BGBiogeosciencesBGBiogeosciences1726-4189Copernicus PublicationsGöttingen, Germany10.5194/bg-14-5217-2017Carbon uptake and biogeochemical change in the Southern Ocean, south of
TasmaniaPardoPaula CondeTilbrookBrontehttps://orcid.org/0000-0001-9385-3827LanglaisClothildehttps://orcid.org/0000-0002-6423-7678TrullThomas WilliamRintoulStephen RichAntarctic Climate and Ecosystem Cooperative Research Centre,
University of Tasmania, Hobart, AustraliaClimate Science Centre, CSIRO Oceans and Atmosphere, Hobart, AustraliaPaula C. Pardo (paula.condepardo@csiro.au)21November201714225217523729May20177June201727September20172October2017This work is licensed under the Creative Commons Attribution 3.0 Unported License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/3.0/This article is available from https://bg.copernicus.org/articles/14/5217/2017/bg-14-5217-2017.htmlThe full text article is available as a PDF file from https://bg.copernicus.org/articles/14/5217/2017/bg-14-5217-2017.pdf
Biogeochemical change in the water masses of the Southern Ocean, south of
Tasmania, was assessed for the 16-year period between 1995 and 2011 using
data from four summer repeats of the WOCE–JGOFS–CLIVAR–GO-SHIP (Key et
al., 2015; Olsen et al., 2016) SR03 hydrographic section (at
∼ 140∘ E). Changes in temperature, salinity, oxygen, and
nutrients were used to disentangle the effect of solubility, biology,
circulation and anthropogenic carbon (CANT) uptake on the
variability of dissolved inorganic carbon (DIC) for eight water mass layers
defined by neutral surfaces (γn). CANT was
estimated using an improved back-calculation method. Warming
(∼ 0.0352 ± 0.0170 ∘C yr-1) of Subtropical
Central Water (STCW) and Antarctic Surface Water (AASW) layers decreased
their gas solubility, and accordingly DIC concentrations increased less
rapidly than expected from equilibration with rising atmospheric CO2
(∼ 0.86 ± 0.16 µmol kg-1 yr-1 versus
∼ 1 ± 0.12 µmol kg-1 yr-1). An increase in
apparent oxygen utilisation (AOU) occurred in these layers due to either
remineralisation of organic matter or intensification of upwelling. The range
of estimates for the increases in CANT were 0.71 ± 0.08 to
0.93 ± 0.08 µmol kg-1 yr-1 for STCW and
0.35 ± 0.14 to 0.65 ± 0.21 µmol kg-1 yr-1
for AASW, with the lower values in each water mass obtained by assigning all
the AOU change to remineralisation. DIC increases in the Sub-Antarctic Mode
Water (SAMW, 1.10 ± 0.14 µmol kg-1 yr-1) and
Antarctic Intermediate Water (AAIW,
0.40 ± 0.15 µmol kg-1 yr-1) layers were similar
to the calculated CANT trends. For SAMW, the CANT
increase tracked rising atmospheric CO2. As a consequence of the general
DIC increase, decreases in total pH (pHT) and aragonite
saturation (ΩAr) were found in most water masses, with the
upper ocean and the SAMW layer presenting the largest trends for
pHT decrease (∼-0.0031 ± 0.0004 yr-1). DIC
increases in deep and bottom layers
(∼ 0.24 ± 0.04 µmol kg-1 yr-1) resulted
from the advection of old deep waters to resupply increased upwelling, as
corroborated by increasing silicate
(∼ 0.21 ± 0.07 µmol kg-1 yr-1), which also
reached the upper layers near the Antarctic Divergence
(∼ 0.36 ± 0.06 µmol kg-1 yr-1) and was
accompanied by an increase in salinity. The observed changes in DIC over the
16-year span caused a shoaling (∼ 340 m) of the aragonite saturation
depth (ASD, ΩAr= 1) within Upper Circumpolar Deep Water
that followed the upwelling path of this layer. From all our results, we
conclude a scenario of increased transport of deep waters into the section
and enhanced upwelling at high latitudes for the period between 1995 and 2011
linked to strong westerly winds. Although enhanced upwelling lowered the
capacity of the AASW layer to uptake atmospheric CO2, it did not limit
that of the newly forming SAMW and AAIW, which exhibited CANT
storage rates (∼ 0.41 ± 0.20 mol m-2 yr-1) twice
that of the upper layers.
Introduction
The Southern Ocean is a key region in terms of climate change and climate
variability, influencing the Meridional Overturning Circulation (MOC) and therefore
modulating the global circulation and oceanic biogeochemical cycles
(Sarmiento et al., 1998, 2004; Orr et al., 2005). Deep waters, formed in the
North Atlantic, spread south and enter the Southern Ocean, where they mix
with deep layers of the Antarctic Circumpolar Current (ACC) and ultimately
upwell between the Southern ACC Front and the Polar Front. The upwelled
waters are eventually transformed into bottom, intermediate and mode waters,
which are exported from the Southern Ocean to ventilate the thermocline and
bottom layers of the major ocean basins. Some of the Southern Ocean waters
subducted into the ocean interior return to the North Atlantic to balance the
southward flux of North Atlantic Deep Water (Speer et al., 2000; Lumpkin and
Speer, 2007; Iudicone et al., 2008).
Within the eastward flow of the ACC, major water exchange between the three
ocean basins takes place. The circumpolar path of the ACC consists of
various narrow jets associated with sharp fronts that separate waters with
different characteristics (Orsi et al., 1995; Belkin and Gordon, 1996).
These jets can reach deep layers and often meander, intensify, merge and
split, conditioned by the topography of the ocean floor, the stratification
of the ACC, and atmospheric variability (Moore et al., 1999; Sokolov and
Rintoul, 2002, 2009; Peña-Molino et al., 2014). Movements in the jets
enhance cross-stream transports and mesoscale activity that can result in
local changes in water mass properties and this may complicate the
computation of long-term changes in water mass properties (Rintoul and
Bullister, 1999; Sallée et al., 2008; Peña-Molino et al., 2014).
Water mass formation and ventilation transport heat, salt, and dissolved
gases from the atmosphere to the ocean interior and other basins (Sarmiento
et al., 2004), with the Southern Ocean contributing ∼ 40 % to the
anthropogenic CO2 (CANT) inventory of the ocean (Sabine et
al., 2004; Gruber et al., 2009; Khatiwala et al., 2009). Circulation and
biological processes drive the redistribution of dissolved inorganic carbon
(DIC) that ultimately affects the capacity of the waters to uptake more
CO2. The uptake of CO2 by the Southern Ocean presents strong
spatiotemporal variability (Lenton et al., 2013), and this can lead to
conflicting results for observational studies, models and atmospheric
inversions, depending on the methodology used (Verdy et al., 2007; Lenton et
al., 2012; Fay et al., 2014). Quantifying long-term changes in the carbon
system is difficult due to the scarcity of data (Lenton et al., 2012;
Kouketsu and Murata, 2014; Fay et al., 2014) and the influence of biological
processes. Notably, long-term trends in CANT concentration are
difficult to estimate due to its small signal (∼ 3 %) with respect
to that of DIC in the ocean.
