Introduction
Oxygen is vital to all aerobic life. Oxygen solubility in seawater is highly
temperature dependent, with salinity playing a secondary role. The [O2]
of a (deep or intermediate) water mass at a particular location is
determined by its initial concentration at the region of sinking, the amount
of respiration it has undergone, and mixing with other water masses. Both
oxygen supply and consumption are ultimately driven by ocean circulation and
biology (Schmittner et al., 2007). Climate models predict that oxygen
concentrations in the ocean will decrease substantially in response to
anthropogenic climate change. Recent expansion of tropical subsurface oxygen
minimum zones have been attributed to this effect (Stramma et al., 2010).
The warming effect on [O2] loss is twofold: (1) less oxygen can be
dissolved at higher sea water temperatures; (2) warmer surface waters may
increase upper ocean stratification, and it is thought that the resulting
decreased ventilation effect exceeds that associated with reduced oxygen
utilization (Sarmiento et al., 1998; Matear et al., 2000; Plattner et al.,
2001; Bopp et al., 2002; Keeling and Garcia, 2002; Keeling et al., 2010). A
global ocean decline in [O2] between 1 and 7 % has been predicted over
the next century (Keeling et al., 2010); over longer timescales (e.g. 100s–1000s of years) a slowdown in ocean overturning has been predicted to
potentially cause an overall decrease in [O2] of 30 %, with declines
in the deep ocean projected to between 20 and 40 % by the year 2800
(Matear and Hirst, 2003; Schmittner et al., 2008; Shaffer et al., 2009).
However, there are large uncertainties associated with coarse-resolution
ocean models in simulating today's and also future [O2] distributions
(e.g. Jin and Gruber, 2003).
The future reduction in ocean overturning is mainly attributed to changes in
surface heat flux and to a lesser extent to surface freshening (Gregory et
al., 2005). Beyond the last couple of decades there are no direct
observations of deep water [O2]. However, palaeoceanographic proxies of
overturning circulation and ocean ventilation as well as redox proxies
provide constraints of changes in deep water [O2] in relation to
specific climatic events.
The effects of large-scale changes in Atlantic circulation on deep water
[O2] are probably best studied during the last glacial period, which
was punctuated by a series of millennial-scale cold events associated with
the advance of large-scale iceberg armadas (Bond and Lotti, 1995) and
thought to involve systematic changes in the northward heat transport
associated with the Atlantic Meridional Ocean Circulation (AMOC) (Stocker
and Johnson, 2003; Barker et al., 2011). Nutrient proxies (benthic
foraminiferal carbon isotopes (δ13C) and Cd/Ca) and ocean
circulation proxies (Pa/Th, 14C ventilation times) (McManus et al.,
2004; Hoogakker et al., 2007; Skinner et al., 2010), provide evidence for
increased deep water nutrients and reduced ventilation and overturning
circulation in the North Atlantic during cold stadial events
(Schmittner and Lund, 2015) and point to decreased deep water [O2].
Redox sensitive proxies are particularly useful to assess qualitative
changes in bottom water [O2] (Nameroff et al., 2002; Pailler et al.,
2002; Jaccard et al., 2009). Recently, Hoogakker et al. (2015) refined a
novel proxy originally proposed by McCorkle and Emerson (1988), where bottom
water [O2] can be reconstructed from the carbon isotope gradient
between bottom water and pore water at the anoxic boundary. Hoogakker et al. (2015)
suggest that bottom water [O2] in the deep northeast Atlantic
(3.1 km) were 45 and 65 µmol kg-1 lower during the last and penultimate
glacials relative to today. Their reconstructions also showed significantly
reduced bottom water [O2] during extreme cold events associated with
large-scale ice rafting and the deposition of ice-rafted debris in the North
Atlantic (Hoogakker et al., 2015). Here we discuss the underlying causes for
millennial-scale reductions in bottom water [O2] in the deep (3.1 km)
North Atlantic. In addition we present new, millennial-scale-resolved,
bottom water [O2] reconstructions in the North Atlantic from the
intermediate depth (1.8 km) core ODP (Ocean Drilling Program) Site 1055, located on the Carolina
Slope off North America.
