplementary figures referred to
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in the main text.
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2. ts01 (Rae_2014_Paleo-data-tables.xls): data tables, giving the stratigraphic tie points, 11B and
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column is labeled, including units.
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C data used in this study. Each worksheet gives a description of the materials within, and each
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INTERPRETATION OF BORON ISOTOPE DATA
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The systematics of the boron isotope proxy in benthic foraminifera are thoroughly described in
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Rae et al. (2011). Here we show how this carbonate system proxy may be used to examine a
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variety of biogeochemical processes over the deglaciation.
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Although boron isotopes are often parameterized in terms of pH, 11B of B(OH)4- (as recorded
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by benthic foraminifera; Rae et al., 2011) is a sensitive parameter of the carbonate system in its
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own right, due to the role of B(OH)4- in balancing conservative alkalinity. Figure S1 shows 11B
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of B(OH)4- and pH as a function of the controlling carbonate system variables, Alkalinity (ALK)
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and Dissolved Inorganic Carbon (DIC); this allows us to examine how common biogeochemical
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processes might be recorded by our 11B data. Note that contours of pH as a function of ALK
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and DIC have a very similar orientation and the same sense of change to 11B of B(OH)4-.
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Furthermore, contours of [CO2], [CO32-], and carbonate saturation state are also closely aligned
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with 11B of B(OH)4-. Thus, while full determination of the carbonate system requires two of its
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parameters to be known, knowledge of 11B of B(OH)4- and pH may still place strong constraints
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on [CO2], [CO32-], and carbonate saturation state.
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During HS1 and the YD we see a decrease in 11B. This has the opposite sense to what may be
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expected to occur as a result of DIC loss from the deep ocean to the atmosphere during these
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periods of atmospheric CO2 rise (Figure S1). In the absence of competing influences, DIC loss
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would result in an increase in deep ocean pH, so other processes must dominate our signal. The
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observed 11B minima also cannot be caused by input of DIC due to changes in productivity, as
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no coherent relationship is seen between 11B and records of productivity from this core ()
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(Figure 2, 7; Gebhardt et al., 2008). The lack of a productivity/remineralisation control on the
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carbonate system at the depth of our core is not surprising, given that most remineralisation
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occurs in relatively shallow subsurface waters and decreases exponentially with depth (Antia et
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al., 2001). The decrease we see in 11B during HS1 and the YD also cannot be attributed to
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enhanced CaCO3 burial. During these periods CaCO3 preservation shows minima in many deep
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ocean sediments (Barker et al., 2010; Gebhardt et al., 2008), while changes in shallow CaCO3
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burial would be slowly transmitted from surface waters to depth, and thus be unlikely to drive
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the rapid changes in deep pH that we observe. In any case, these biogeochemical processes
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alone would not significantly alter the deep water radiocarbon signature.
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As well as tracking different biogeochemical processes, 11B of B(OH)4- and pH may also track
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changes in circulation, due to the different biogeochemical histories of each water mass. Water
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mass mixing, via deep water formation, has the dominant effect on the excursions observed in
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our 11B record, with large scale convective mixing acting to homogenise the pH profile of the
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water column and bring low pH waters from intermediate depths to the depth of our core. This is
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illustrated schematically in Figure 1d, and quantitatively by our earth system modelling
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experiments (Figure 5 & 6). Deepwater formation is also consistent with the observed increase
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in deep 14C and minima in benthic-planktic 14C.
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SUPPLEMENTARY FIGURES
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Figure S1: δ11B of borate (as recorded by benthic foraminifera Rae et al. (2011)) and pH as a
function of the controlling carbonate system variables, Alkalinity (ALK) and Dissolved
Inorganic Carbon (DIC). Ocean temperature, salinity and pressure will slightly alter these 11B
values, but for a likely range of conditions at a given site, the effect of changes in these
parameters is < 0.1 ‰. This figure is plotted for conditions appropriate for our core site today, T
= 1.5 C, S = 35, P = 3600 dbar, with carbonate system calculations carried out using
CO2SYS.m (van Heuven et al., 2009) using parameters as recommended by Dickson et al.,
(2007), the seawater boron concentration of 432.6 mol/kg (Lee et al., 2010), 11Bsw of 39.61
‰ (Foster et al., 2010) and B = 27.2 ‰ (Klochko et al., 2006).
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Figure S2: MD02-2489 age model based on ties between δ18O of planktic foraminifera
(Gebhardt et al., 2008) and NGRIP ice core δ18O (NGRIPmembers, 2004; Johnsen et al., 1997)
on the GICC05 timescale (Rasmussen et al., 2006; Andersen et al., 2006; Svensson et al., 2006;
Rasmussen et al., 2008)) and East Asian speleothem δ18O (Wu et al., 2009; Dykoski et al., 2005;
Wu et al., 2009). Bold lines in the ice core and speleothem records show 5 point running means.