Once CO2 dissolves in the ocean (DIC) it begins the process of ocean
acidification, i.e. decreases the pH and the saturation state of calcium
carbonate (CaCO3) minerals such as calcite and aragonite (Feely et al.,
2004; Bates et al., 2014), with potential to disrupt ecosystems and
biological processes (Doney et al., 2009).
Numerous studies have documented warming and freshening of deep and bottom
layers of the Southern Ocean in recent decades (see reviews by Jacobs, 2006
and van Wijk and Rintoul, 2014). The abyssal waters of the
Australian–Antarctic Basin (A–AB) show the greatest freshening in the last
40 years (van Wijk and Rintoul, 2014). This freshening was accompanied by a
warming of the deep–bottom layers, leading to a contraction in bottom waters
by more than half their volume in the basin (Purkey and Johnson, 2012; van
Wijk and Rintoul, 2014). Subsurface to intermediate layers have also warmed
and freshened south of Australia (Bindoff and Church, 1992; Wong et al.,
1999; Aoki et al., 2005). The solubility of gases and their distribution in the
ocean depend on the dynamics and properties of water masses, and the
thermohaline changes that have occurred south of Tasmania (Fig. 1) have
implications in the carbon system of the region.
Transects of the four summer repeats of the G0-SHIP hydrodynamic line
SR03 in the Southern Ocean south of Tasmania for the period 1995–2011 and
the main hydrography features of the region. ZC = Zeehan Current.
EAC = East Australian Current. Black arrows indicate the flow of the
Antarctic Circumpolar Current (ACC), with STF = Subtropical Front;
SAF = Sub-Antarctic Front, PF = Polar Front and SACCF = Southern
ACC Front.
A reduction of the carbon sink of the Southern Ocean was observed between the
1980s and the early 2000s (Le Quéré et al., 2007; Lovenduski et
al., 2008), with more recent studies suggesting a recovery and even
intensification of the CO2 uptake by 2011–2012 (Zickfeld et al., 2008;
Fay et al., 2014; Landschützer et al., 2015). DeVries et al. (2017), used
a global inverse model to postulate that changes in circulation are
responsible for most of the variability in the oceanic CO2 uptake, with
the weakening of the upper-ocean circulation being responsible for the
increase in oceanic carbon uptake over the past decade. Ocean acidification
has been observed in the whole Southern Ocean (Lauvset et al., 2015) and
locally in the Atlantic and Pacific sectors (Williams et al., 2015; Hauri et
al., 2015). South of Tasmania, McNeil et al. (2001) reported an increase in
CANT uptake between 1968 and 1996 for the region
45–50∘ S. These authors reported, for the first time,
CANT accumulation in the AABW and highlighted the importance of
the formation of bottom and mode waters as a mechanism for transporting
CANT to the ocean. In terms of ocean acidification, we are not
aware of any study about trends in ocean acidification in the water masses
south of Australia.
Considering the lack of observational estimates for recent biogeochemical
changes in the A–AB as well as the large changes in CO2 uptake and
storage suggested by recent atmospheric and surface observations in the
Southern Ocean (e.g. Fay et al., 2014; Landschützer et al., 2015) there
is a need to provide a full ocean depth observational perspective on how the
ocean is changing. The aim of this paper is to provide the first estimates of
biogeochemical change in the water masses south of Tasmania, for the period
1995–2011, disentangling the effects that solubility, circulation, biology
and CANT uptake have on the variability of DIC. We use data from
four summer repeats of the (Key et al., 2015; Olsen et al., 2016)
WOCE–JGOFS–CLIVAR–GO-SHIP hydrographic section SR03 (Fig. 1; Table 1), one
of the most revisited sections in the Southern Ocean. Trends in oxygen
(O2), nutrients, and the carbon system parameters (i.e. DIC, total
alkalinity (TA), anthropogenic carbon (CANT), total pH
(pHT) and percentage aragonite saturation (ΩAr)) were estimated for the period 1995–2011, when both DIC and
TA measurements are available. CANT estimates were obtained with
a back-calculation method (Pardo et al., 2014). The changes were evaluated in
the different water mass layers of the section defined by neutral surfaces
(γn, McDougall et al., 1987).
Repeats of the GO-SHIP SR03 line on board the Aurora
Australis from 1995 to 2011 with the expocode from GLODAPv2 database.
EXPOCODEDates (dd/mm/yyyy)REF.09AR1994121320/12/1994–1/02/1995199509AR2001102929/10/2001–11/12/2001200109AR2008032223/03/2008–15/04/2008200809AR201101244/01/2011–31/01/20112011Hydrography of the region
The dynamical structure of the region south of Tasmania (Fig. 1) is
characterized by a number of fronts that separate the major water masses of
the region (Sokolov and Rintoul, 2002, 2007, 2009). At the northern end of the
section, the presence of the weak Subtropical Front (STF, Fig. 1) separates
warm, salty subtropical surface waters from cooler and fresher sub-Antarctic
surface waters (Deacon, 1937). The northern end of the section is a complex
mixing zone where waters transported down the east coast of Tasmania in a
series of mesoscale eddies from the East Australian Current (EAC, Fig. 1) mix
into the Subantarctic Zone, and also meet Zeehan Current (ZC, Fig. 1) waters
transported down the west coast of Tasmania (Boland and Church, 1981; Baines
et al., 1983; Speich et al., 2002; Davis, 2005; Ridgway et al., 2007; Sloyan
et al., 2016). The EAC transported south of Tasmania forms a zonal jet
towards the south-east Indian Ocean, known as the Tasman Outflow, that reaches
the bottom of the Tasman slope and that is maintained all year round (Rintoul
and Bullister, 1999; Ridgway et al., 2007). The encounter between these
currents presents high variability at the northern end of the section (Fig. 1).
Farther south (Fig. 1), the Sub-Antarctic Front (SAF) and the Polar Front
(PF) are regions of maximum transport in the ACC (Rintoul and Bullister,
1999; Sokolov and Rintoul, 2002). North of the SAF, deep winter convection
generates Sub-Antarctic Mode Water (SAMW), a relatively uniform water mass
that occupies subsurface layers down to ∼ 600 m deep (McCartney, 1977;
Rintoul and Bullister, 1999). SAMW constitutes the main part of the upper
limb of the MOC, ventilating the thermocline of all the ocean basins (e.g. Speer
et al., 2000; Sloyan and Rintoul, 2001). The SAF coincides with the deepening
of the salinity minimum at intermediate depths (Whitworth III and Nowlin Jr.,
1987), which is the signature of the Antarctic Intermediate Water (AAIW). The
AAIW underlies the SAMW and also ventilates the global thermocline layers of
the ocean. It is mainly formed in the south-east Pacific and it is
continuously transformed on its way to the region south of Tasmania (Hanawa
and Talley, 2001). South of the SAF, colder and fresher Antarctic Surface
Water (AASW) covers the surface ocean. AASW originates from progressive
warming of Winter Water (WW, Mosby, 1934), which can be perceived even in
summer as a remnant layer of cold water at the base of the AASW (Rintoul et
al., 1997).