Methods
Age models
The age models of both cores were constructed by correlating planktonic
(surface dwelling) foraminiferal oxygen isotope (δ18Op)
records with North Greenland Ice Core Project (NGRIP) δ18Oice (NGRIP Members, 2004). Both marine and ice core records
show a series of oscillating cycles of rapid warmings followed by gradual
cooling (e.g. Dansgaard–Oeschger cycles), culminating in extreme cold events
that are associated with the deposition of massive layers of ice-rafted
debris (IRD) in the North Atlantic (e.g. Heinrich events) (Heinrich, 1988;
Johnsen et al., 1992; NGRIP project Members, 2004; Shackleton et al., 2000,
2004). Typically, six Heinrich layers, H1–H6, have been described for
Marine Isotope Stage (MIS) 3 (29–60 ka BP, before present) and a further five, H7–H11,
over MIS 4 and 5 (between 60 and 130 ka BP). However, outside the
Labrador Sea such IRD layers contain conspicuously less detrital carbonate
(a defining criterion for a Heinrich layers) and are labelled C19–C25
(Chapman and Shackleton, 2002). For the interval 0–60 ka the GICC05 (Greenland Ice Core Chronology 2005) age
model was applied, whose ages are very similar to that of the SFCF 2004 age
model as was used previously in Hoogakker et al. (2015). Thornalley et al. (2013)
apply a revised chronology prior to 60 ka, based on the
speleothem-tuned age model of Barker et al. (2011), and to aid comparison
between the two sites the same chronology was applied to MD95-2042 between
60 and 123 ka. Based on these age models, results of core MD95-2042 cover
the last 150 kyr, whilst those of core ODP Site 1055 cover the interval 85–59 ka BP (Fig. 2).
Age models of MD95-2042 and ODP 1055 established by tying planktonic
foraminifera oxygen isotope changes of Globigerina bulloides
(MD95-2042; Shackleton et al., 2000) and Globigerinoides ruber (ODP Site 1055) to
those of NGRIP. Several Dansgaard-Oeschger interstadial events are numbered
in the NGRIP records.
Sea water [O2]
The biogeochemical cycles of oxygen and carbon are stoichiometrically linked
through photosynthesis and respiration. Photosynthesis uses carbon dioxide
(CO2), water, sunlight and nutrients to make organic material and
oxygen. The breakdown of organic material, in well-oxygenated environments,
uses oxygen and produces CO2. During photosynthesis, organisms
preferentially take up 12C, which is isotopically light compared with 13C, causing an
overall enrichment of the carbon isotopic composition (δ13C) of
DIC (dissolved inorganic carbon)
in surface waters (Kroopnick, 1985; Gruber et al., 1999). When organic
material is broken down, the release of light 12C causes a depletion in
seawater δ13C DIC. Globally there is a strong linear
relationship between deep water [O2] and δ13C, where a
50 µmol kg-1 decrease in [O2] corresponds to a
0.34 ‰ decrease in seawater δ13C DIC (Fig. 3),
with R2 between 0.78 and 0.85. However, within the North and South
Atlantic and the Southern Ocean the data are distributed within a cloud,
displaying a much weaker relationship. Some of the increased variability in
[O2] in the Atlantic basins and Southern Ocean is probably related to
seawater temperature differences; colder seawater can contain more dissolved
oxygen but also mixing of water masses. Furthermore, δ13C-DIC
distributions in the oceans are also affected by temperature-dependent
fractionation during air–sea gas exchange (Lynch-Stieglitz et al., 1995) and
degree of surface water equilibration with the atmosphere (Schmittner et
al., 2013) at source waters, biology, and also mixing with other water masses
(Gruber et al., 1999). During glacial times bottom water δ13C
estimates derived from benthic foraminiferal calcite δ13C in
the deep (> ∼2.5 km) Atlantic became more depleted
(Curry and Oppo, 2005; Oliver et al., 2010), but uncertainties related to
preformed δ13C, air–sea fractionation, and terrestrial biomass
contributions to deep water δ13C-DIC precludes the use of
bottom water δ13C-DIC inferred from benthic foraminifera in the
past as a reliable bottom water [O2] proxy.