The solid grey lines bars indicate ties used to provide age control. Note that the gradual BøllingAllerød boundary at ~14.6 ka in G. bulloides relative to N. pachyderma (s.) may be explained by
bioturbation, as there is a pronounced increase in G. bulloides abundance at the BA; hence we
use the N. pachyderma record for this tie point. Conventional 14C dates are used in the younger
portion of the core, with a reservoir age of 850 years (R = 450 years (Southon & Fedje, 2003;
Key et al., 2004; Hutchinson et al. 2004), and the Marine09 curve from IntCal09 (Reimer et al.,
2009). Inset plot shows MD02-2489 core-depth vs. calendar age, with the slope giving
sedimentation rate. The inferred increase in sedimentation rate during the Bølling-Allerød may
have resulted from enhanced biological productivity, widely seen in the North Pacific at this time
(Kohfeld & Chase, 2011).
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Figure S3: a) Radiocarbon plateau-tuned age model for core MD02-2489 (see Sarnthein et al.,
2007). Boxes indicate plateau structures in the planktic 14C record that are tied to the Lake
Suigetsu reference record (Ramsey et al., 2012), as described in Sarnthein et al. (2013). b)
Benthic 11B and benthic-planktic 14C offsets plotted on both the 18O-based age model (solid
lines) and 14C plateau-tuned age model (dashed lines), along with atmospheric CO2 (Monnin et
al., 2001; Lourantou et al., 2010), synchronised to the GICC05 timescale by Lemieux-Dudon et
al. (2010). The younger portion of both age models uses conventional 14C dating, with a
reservoir age of 850 years (R = 450 years (Southon & Fedje, 2003; Key et al., 2004;
Hutchinson et al. 2004), and the Marine09 curve from IntCal09 (Reimer et al., 2009). The age
models show a good agreement, with both aligning the major benthic-planktic 14C excursion with
mid HS1.
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Figure S4: Updated age model for ODP 887. The early deglacial portion of the age model is
based on tying % CaCO3 data at ODP 887 (open squares; 54.37 N, 148.45 W, 3647 m) to %
CaCO3 data in our adjacent core MD02-2489 (filled squares; 54.39 N, 148.92 W, 3640 m;
Gebhardt et al., 2008), as shown in bottom panel. An XRF record of calcium intensity from
MD02-2489 (+ symbols) is also shown. Vertical grey lines indicated tie points. The LGM, late
deglacial and Holocene portions of the age model use conventional 14C dating (triangles in
bottom panel), with a reservoir age of 1800 ± 250 years for the LGM and 850 ± 100 years for the
late deglacial and Holocene. This age model gives good agreement in % total organic carbon
(mid panel) and planktic 14C (top panel) between our cores. The inset figure shows the resulting
sedimentation rate (open diamonds) and comparison to the previous ODP 887 age model (filled
circles; Galbraith et al., 2007).
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Figure S5: Comparison of deglacial North Pacific radiocarbon records. Sites include: the deep
NE Pacific (MD02-2489 at 54.39 N, 148.92 W, 3640 m, red stars; and ODP 887 (Galbraith et
al., 2007) at 54.37 N, 148.45 W, 3647 m, orange diamonds); mid-depths on the Oregon Margin
(W8709A-13PC at 42.1°N, 125.8°W, 2710 m, purple triangles; Lund et al., 2011); and a
compilation of intermediate depth NW Pacific cores (978 – 2800 m, green circles; Okazaki et al.,
2010). Planktic 14C data from MD02-2489 are also shown (yellow stars), with the dominantly N.
pachyderma (s.) samples recording shallow subsurface waters. The dot-dash line for W8709A13PC is an error-weighted mean through samples from the same core interval. The dashed line
for the NW Pacific compilation is a 3-point running mean. a) Benthic-planktic 14C offsets. Error
bars are the quadratic addition of benthic and planktic analytical uncertainties at 1 . b) 14C
data, with 1  error ellipses combining analytical and calendar age uncertainties. c) Difference
between foraminiferal 14C records and the 14C of the contemporaneous atmosphere (14C),
using atmospheric 14C from IntCal09 (Reimer et al., 2009), spliced with the Hulu Cave 14C
record for the 15 – 26 ka interval (Southon et al., 2012). Grey lines indicate constant ventilation
ages in calendar years relative to the contemporaneous atmosphere. d) Difference between
foraminiferal 14C age and that of the contemporaneous atmosphere (foram-atmosphere reservoir
age) with 1  Monte Carlo error arrays combining uncertainties in the foram 14C data,
atmospheric 14C record, and calendar age error. Note the large radiocarbon gradients at the
LGM, which collapse to a range of only ~60 ‰ or 500 yr at 17.3-16.0 ka. This is smaller than
the spread seen across a similar range of locations and depths in the modern North Atlantic.
Although benthic-planktic offsets at the Oregon margin site remain larger than those from this
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study during the deep water formation event, this is likely the result of surface waters at this
more southerly site staying stratified and giving young planktic 14C ages. Benthic 14C ages
during this interval are similar to those from this study, consistent with reduced regional
stratification and local deep water formation.
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