The Southern ACC Front (SACCF, Fig. 1) is a deep front located south of the
PF (Orsi et al., 1995) and can coincide with the southern boundary of the
ACC, which is represented by the southern limit of the oxygen minimum (Orsi
et al., 1995). The oxygen minimum is related to the advection of Upper
Circumpolar Deep Water (UCDW) that originates in the Indian and Pacific
Oceans (Callahan, 1972), from where it spreads south, mixes with deep layers
of the ACC and ultimately upwells near the Antarctic continent as part of the
lower cell of the MOC. Lower Circumpolar Deep Water (LCDW) is below UCDW and
is identified by a salinity maximum and contributes to upwelling around
Antarctica. The precursor of the LCDW is North Atlantic Deep Water (NADW),
which originates in the Labrador and Nordic seas of the North Atlantic polar region
(Dickson and Brown, 1984) that flows southward to enter the ACC as part of
the MOC (Callahan, 1972; Orsi et al., 1995; Johnson, 2008).
At the southern end of the section, the Antarctic Slope Front forms the
boundary between cold and fresh shelf water and relatively warm and salty
waters offshore (Jacobs, 1991). Polynyas along the Adélie and George V
Land coast (Fig. 1) contribute to the formation of Adélie Land Bottom
Water (ALBW), which is a mixture of High-Salinity Shelf Water (HSSW)
resulting from brine rejection during ice formation and ultra-modified LCDW
(Foster and Carmack, 1976; Rintoul, 1998; Marsland et al., 2004). ALBW
constitutes ∼ 25 % of the total volume of water < 0 ∘C
in the ocean (Rintoul, 1998). The bottom waters near the southern end of
the section also contain a component of Ross Sea Bottom Water (RSBW) that
originates to the east and is deflected westwards towards the A–AB as it is
transported down the continental slope. The RSBW is modified when it arrives
at the location of the SR03 section by mixing with deep layers of the ACC and
recently formed ALBW (Gordon and Tchernia, 1972; Rintoul, 1998). The ALBW and
the modified RSBW together ventilate the abyssal layers of the A–AB before
spreading north to ventilate the Indian and Pacific basins (Mantyla and Reid,
1995; Fukamachi et al., 2010).
Data and methodData
The SR03 hydrographic section between Tasmania and Antarctica (Fig. 1) was
occupied between 1991 and 2011. Measurements of total alkalinity (TA) were
only available from the beginning of 1995, and our evaluation of
biogeochemical changes is limited to four summer sections occupied for the
period 1995–2011 (Table 1). A winter cruise in 1996 was not considered in
this study in order to minimise seasonal biases.
Water column salinity, temperature, pressure and dissolved oxygen (O2)
were collected from the conductivity–temperature–depth (CTD) device with an
accuracy of ±0.002 for salinity and temperature, ±0.015 % of
full-scale range for pressure and ±1 % for O2, according to
WOCE standards (Joyce and Corry, 1994). Samples from the Niskin bottles were analysed
for DIC by coulometry and TA by open cell
potentiometric titration (Dickson et al., 2007). Certified reference material
provided by A. Dickson, Scripps Institution of Oceanography, were used as
reference standards for DIC and TA. The precisions of DIC and TA measurements
improved slightly on more recent sections, and for all sections were better
than ±2 µmol kg-1, for both variables, based on
analysis of duplicate samples and certified reference material. Samples for
dissolved O2 were measured using modified Winkler titrations (Hood et al
2010), with an estimated accuracy and precision of ±0.3 % for the
sections. For the 1995 cruise, sensor-based O2 was used instead of
sampled O2 because of the poor quality of many of the Winkler
measurements.
Other variables of the dissolved CO2 system were calculated from the DIC
and TA measurements. We calculate pHT and ΩAr
(from measured DIC and TA) using the CO2sys program from Lewis and Wallace (1998)
adapted to MATLAB by van Heuven et al. (2011). We use the constants
for the carbonic acid from Mehrbach et al. (1973) refit by Dickson and
Millero (1987), the CO2 solubility equation from Weiss (1974), dissociation
constants for sulfate from Dickson (1990) and borate constant from Uppstrom (1974). Aragonite saturation states (ΩAr) are calculated
because it is a less stable form of CaCO3 than calcite and is the
predominant biogenic form of CaCO3 precipitated by calcifying organisms.
Data presented here were interpolated to a regular grid, and the water mass
layers were defined as layers between neutral surfaces (γn)
determined using potential temperature and salinity and published literature
values (Table 2, Sect. 2). The first 50 m of the water column were
eliminated in order to reduce the short-timescale variability in surface
properties. The use of γn to identify the water mass layers
reduces the variability due to isopycnal heave caused for example by eddies
and internal waves (McDougall et al., 1987; Bindoff and McDougall, 1994;
Jackett and McDougall, 1997). The upper ocean layers south of the SAF were
divided into the AASW layer to the north of the PF, and the AASWupw
layer that is composed of surface waters south of the PF (Fig. 1). ADLBW and
RSBW were included in the AABW layer (Table 2). The ACC fronts along the SR03
section (Table 3) were defined as a function of hydrographic variables
following Sokolov and Rintoul (2002).
Definition of the water mass layers between neutral
surfaces (γn) and references used to
accordingly decide the limits of the layers. We also consider limits in
salinity, depth and position of the ACC fronts (see Table 3) to
differentiate better some of the layers.
LayerDefinitionComplete nameReferenceSTCWΥn < 26.85 & S ≥ 34.3Subtropical Central WaterRintoul (1998)AASWΥn < 27.7 & S < 34.3Antarctic Surface WaterRintoul (1998);Williams et al. (2015)AASWupwΥn≥ 27.7 & depth ≤ 300 mAntarctic Surface Water–SAMW26.85 ≤Υn < 27.2 & S ≥ 34.3Sub-Antarctic Mode WaterRintoul (1998);Rintoul and Bullister (1999)AAIW27.2 ≤Υn < 27.7 & S ≥ 34.3Antarctic Intermediate WaterRintoul and Bullister (1999)UCDW27.7 ≤Υn < 28.18 & depth > 300 mUpper Circumpolar Deep waterWilliams et al. (2015)LCDW28.18 ≤Υn < 28.25Lower Circumpolar Deep WaterLacarra et al. (2011);Williams et al. (2015)AABWΥn≥ 28.25Antarctic Bottom WaterRintould and Bullister (1999);Williams et al. (2015)Estimates of anthropogenic carbon (CANT)
CANT was estimated using a back-calculation method (Chen and
Millero, 1979; Gruber, 1998) combined with an optimum multi-parameter (OMP)
analysis (Tomczak, 1981), described by Pardo et al. (2014). This technique
has the advantage of considering water mass mixing and the temporal
variability of the air–sea CO2 disequilibrium. The accuracy of the
method is ±6 µmol kg-1 (Pardo et al., 2014).