(a) Global relationship between seawater [O2] and
δ13C of DIC. (b) Cross-plots of seawater [O2] and
δ13C for intermediate (1000–1500 and 1500–2000 m) and deep
(2000–3000 and 3000–4000 m) waters. The data used to create this figure can be
found in the Supplement and were obtained from
http://www.nodc.noaa.gov/OC5/SELECT/dbsearch/dbsearch.html. Only World Ocean
Database quality-controlled data with accepted values (e.g. flag 0) are included.
Here we apply the refined bottom water to pore water (at the anoxic
boundary) δ13C gradient as a quantitative bottom water
[O2] proxy (Hoogakker et al., 2015). This proxy was originally proposed
by McCorkle and Emerson (1988), who observed that the carbon isotope gradient
between bottom water and pore water at the anoxic boundary ([O2]=0)
in sediments decreases with decreasing bottom water [O2]. These changes
are attributed to changes in the amount of organic material that can be
remineralized; e.g. more organic material can be remineralized under higher
bottom water [O2], releasing more 12C into the pore waters,
increasing the bottom water to anoxic pore water δ13C gradient
(Δδ13Cbw-ab_pw), as supported by
pore water δ13C and [O2] models (McCorkle and Emerson,
1988; Gehlen et al., 1999). Hoogakker et al. (2015) furthermore show that
additional observations of Δδ13Cbw-ab_pw, inferred
from the difference in δ13C between bottom water
and foraminifera living at the anoxic boundary (Globobulimina spp.) as well as between
bottom water suspension feeding foraminifera (Cibicidoides wuellerstorfi) and anoxic boundary dwelling
foraminifera (Globobulimina spp.), all fit the original observations exceptionally well at
[O2] between 55 and 235 µmol kg-1. At higher
(> 235 µmol kg-1) [O2], additional light carbon is added to the pore water
from other remineralization reactions. These observations confirm that
δ13Cbw-ab_pw can be approximated by the
δ13C difference between the test carbonate δ13C of
benthic foraminiferal species that live in bottom water (e.g. C. wuellerstorfi) and in the
sediment at the dysoxic/anoxic boundary (e.g. Globobulimina spp.) at bottom water
[O2] values of 55–235 µmol kg-1, where a 0.39 ‰
increase in Δδ13Cbw represents a 50 µmol kg-1
increase in bottom water [O2] (Hoogakker et al., 2015). Hereafter we refer
to this carbon isotope gradient simply as Δδ13C.
Benthic foraminifera Δδ13C at deep site MD95-2042
and intermediate ODP Site 1055 and their planktonic foraminifera oxygen
isotopes. Original benthic foraminifera δ13C records (MD95-2042
from Shackleton et al., 2000) of epifaunal C. wuellerstorfi (red
circles) and deep infaunal G. affinis (blue circles) are also shown
intercalated between the Δδ13C records. Several Heinrich
events and cold events are shown.
Reconstructed bottom water [O2] at deep site MD95-2042 and
intermediate ODP Site 1055 shown with their planktonic foraminifera oxygen
isotope records (Shackleton et al., 2000; Thornalley et al., 2013). Heinrich
events 1, 3, 4, 5, 5a, and 6 and cold events 19, 20, and 21 are shown.
Results
Both records show relatively well-oxygenated water masses for the periods
covered, with Δδ13C values of 1.45 ‰
and higher (Fig. 4) amounting to bottom water [O2] of 144 µmol kg-1
and higher (Fig. 5). Typically, seawater is considered hypoxic when
[O2] values of 60 µmol kg-1 or less are recorded, although the median
lethal [O2] varies between different organisms; temperature and
CO2 also influence this threshold (Keeling et al., 2010). At MD95-2042,
the LGM (last glacial maximum), MIS 6, and extreme cold events are associated with lower [O2]
(Hoogakker et al., 2015), with Heinrich event 4 showing the lowest Δδ13C and thus bottom water [O2] (Fig. 4). At the
shallower northwest Atlantic ODP Site 1055, MIS 4 and cold events C19, C20,
and C21 are associated with a lower Δδ13C and bottom water
[O2]. From ∼62 ka BP there is gradual increase in
Δδ13C, including the latter parts of Heinrich event 6
at ODP Site 1055, although Δδ13C was lower
compared with warm interstadials (Fig. 4).
Hoogakker et al. (2015) calculate that the total error associated with
bottom water [O2] reconstructions using this method at mid to low
latitudes is 17 µmol kg-1. This error includes uncertainties associated
with variations in the δ13C of organic carbon of ±1 ‰ (see Hoogakker et al., 2015 Supplement
for details), which seems a reasonable assumption for the low to mid
latitude ocean (Goericke and Fry, 1994). Because of decreased [CO2(aq)]
during full glacial conditions, δ13Corg was
enriched by 2 ‰ (Rau et al., 1991) causing an initial
overestimation of glacial bottom water [O2] and correction of
10 µmol kg-1 (Hoogakker et al., 2015). The study of Rau et al. (1991) is of too
low resolution to decipher any possible millennial-scale oscillations in
δ13Corg; however, generally δ13Corg appears
lighter prior to the LGM. It is also important to note that within the North
Atlantic Heinrich belt, organic carbon δ13C values are depleted
during glacial times compared to the Holocene, with the lightest values (up to
-28 ‰) during Heinrich events 4, 2, and 1 (Huon et al., 2002;
Schouten et al., 2007). Both Huon et al. (2002) and Schouten et al. (2007)
attribute these depletions in organic δ13C to increased input
of terrestrial organic material from either ice-rafted debris or wind-blown
sources. It is therefore possible that estimates of [O2] during
Heinrich events and cold events C20 and C21 are overestimated. However, as
terrestrial plant remains are generally much older in age (Schouten et al.,
2007), it is possible that they are largely refractory (insoluble and
non-hydrolyzable) and may not have degraded substantially. Because of this
unknown we consider estimates of bottom water [O2] during these
Heinrich events and cold events C20 and C21 as maximum estimates (Fig. 5).
Lowest bottom water [O2] associated with Heinrich and extreme
cool events and the difference with modern time at intermediate North Atlantic ODP
Site 1055 (254 µmol kg-1 today) and deep North Atlantic site
MD95-2042 (245 µmol kg-1 today). Note that [O2] at
MD95-2042 during cold event C20 (indicated with *) is not significantly
different from that at modern time.
Bottom water [O2] in µmol kg-1 (±17 µmol kg-1)
Event
ODP 1055
Diff. with modern
MD95-2042
Diff. with modern
C21
178
76
C20
230
24
230
15*
C19
213
41
H6
224
30
170
75
H5a
206
39
H5
209
35
H4
144
101
H3
181
64
H1
166
79
Discussion
Millennial-scale climate oscillations are a common feature of the last
glacial as well as the transition from the previous interglacial (Eemian) to
glacial in the North Atlantic (Fig. 2). Within the North Atlantic IRD
belt, ice rafting becomes a common feature during millennial-scale cooling
events when sea level falls below -45 m (Chapman and Shackleton, 2002).
Decreased benthic foraminiferal δ13C from deep (below 2.5 km)
sites in the North Atlantic provide evidence for widespread changes in
bottom water carbonate chemistry during these events (Shackleton et al.,
2000; Sarnthein et al., 2000; Chapman and Shackleton, 2002; Thornalley et
al., 2013). Reconstruction of [CO32-] support the inferred changes
in deep bottom water carbonate chemistry (Yu et al., 2008). High-resolution
intermediate depth North Atlantic records from the northeast Atlantic also
generally show lower benthic δ13C during Heinrich events
(Sarnthein et al., 2000; Chapman and Shackleton, 2002; Rasmussen et al.,
2003; Peck et al., 2006; Dickson et al., 2008; Thornalley et al., 2010)
whereas ODP Site 1055 from the northwest Atlantic, featured here, shows
hardly any change (Evans and Hall, 2008; Thornalley et al., 2013). During
glacial times, reconstructed [CO32-] at North Atlantic sites above
2.8 km all show increased concentrations (Yu et al., 2008); to date, no
inferences have been made with regards to millennial-scale climate
oscillations.