Back-calculation methods assume the ocean is at steady state for dynamical
and biological processes and estimate CANT
(CANT_BC) as an excess of DIC in the ocean resulting from
the increase in atmospheric CO2 due to anthropogenic emissions as follows:
CANT_BC=DIC-DICBIO-DICπ,
where DICBIO is the biological contribution to DIC, and DICπ
is the preformed concentration of DIC in pre-industrial times.
The term DICBIO is calculated (Chen et al., 1982; Ikegami and
Kanamori, 1983) by the following:
DICBIO=AOURC+12TA-TA0+AOU⋅1RN+1RP,
where RC, RN and RP are the
stoichiometric ratios of carbon, nitrate and phosphate respectively, referred
to O2 consumption by respiration–remineralisation processes that are
considered constant (1.45, 9 and 125, respectively, Broecker, 1974; Anderson
and Sarmiento, 1994; Martiny et al., 2013).
The term AOURC represents the remineralisation
of organic matter, with the apparent oxygen utilisation (AOU) defined as
AOU = OSAT – O2, i.e. the difference between the
saturation of oxygen (OSAT) at the potential temperature (θ) and salinity of the measured O2, and the measured O2. The term
12TA-TA0+AOU⋅1RN+1RP represents the
dissolution or precipitation of CaCO3, with TA0, the
preformed alkalinity, obtained in the water formation sites using regional
parameterisations as functions of salinity, θ and phosphate (Pardo et
al., 2011; Vázquez-Rodríguez et al., 2012; Supplement Sect. S3,
Table S3).
The preformed pre-industrial term, DICπ , is the total
concentration of carbon dioxide in seawater saturated with respect to the
pre-industrial pCO2 (DICSATπ) and corrected for an air–sea
CO2 disequilibrium (CDISπ) term:
DICπ=DICSATπ+CDISπ.
CDISπ is time dependent:
CDISπ=CDIS-δCDIS,
where CDIS is the current disequilibrium between the ocean and the
atmosphere pCO2, and δCDIS is the change in the disequilibrium
from pre-industrial to current times. CDIS was obtained by similar
parameterisations to those used for TA0, combined with monthly mean
values of atmospheric CO2 values from the NOAA network (Dlugokencky et
al., 2016) (see Sect. S3 and Table S3). The δCDIS
values were obtained from results of the 1/10∘ resolution carbon
model OFAM3-WOMBAT (Sect. S1).
Interior values of preformed variables (TA0 and DICSATπ)
and CDISπ were obtained using an OMP
analysis to mix endmembers as described in Sects. S2 and S3,
using the endmembers in Table S1 in the Supplement. The OMP analysis is based on the
assumption that a property measured in a certain point is the result of
linear mixing between endmembers, known as source water types (SWTs). A
system of equations is created for each measurement point and is solved to
obtain the fractions of the different SWTs (Sect. S2). The
application of the OMP analysis requires good regional hydrodynamic knowledge
as the results are strongly dependent on the definition of the SWTs (Tomczak,
1981). We used 11 SWTs to characterise the biogeochemical properties of the
waters in the SR03 section and the SWT properties were assumed to be constant
with time (Table S1, Fig. S1).
Small negative values of CANT_BC can occur due to an
overestimation of the CDISπ term (Eq. 4), which acts as measure of
the age of the water mass and has high values in old deep layers (see Table S1). These negative values of CANT_BC
were found at some points in deep waters of the section (mainly UCDW and LCDW
layers) and were small (between 0 and -2 µmol kg-1, i.e.
less than the accuracy of the methodology) and changed to zero for our
analysis.
Changes in the carbon system
Changes in carbon system parameters ∂DIC∂t,∂TA∂t,∂CANT_BC∂t,∂pH∂tand∂ΩAr∂t were estimated using linear
regressions with time for the period 1995–2011 in each one of the water mass
layers defined by their γn condition (Table 2). The trends
were estimated using all the points in each water mass layer, and only those
linear trends with p < 0.05 and r > = 0.2 were considered
statistically significant for discussion. We show the value of the root mean
square error (RMSE or square root of the variance of the residuals), which
can be interpreted in large part as unexplained variance caused by short-timescale processes and the different seasonal timings of the cruises. RMSE has
the same units as the response variable.
With respect to the total change in DIC ∂DIC∂t, our
goal is to disentangle the effects that solubility, circulation, biology and
CANT uptake have on the variability of DIC. The total change in
DIC ∂DIC∂t in a water mass is due to changes in
the atmosphere–ocean interchange, biological processes and circulation
processes. In order to account for the change in DIC due the atmosphere–ocean
interchange and biological processes, we compare ∂DIC∂t to ∂CANT_BC∂t
and ∂DICBIO∂t, Eq. (2). The change in
DIC not explained by ∂CANT_BC∂t
or ∂DICBIO∂t will then be due to circulation
processes.
In order to compare ∂DIC∂t to ∂CANT_BC∂t we need to consider how the change
in ∂DICBIO∂t and DICπ∂DIC∂tπ (terms of Eq. 1,
Sect. 3.2) affect the changes in CANT_BC.
The term ∂DIC∂tπ can be expressed as follows:
∂DICπ∂t=∂DICSATπ∂t-∂CDIS∂tπ.
The terms ∂CDIS∂tπ and ∂DICSATπ∂t reflect changes in the properties of the
water masses over time, primarily temperature and salinity change due to
mixing and heating or cooling. CDISπ and DICSATπ are
defined at the ocean surface (in each of the SWTs: Table S1) and are
calculated at each point in the ocean interior using the OMP analysis
(Sects. S2 and S3). Because a change in temperature and/or
salinity in the water is solved by the OMP analysis as a change in the SWTs
fractions, this also produces varying CDISπ and DICSATπ. No significant trends were obtained for CDISπin any of the
layers. CANT_BC (Eq. 1) is not affected by the changes in
solubility occurring from one voyage to another thusneitheris∂CANT_BC∂t, since any change in
temperature or salinity is cancelled out by the subtraction of DICπ
with respect to DIC in Eq. (1) (Sect. 3.2) and by O2 respect
to OSAT in the DICBIO term (Eq. 2). However,
∂DIC∂t based on the measured DIC in the sections will be affected
by changes in the solubility over time, and this difference needs to be
accounted for when comparing ∂DIC∂t with
∂CANT_BC∂t to obtain a better
approximation for the change in CANT: ∂CANT∂t=∂CANT_BC∂t+∂DIC∂tπ.
Location of the ACC fronts south of Tasmania for each of
the cruises following the definitions of Sokolov and Rintoul (2002) for
hydrographic data. * The range in the location of the SAF for the
2011 cruise could be related to a diversification of the PF or SAF into
different jets (Sokolov and Rintoul, 2009) but also to crossing the meander
of the front twice.