During most of MIS 5, including the transition to
glacial conditions, the deep northeast Atlantic was well oxygenated (Fig. 5).
Between 126 and 109 ka BP G. affinis was absent, probably because a reduced
organic carbon flux and deep or weakly developed anoxic boundary meant its
microhabitat conditions were not met, similar to Holocene conditions
(Hoogakker et al., 2015). Following this period Δδ13C
is > 2.25 ‰, indicating well-oxygenated
(> 235 µmol kg-1) waters. It is not until after
∼76 ka BP, coincident with Atlantic cold event C20, that
Δδ13C of < 2.25 ‰ are
measured (Fig. 4). Applying the Δδ13C : [O2]
calibration equation of Hoogakker et al. (2015), we calculate that during
Atlantic cold event C20 bottom water [O2] at the Iberian Margin was
230±17 µmol kg-1 (Fig. 5, Table 1). Note that if we had used
the present-day δ13C : [O2] relationship as defined in
Fig. 3, bottom water [O2] would be drastically underestimated, with
bottom water [O2] of ∼120 µmol kg-1 during event
C20. At the Blake Ridge location (ODP Site 1055), Δδ13C
fell below 2.25 ‰ during North Atlantic cold events C21
and C20, giving bottom water [O2] of 179 and 230±17 µmol kg-1
respectively (Table 1). Interestingly, during North Atlantic cold
event C20 both the deep northeast Atlantic record and intermediate northwest
Atlantic record show the same bottom water [O2] (Fig. 5).
Between 76 and 64 ka BP, roughly coincident with MIS 4, the record of
MD95-2042 does not resolve millennial-scale oscillations, mainly because C. wuellerstorfi
was not abundant during this time. In the few instances it did occur Δδ13C was > 2.25 ‰, suggestive of
well-oxygenated conditions. At the intermediate depth ODP Site 1055 Δδ13C follows G. ruber δ18O, where lighter δ18O values are associated with Δδ13C > 2.25 ‰
and heavier δ18O values, corresponding to millennial-scale cool events, with Δδ13C < 2.25 ‰. During North Atlantic cold
event C19, reconstructed bottom water [O2] at ODP Site 1055 was
213±17 µmol kg-1, and the cold period that follows is
characterized by bottom water [O2] of 194±17 µmol kg-1
(Fig. 5, Table 1).
During the later part of MIS 4 and MIS 3, the deep record of MD95-2042 is
characterized by bottom water [O2] variations that follow Greenland
climate trends, with high Δδ13C (> 2.25 ‰) values during interstadials, whereas low bottom
water [O2] characterize Heinrich events, with H6, H4, and H1 showing
lowest bottom water [O2] of 170, 144, and 166±17 µmol kg-1
respectively (Fig. 5). Obviously these values still mean well-oxygenated
bottom water masses, but they are lower compared with warm interstadial
intervals (> 235 µmol kg-1) as well as the LGM (200±17 µmol kg-1). At the intermediate location ODP Site 1055, early H6 shows
slightly lower bottom water [O2] of 224±17 µmol kg-1
followed by an increase to > 235 µmol kg-1 (Fig. 5,
Table 1).
Causes for millennial-scale bottom water
[O2] changes
The glacial decreased bottom water [O2] values at the Iberian Margin to
200±17 µmol kg-1 (LGM) and 180±17 µmol kg-1 (MIS 6)
(compared with 245 µmol kg-1 today) have been largely attributed to
ocean circulation changes, with a shift in bottom water mass from NADW to
SSDW (Hoogakker et al., 2015).
Over the transition from MIS 5 to early MIS 4, a mode change has been
suggested in the AMOC
(Bereiter et al., 2012; Thornalley et al., 2013; Barker and Diz, 2014,
Böhm et al., 2015). Bereiter et al. (2012) suggest that during MIS 5
AMOC was strong, characterized by southward flow of NADW to the deep South
Atlantic. This would imply that NADW and NAIW (North
Atlantic Intermediate Water) influenced bottom
waters at the deep and intermediate sites respectively. Several studies have
shown that most cold events within MIS 5 are associated with decreased
benthic foraminifera δ13C (Shackleton et al., 2000; Oppo et
al., 2001; Evans and Hall, 2008; Hodell et al., 2009), which have often been
interpreted to reflect AMOC changes. Guihou et al. (2010), using the
kinematic overturning circulation proxy 231Pa /230Th, show that
AMOC export from the North Atlantic was reduced during the cold events of
MIS 5. However, Guihou et al. (2011) further show that cold events within MIS
5 and MIS 4 could be associated with stronger AMOC export at shallow depths,
which agrees with grain size results of Thornalley et al. (2013), suggesting
more vigorous near-bottom flow speeds during millennial cold events at
the intermediate ODP Site 1055. These results confirm inferences of possible
strengthened open ocean convection south of the Greenland–Scotland Ridge
driving a strong intermediate depth Atlantic Overturning Circulation cell
(Thornalley et al., 2013). It would then be somewhat surprising to find
lower bottom water [O2] during these events as more vigorous North
Atlantic Intermediate Water flow is generally associated with better
ventilation, although changes in the mode of water mass formation can alter
the extent to which newly formed intermediate/deep waters have equilibrated
with the atmosphere.
During the glacial, AMOC was considerably different. Rahmstorf (2002)
proposed, based on a benthic foraminifera δ13C synthesis of
Sarnthein et al. (1994), that a deep North Atlantic overturning cell with
active deep and intermediate water formation in the North Atlantic and
Greenland–Iceland–Norwegian (GIN) seas occurred during warm interstadials,
active intermediate convection occurred during stadial events, whereas
Heinrich events were associated with a significant reduction in overturning
strength. Using 231Pa /230Th as a kinematic overturning proxy,
McManus et al. (2004) suggest that the meridional overturning circulation
was significantly reduced during Heinrich Stadial 1. However, the picture
appears more complicated. Bottom flow speed reconstruction from the deep
(3.5 km) northwest Atlantic suggests that flow speed changes at this depth
follow an Antarctic temperature signal, showing slowdowns in bottom flow
vigour coincident with Antarctic warmings (Hoogakker et al., 2007),
which have also been linked with bottom water changes (Gutjahr et al.,
2010). Both Hoogakker et al. (2007) and Roberts et al. (2010) suggest that
perturbations associated with millennial cool events likely only influenced
the shallow overturning cell in the North Atlantic. 231Pa /230Th
reconstructions covering the intermediate northeast Atlantic over H1 however
do not show evidence for a weakened shallow overturning cell (Gherardi et
al., 2009). Since then it has emerged that glacial Antarctic Bottom Waters
and glacial Antarctic Intermediate Waters might show a see-saw pattern in
the North Atlantic during Heinrich events, where deep waters show an
increase in the contribution of high-nutrient, low-[O2] glacial
Antarctic Bottom Waters, and intermediate waters show a decreased
contribution of Antarctic Intermediate Water and increased contribution of
possibly well-ventilated high-[O2] Glacial North Atlantic Intermediate
Water (Gutjahr et al., 2008, 2010; Huang et al., 2014;
Piotrowski et al., 2005). Whilst changes in bottom water mass may thus have
some part to play in the bottom water [O2] changes at deep sites during
Heinrich events, they cannot however explain lower bottom water [O2] at
the intermediate depth site.
In terms of biological mechanisms driving North Atlantic seawater [O2]
changes during Heinrich events, the picture is not clear. Model simulations
suggest that export production during Heinrich events was globally reduced
(Schmittner et al., 2005; Mariotti et al., 2012; Menviel et al., 2014).