Trends in the water mass layers for the period 1995–2011 of
dissolved inorganic carbon (∂DIC /∂t),
anthropogenic carbon (∂CANT/∂t), silicate
(∂SiO4/∂t)
total pH (∂pHT/∂t), aragonite saturation (∂ΩAr/∂t) and percentage of change in the
aragonite saturation (% ∂ΩAr/∂t). RMSE = root
mean square error. * Trends in CANT for the AASW and
AASWupw layers considering an approximate value for
the increase in DIC due to the advection of old deep waters to the section
(see Sect. 4.4.1 in the text). ** The value of ∂CANT∂t in the AABW layer
is considered negligible because it falls below the accuracy of the
back-calculation method.
The term ∂DICBIO∂t can be influenced by
changes with time of alkalinity due to changes in the rate of carbonate
precipitation–dissolution and of AOU due to changes in the rate of
remineralisation and in circulation. In the present study only surface waters
of the SR03 section present changes DICBIO between 1995 and 2011.
Numerous studies have reported a strong influence of biological communities
in the seasonal cycle of dissolved O2 in surface waters (Bender et al.,
1996; Moore and Abbott, 2000; Sambrotto and Mace, 2000; Trull et al., 2001a).
Inter-annual variability in O2 in upper layers of the Southern Ocean have
also been related to changes in the entrainment of deeper waters into the
mixed layer due to the mixed layer depth variability (Matear et al., 2000;
Verdy et al., 2007; Sabine et al., 2008; Sallée et al., 2012). Although
some studies found long-term decreases in O2 due to circulation in deep
waters of the Weddell Sea (van Heuven et al., 2014) and for the first 1000 m
of the global ocean (Helm et al., 2011), significant long-term trends in
O2 due to circulation and remineralisation processes have not yet been
reported for surface waters of the Southern Ocean. Thus, the term
∂DICBIO∂t may also contribute to variation in
∂CANT_BC∂t, since part of the
changes in AOU with time reflect changes in circulation that we cannot
separate from those in remineralisation. We consider the best approximation
for the change in CANT as a range depending on the possible
effect of biology and circulation processes on ∂DICBIO∂t. If the value of ∂DICBIO∂t is due to the variability in the remineralisation
rates and the change in solubility is considered, the estimate ∂CANT_BC∂t will be the lower
limit of the range (lower limit of ∂CANT∂t=∂CANT_BC∂t+∂DIC∂tπ). For
the upper limit of the range, we consider that the value of ∂DICBIO∂t is due to changes in circulation and the upper
limit of the range is obtained by ∂CANT∂t=∂CANT_BC∂t+∂DIC∂tπ+∂DICBIO∂t. We assume that the changes in DICBIO due to
circulation do not affect the amount of DIC in the layer. This assumption is
one of the caveats of the methodology, since we cannot know how much of the
change in DIC is associated with changes in circulation, i.e. how much of the
change in DIC is a change in non-anthropogenic DIC. We will discuss this more
in Sect. 6.2.
Results
The different trends in biogeochemical properties are summarised in Table 4.
The biogeochemical changes between 1995 and 2011 are presented for each of
the water mass layers and the effect of changes in solubility, biological
processes and circulation in the estimates of ∂CANT∂t and ∂DIC∂t are
considered along with changes in the aragonite saturation depth and
CANT storage.
Trends of anthropogenic carbon from estimates of the
back-calculation method and of the terms DICBIO and DICπ (see Sect. 3.2 of the text). RMSE = root mean square error. Bottom:
schematic of the estimation of the upper and lower limits of ∂CANT∂t (see Sect. 3.3).
In the STCW layer, DIC increased between 1995 and 2011 (Fig. 2a, b) at a rate
of 0.86 ± 0.07 µmol kg-1 yr-1∂DIC∂t,Table4,
leading to a decrease in pHT of
-0.0027 ± 0.0001 yr-1∂pHT∂t,Table4,Fig.2e,f. The trend
in pHT is similar to the one found by Lauvset et al. (2015)
between 1991 and 2011 for the IO-STPS (Indian Ocean subtropical
permanently stratified) biome (-0.0027 ± 0.0005 yr-1). We found
a decrease in DICπ∂DIC∂tπ=-0.34±0.06µmolkg-1yr-1,Table5
in the STCW layer due to a negative trend of DICSATπ
resulting from a decrease in solubility that resulted from a temperature
increase (calculated from the section data) in the STCW layer of
0.0335 ± 0.0130 ∘C yr-1 (not shown). The increase in
temperature agrees with the warming trend observed south of Tasmania of 0.2
to 0.3 ∘C decade-1, obtained from satellite data (Armour and
Bitz, 2015) and from combined data and models (0.5 ∘C every
30 years, Aoki et al., 2015). For θ= 16 ∘C and S = 35.1
(definition of SWTSTW16 in the OMP analysis, Table S1), a change
in temperature of 0.03 ∘C yr-1 would lead to a decrease in
DICSATπ of -0.27 µmol kg-1 yr-1,
which is similar to the value obtained for ∂DIC∂tπin the STCW layer (Table 5). The difference
between these trends is related to the mixing of the different SWT fractions
within the STCW layer, established from the OMP analysis (see Sect. 3 and
Sect. S2). When the solubility change is incorporated into ∂CANT_BC∂t, i.e. ∂CANT_BC∂t+∂DIC∂tπ (Table 5, Fig. 2c, d), we obtain a value of
0.71 ± 0.08 µmol kg-1 yr-1.
Distribution of DIC (a, b),
CANT_BC(c, d) and
pHT(e, f) in the SR03 section south of Tasmania
for the years 1995 (a, c, e) and 2011 (b, d, f). Red lines in plots (a) and (b),
pink lines in plots (c) and (d), and green lines in plots (e) and
(f) indicate the neutral surfaces that define the different water masses
(γn). The position of the fronts (coloured
triangles and grey dotted lines) in each of the cruises is also shown.
There is an increase in DICBIO in the STCW layer ∂DICBIO∂t,Table5 that also affects the estimates of
∂CANT_BC∂t. The increase in
DICBIO is due to an increase in AOU (no changes were found in TA, Eq. 2 in
Sect. 3.2) due to a decrease in O2 in the layer. We
cannot separate the effects of circulation and biology on the AOU change, and
∂CANT∂t in Table 4 should be considered
a range. If the changes in AOU are only due to the variability in the
remineralisation rates, the calculated lower limit of ∂CANT∂t is 0.71 ± 0.08 µmol kg-1 yr-1
(Table 4, ∂CANT∂t=∂CANT_BC∂t+∂DIC∂tπ in Table 5). If the changes in AOU
are due to changes in circulation, the upper limit value of 0.93 ± 0.11 µmol kg-1 yr-1∂CANT∂t=∂CANT_BC∂t+∂DIC∂tπ+∂DICBIO∂t will explain the increase in DIC
in the STCW layer ∂DIC∂t≈∂CANT∂t,Table4. The increase in CANT
found in this layer is comparable to the range of increases (0.8–1.3 µmol kg-1 yr-1)
found by Carter et al. (2017) in
the Pacific Ocean (P16 WOCE, CLIVAR and GOSHIP lines) for the past 2 decades (1990s–2000s and 2000s–2010s).
Changes in the AASW layer are summarised in Table 4. DIC increased at a
similar rate of 0.85 ± 0.14 µmol kg-1 yr-1 to the
STCW layer and the trend is similar to the values found by Williams et
al. (2015) for the AASW layer in the Pacific sector of the Southern Ocean
(12–18 µmol kg-1 for the period 1992–2011 and
3–5 µmol kg-1 for the period 2005–2011). The increase in DIC
in the AASW layer results in a pHT decrease of
-0.0035 ± 0.0002 yr-1, close to Williams et al. (2015) estimates
for surface waters (∼-0.0023 ± 0.0009 yr-1) and Lauvset et
al. (2015) estimates of -0.0021 ± 0.0002 yr-1 for the Southern
Ocean seasonally stratified, SO-SPSS, biome. The AASW layer for our sections
warmed at a similar rate (0.0369 ± 0.0109 ∘C yr-1) to the
STCW layer, reducing the solubility and influencing
∂DIC∂tπ (Table 5) due to changes in DICSATπ.
The DICBIO also increased with time (0.50 ± 0.16 µmol kg-1 yr-1, Table 5) due to an
increase in AOU. Following the same reasoning as for the STCW layer and
considering the trend in CANT of
0.70 ± 0.06 µmol kg-1 yr-1 obtained by the
back-calculation method ∂CANT_BC∂t;Table5, the best estimation of ∂CANT∂t in the AASW layer is a range of
0.35 ± 0.14 to 0.85 ± 0.22 µmol kg-1 yr-1
(Table 4). Our values are within the range of values found by Williams et
al. (2015) for the AASW in the Pacific sector between 2005 and 2011, and the
upper limit is similar to a CANT increase of 0.73–0.86 µmol kg-1 yr-1
for waters South of Tasmania for the period 1968–1996 found by McNeil et al. (2001).
DIC in the AASWupw layer increased at a rate of
0.61 ± 0.10 µmol kg-1 yr-1, and the
pHT decreased by -0.0015 ± 0.0004 yr-1 (Table 4). We
were not able to detect a statistically significant trend in DICπ
(i.e. solubility) or CANT from the estimates of the
back-calculation method ∂CANT_BC∂tTable5. However, we found an increase in DICBIO of
0.42 ± 0.28 µmol kg-1 yr-1 (Table 5) that is due
to an increase in AOU. Considering the different drivers of the AOU increase
(e.g. biology, circulation), the optimal estimation of ∂CANT∂t for this layer is a value between 0 and
0.42 ± 0.28 µmol kg-1 yr-1 (Table 4).
Mode waters and intermediate layers (SAMW and AAIW)
The increase in DIC in the SAMW layer
(1.10 ± 0.14 µmol kg-1 yr-1) for the period
1995–2011 is higher than that of upper ocean layers, and pHT
decreases over the same period by -0.0031 ± 0.0003 yr-1 (Table 4).
The DIC increase is explained almost entirely by ∂CANT_BC∂t of
0.92 ± 0.09 µmol kg-1 yr-1. No significant trend
was found in DICBIOor DICπi.e,∂CANT∂t=∂CANT_BC∂t. In the AAIW layer the DIC trend of
0.40 ± 0.15 µmol kg-1 yr-1 results in a
pHT decrease of -0.0017 ± 0.0002 yr-1 and is also
explained by the increase in CANT
(0.42 ± 0.06 µmol kg-1 yr-1, ∂CANT∂t=∂CANT_BC∂t, Tables 4 and 5). As with SAMW, no changes in solubility
∂DIC∂tπ or biology/circulation processes
∂DICBIO∂t were detected in the AAIW layer.
The values found in the SAMW and AAIW layers are very similar to the mean
decadal changes found by Murata et al. (2007) between the 1990s and the 2000s
in the subtropical Pacific Ocean (∼ 1 µmol kg-1 yr-1 for the SAMW layer and
0.4 µmol kg-1 yr-1 for the AAIW). Waters et al. (2011)
used data from the P18 line along ∼ 110∘ W and estimated an
increase in CANT of 0.89 ± 0.4 µmol kg-1 yr-1
for the SAMW and 0.64 ± 0.2 µmol kg-1 yr-1 in
AAIW for the period 1994–2008, which are also comparable to
our results.
Deep–bottom layers (UCDW, LCDW and AABW)
The UCDW layer shows an increase in DIC of
0.29 ± 0.02 µmol kg-1 yr-1 between 1995 and 2011
and a change in pHT of -0.0013 ± 0.0001 yr-1, which
is similar to the change in DIC (0.20 ± 0.02 µmol kg-1 yr-1) and pHT
(-0.0012 ± 0.0002 yr-1) for LCDW. No statistically significant
changes in time were detected for CANT_BC or in the
DICBIO and DICπ terms for any of these two layers.
The AABW layer also shows an increase in DIC
(0.24 ± 0.02 µmol kg-1 yr-1) during the period
1995–2011 with an associated decrease in pHT of
-0.0013 ± 0.0002 yr-1 (Table 4). The increase in
CANT∂CANT_BC∂t of
0.07 ± 0.01 µmol kg-1 yr-1 is low and this trend
indicates an increase in CANT of
∼ 1 µmol kg-1, which is less than the accuracy of the
back-calculation method (±6 µmol kg-1).
Changes in the aragonite saturation (∂ΩAr∂t) and
CANT storage
There are statistically significant decreases in ΩAr in the
STCW, AASW and SAMW layers (∼-0.010 ± 0.001 yr-1, Table 4)
similar to the trends observed at open-ocean time series sites in recent
decades (Bates et al., 2014). The decrease in ΩAr found for
the AASW layer (-0.61 ± 0.19 % yr-1, Table 4) is also similar
to the values obtained by Williams et al. (2015) for the Pacific sector of
the Southern Ocean (-0.47 ± 0.10 % yr-1 for the period 1992–2011
and -0.50 ± 0.20 % yr-1 for the period 2005–2011).
Accompanying the decrease in ΩAr with time along SR03 is
the shoaling of the aragonite saturation depth (ASD, ΩAr= 1, Fig. 3) at a mean rate of -13 ± 3 m yr-1. The
shoaling of the ASD is not uniform over the section. North of the PF, the ASD
shoals at a rate of -6 ± 4 m yr-1 while the rate is 3.5 times
greater south of the PF (-21 ± 4 m yr-1). North of the PF the
shoaling mostly affects the AAIW layer (Fig. 3a). South of the PF, from
∼ 62∘ S, the movement of the ASD follows the upwelling path of
the UCDW layer (Fig. 3) with a shoaling of ∼ 340 m over the 16-year
period.