Interestingly, while Mariotti et al. (2012) suggest an overall decrease in
export production in the North Atlantic, model simulations by
Menviel et al. (2014) show increases across large areas in the Atlantic. According to
Salguiero et al. (2010) there were no changes in productivity in the
northeast Atlantic at MD95-2042. However, for the subtropical northeast
Atlantic, McKay et al. (2014) inferred increased primary production in
surface waters during H1, causing low-oxygen conditions in the underlying
(2.5 km) sediments. Furthermore, several studies from deep locations in the
Atlantic, including the Blake Outer Ridge, Bermuda Rise, the Tobago Basin and
equatorial region have documented conspicuous increases in opal sediments
during Heinrich events and extreme cold events of MIS 5 (Hoogakker et al.,
2007; Keigwin and Boyle, 2008; Gil et al., 2009; Griffiths et al., 2013;
Meckler et al., 2013). This could imply a change in productivity at
oligotrophic gyre locations in the North Atlantic with increased
contribution from opal producers, possibly at the expense of carbonate
(foraminifera, coccolith, pteropod, aragonite) producers (Brzezinski et al.,
2002; Griffiths et al., 2013). Recent work by Hoogakker et al. (2013)
suggests weaker summer stratification in the northwest Atlantic during H5,
which could be associated with a deeper mixed layer potentially enhancing
silicate available to surface waters. In combination with an increased dust
flux (López-Martinez et al., 2006), iron fertilization could have supported
diatom productivity. More importantly, whilst export of diatoms to the deep
ocean is not that efficient, accumulation of diatom deposits in sediments
during Heinrich events (Lippold et al., 2009; Griffiths et al., 2013) could
provide evidence that more organic-rich material was exported to greater
water depths during these episodes. Based on this evidence we propose that
the lower bottom water [O2] values at intermediate ODP Site 1055 during
extreme millennial-scale cool events were driven by increased export
production. The model simulation of Marriotti et al. (2012) and Menviel et
al. (2014) also suggests an increase in South Atlantic export production, in
agreement with an earlier proxy study by Anderson et al. (2009). In their
study, Anderson et al. (2009) found the highest opal fluxes in the Southern
Ocean, which were coincident with bottom [O2] minima at MD95-2042 of H6,
H4, and H1. This implies that biological mechanisms also played a role in
decreasing bottom water [O2] at the deep site, either by changing the
[O2] of SSDW in the Southern Ocean or through increased export across
the Atlantic.
Our reconstructed bottom water [O2] changes across Heinrich events and
extreme cool events of MIS 5 agree with a modelling study of Schmittner et
al. (2007), who show that intermediate and deep waters of the North Atlantic
were associated with lower bottom [O2] during such events.
Although the UVic (University of Victoria) model simulations depict the main features of modern
oxygen distributions, the North Atlantic results have higher values than
observations, whereas large parts of the South Atlantic and Indian/Pacific
oceans have lower [O2] values compared with observations (Schmittner et al.,
2007). Furthermore, while compared with modern time, the model simulations of
Schmittner et al. (2007) predict a decrease in bottom water [O2] of
60–90 µmol kg-1 at the longitude of the intermediate site 1055 and
90–120 µmol kg-1 at the longitude of deep site MD95-2042 during meltwater
events, our reconstructions suggests more modest decreases in the range of
24–76 µmol kg-1 (9–30 %) for the intermediate site, and
15–101 µmol kg-1 (5–40 %) at the deep site (Fig. 5). The larger
amplitude changes in seawater [O2] simulated by Schmittner et al. (2007)
may be the result of the prescribed pre-industrial boundary
conditions with strong AMOC; had they used glacial boundary conditions
with weaker AMOC, the oxygen changes at the deep site might have been
smaller. However, it is noted that the model outputs depict a particular
(extreme) point in model time, whereas reconstructions from deep sea
sediments represent an averaged view where extremes have been smoothed out
by bioturbation. Our reconstructions agree with model simulations suggesting
an overall decrease in North Atlantic [O2] during glacial
millennial-scale cold events.