Approximated rates of CANT storage based on the trends from Table 4.
Aragonite saturation depth (ASD, ΩAr= 1)
for the 1995 and 2011 cruises. (a) Distribution of ASD (blue
dotted and solid lines) and γn (green and yellow
lines) in latitude in the SR03 section. (b) Section zoomed for the first
1500 m of the water column. The position of the fronts (triangles) in 1995
(yellow) and 2011 (green) is also shown. The blue palette in the background
indicates the distribution of ΩAr for the
2011 cruise.
The storage rate of CANT (Table 6) for the surface and
intermediate water mass layers is obtained from ∂CANT∂t
(Table 4), with the most storage in SAMW and
AAIW due to both their greater thickness and ∂CANT∂t values. The rate of increase of the
CANT storage in the whole longitude band of the SR03 section is
0.30 ± 0.24 mol m-2 yr-1, calculated by computing the mean
of the storage rates of the layers weighted by the mean volume occupied by
each of the layers for the period 1995–2011 (Table 6)
Discussion
Our results are indicative of a scenario of increased transport of deep
waters into the section and enhanced upwelling at high latitudes for the
period between 1995 and 2011 linked to strong westerly winds. Several studies
have reported a trend in the Southern Annular Mode (SAM) toward its positive
phase from the 1960s until the 2000s (Thompson and Solomon, 2002; Marshall,
2002, 2003; Lenton and Matear, 2007; Sallée et al., 2008). According to
these studies, the positive phase of the SAM is correlated with an
intensification and southward movement of the subpolar westerly winds that
ultimately lead to the enhancement of northward Ekman transport, meridional
overturning and upwelling south of the ACC. Also, surface warming and more
intense and frequent pulses in the extension of the EAC on long timescales
have been related to a poleward movement of the westerly winds (Rintoul and
Sokolov, 2001; Ridgway, 2007; Hill et al., 2011). From the 2000s on, the SAM
index no longer presents a positive trend but, despite exhibiting
considerable inter-annual variability (Fig. S2),
remains in its positive phase, favouring strong winds over the
region.
In the northern part of the SR03 section, the area occupied by the STCW has
high variability due to the encounter between the EAC and the ZC in the north
of the section (Ridgway et al., 2007; Herraiz-Borreguero and Rintoul, 2011;
Sloyan et al., 2016). The warming of the STCW layer found in this study
(0.0335 ± 0.0130 ∘C yr-1) could be linked to variability
in the extension of subtropical waters but it could also be related to
atmospheric warming. Aoki et al. (2015) related the 30-year warming found
north of the SAF in the South Pacific and Indian oceans to the
intensification of the subtropical gyres, which promote the arrival of warmer
waters. In the AASW layer that extends approximately between the SAF and the
PF we found a similar warming (0.0369 ± 0.0109 ∘C yr-1)
to that of the STCW. This could indicate that the increase in temperature
found in the upper layers of the section could be most likely due to ocean
heat uptake and atmospheric warming.
Due to the surface warming, the increase in DIC found in the STCW layer
(Table 4) is lower than expected from the increase in atmospheric CO2
(∼ 1 ± 0.12 µmol kg-1 yr-1). Nevertheless,
at least 83 % of the increase in DIC in the STCW layer is explained by the
increase in CANT (Table 4). As for the AASW layer, our results
indicate that temperature does affect the estimate of
∂DIC∂t, but the effect of the increase in DICBIO (due to an
increase in AOU) overweigh that of solubility (Table 5).
The seasonal to inter-annual variability of the AASW layer is also influenced
by the variability of the positions of the SAF and PF (Fig. 1, Table 3),
which
is highly conditioned by the flow of the ACC over the South-East Indian Ridge
(Fig. 1). A close relationship between phytoplankton blooms and regions where
the ACC fronts interact with large topographic features has been noticed
(Moore et al., 1999; Moore and Abbott, 2000). A variability in the
remineralisation rates due to phytoplankton bloom variability could explain
the changes in DICBIOobserved in the AASW layer. Nevertheless, no
changes in nutrients (nitrates or phosphates) are measurable in this layer
that could indicate intense biological activity.
Comparison between ∂CANT∂t from the present
study and the values of ΔCANT from
the two-regression and eMLR methods. RMSE = root mean square
error. ** The value of ∂CANT∂t in the AABW layer
is considered negligible because it falls below the accuracy of the
back-calculation method.
Furthermore, the AASW layer is also affected by the upwelling of deep waters
south of the PF, and an intensification of the upwelling could increase the
content of low-O2 DIC-rich waters in the AASW layer, leading to an
increase in AOU. The increase in DICBIO found in the AASWupw
layer (Table 4), south of the PF, is similar to the increase obtained for the
AASW layer, which indicates the likelihood that the upwelling of deep waters
results in the increase in AOU. The increase in DICBIO in the
AASWupw layer coincides with an increase in salinity of 0.0029 ± 0.0001 yr-1
(not shown), which is consistent with increased transport of
saltier waters from the deep ocean to subsurface layers. Besides, we also
found an increase in dissolved silicate of
0.36 ± 0.06 µmol kg-1 yr-1∂SiO4∂tTable4 that could be related to the upwelling
enhancement as well (Tréguer, 2014).
The influence of the upwelling on the DIC budgets (as non-anthropogenic DIC)
is clearer in the AASWupw layer than in the AASW layer. For the AASW
layer, the lower limit of ∂CANT∂t (i.e.
the change in DICBIO is assumed to be due to biological processes,
Table 4) indicates that at least 41 % of the increase in DIC in the layer
is explained by the increase in CANT while the lower limit of
∂CANT∂t is zero for AASWupw (Table 4),
meaning that the effect of the upwelling over AASW is lower than over
AASWupw. Matear and Lenton (2008), using carbon models, concluded that
the uptake of CO2 by the waters north of the PF is more influenced by
the wind variability than by other processes such as the upwelling. An
intensification of the winds (due to a positive phase in SAM) could
contribute to the increase in CANT found in the AASW layer.
Considering the upper limits of ∂CANT∂t in both layers (i.e. the change in DICBIO is assumed to be due to
circulation processes), the increase in CANT in the AASWupw
layer represents no more than 69 % of the increase in DIC (upper limit of
∂CANT∂t, Table 4), while the upper
limit of ∂CANT∂t for the AASW equals
the increase in DIC. Thus, in the AASWupw layer, at least ∼ 30 % of the
increase in DIC (∼ 0.18 µmol kg-1 yr-1) is still
not explained and is most probably related to the upwelling of DIC-rich
waters. The increase in non-anthropogenic DIC could be even higher, since we
assume that the change in DICBIO due to circulation does not affect
DIC (see Sect. 3.3).
In terms of the change in oxygen, Helm et al. (2011) found an average
decrease in the concentration of O2 between 100 and 1000 m from 1970 to
1992 of ∼-0.23 µmol L-1 for the Southern Ocean
(27 % of the estimated global average change,
-0.93 ± 0.23 µmol L-1) . Considering the volume of
the first 1000 m of the water column of the Southern Ocean to be
19 400 × 10-9 l (obtained using ETOPO1,
Amante and Eakins, 2009) and the volume of the first 1000 m of the SR03
section to be 2700 × 10-9 l, the decrease in O2 found by
Helm et al. (2011), if constant in time, would correspond to a decrease of
∼-1.7 µmol L-1 yr-1. We only found changes in
oxygen within the surface water mass layers (STCW, AASW and
AASWupw) that approximately fill the first 300 m of the water
column of the SR03. Then, the decrease of
∼-1.7 µmol L-1 would correspond to an average change
in O2 of ∼-0.32 µmol kg-1 yr-1 for
surface waters of the SR03. This means that values of
∼ 0.20 µmol kg-1 yr-1 due to circulation
processes can be expected in ∂DICBIO∂t for surface waters, which is
comparable to the average of our findings (Table 5),
0.32 ± 0.24 µmol kg-1 yr-1, and could indicate
that the change in O2 is related to circulation processes.
The variability of the SAMW and AAIW layers south of Tasmania has been
related to variability in the northward Ekman transport that drives the
northward movement of AASW (Rintoul and England, 2002; Sallée et al.,
2006, 2012). A scenario of intensification of the upwelling near the
Antarctic Divergence would lead to an increase in the northward Ekman
transport, conditioning the properties of these water mass layers,
particularly for SAMW, which is mostly formed north of the SAF. There is a
significant freshening of the SAMW layer (-0.0026 ± 0.0001 psu yr-1,
not shown) between 1995 and 2011 that could be related to higher
inputs of AASW into the SAMW layer and is consistent with the increase in Ekman
transport. Besides, an intensification of the winds due to the positive trend
of the SAM favours the ventilation and thus the increase in CANT
uptake by both water mass layers (Matear and Lenton, 2008). Our results
indicate that the change in DIC in the SAMW and AAIW layers is driven mostly
by the uptake of atmospheric CO2∂DIC∂t≈∂CANT∂t,Table4. The
increase in CANT in the SAMW layer is higher than that found for
the upper ocean layers and closer to the expected values from the increase in
atmospheric CO2 (∼ 1 µmol kg-1 yr-1). The
smaller increase in CANT in the AAIW layer compared to the SAMW
layer (Table 4) agrees with lower ventilation of the AAIW layer south of
Tasmania (see Sect. 2) due to the fact that this layer carries recently
ventilated waters mixed with older waters ventilated far out the SR03
section. The lack of measurable long-term changes in DICBIO and
DICπ in both AAIW and SAMW layers indicate that circulation and
biological processes do not have a large effect on ∂DIC∂t.
Deep to bottom layers of the section show significant trends for DIC that are
not explained by the increase in CANT. These trends are most
likely due to the advection of old and DIC-rich waters. Concretely for deep
waters (UCDW and LCDW), the trends could result from an intensification of
upwelling at high latitudes being offset by enhanced transport of old and
CO2-rich waters to replace the upwelled waters, since the increase in
DIC follows the upwelling path of the UCDW and LCDW layers (Fig. 4). We
separated the UCDW layer into two latitudinal sectors: north and south of the
SAF (Fig. 4). The increase in DIC in the UCDW layer north of the SAF is
0.44 ± 0.04 µmol kg-1 yr-1 while that south of the SAF is
smaller at 0.26 ± 0.04 µmol kg-1 yr-1 (not
shown), consistent with a greater supply of waters from the north at depth. A
decrease in O2 in the UCDW to the north of the SAF occurs mostly in the
upper to middle parts of the UCDW layer (Fig. 4), and this is not observed
south of the SAF. The decrease in O2 is also in agreement with the
arrival of waters from Indian–Pacific origin since these waters provide the
characteristic oxygen minimum zone that defines UCDW (Callahan, 1972; Talley
2013). Another feature that agrees with the hypothesis of upwelling
intensification is the shoaling of the ASD following the path of upwelling of
the UCDW layer. This feature was also described by Bostock et al. (2013) in
an oceanic climatology of ΩAr and could be due to the
naturally lower buffer capacity of the UCDW layer (low value of TA / DIC ≈ 1.043) with
respect to upper layers (TA / DIC ≈ 1.06 in the
AASW layer). However, the greatest shoaling of the ASD in the UCDW layers
compared to the AAIW layer (Table 6) is consistent with the upwelling of
UCDW, as both water masses have similar TA / DIC ratios (TA / DIC ≈ 1.043
for the UCDW layer and TA / DIC ≈ 1.042 for the AAIW layer).
Furthermore, the increase in SiO4 found in deep–bottom layers (Table 4)
could also indicate the arrival of old waters to the section that are
progressively enriched in SiO4 (e.g. Callahan, 1972).
Distribution of (a, b) dissolved inorganic carbon (DIC)
and (c, d) dissolved oxygen (O2) in the 1995 and
2011 cruises for deep–bottom layers of the section. The location of the
Sub-Antarctic Front (SAF) is also shown as well as the neutral surfaces
(γn, white lines) limiting the UCDW, LCDW and
AABW layers.
Statistically significant decreases in pHT with time were
observed in all water mass layers (Table 4), with the greatest change in
surface water masses coinciding with the greatest DIC changes. The decrease in pHT in the STCW and SAMW layers is related to the increase in
the uptake of CANT, while for the AASW and AASWupw layers the
pHT change appears to be linked to the upwelling of DIC-rich
waters at high latitudes. At deep layers, the tongue of water of
pHT= 7.9 off the shelf is reduced in 2011 compared to 1995
(Fig. 2e, f), which is consistent with the advection of DIC-rich waters in the
section due to the enhanced upwelling. The different rates of pHT
change in the water masses is in part related to the buffering capacity of
the waters. The AASW layer has lower temperature than the STCW layer (∼ 2.3 ∘C
for the AASW layer compared to ∼ 11.0 ∘C for the
STCW layer, mean values for the period 1995–2011) and lower buffer capacity
than the STCW (TA / DIC ≈ 1.058 for the AASW layer versus TA / DIC ≈ 1.095
for STCW). For similar increases in DIC (Table 4) the
decrease in pHT in the AASW is expected to be higher than in the
STCW layer.
In terms of carbon, previous studies concluded that the intensification of
the upwelling (as a consequence of the SAM variability) caused a reduction in
the uptake of CO2 by the Southern Ocean between the 1980s and 2000s due to
the outgassing of CO2 near the Antarctic Divergence (Le Quére et al.,
2007; Lovenduski et al., 2008). Landschützer et al. (2015) showed that
the efficiency of the Southern Ocean CO2 sink declined through the
1990s, and the trend reversed from about 2002, although the reversal in the
sink efficiency was not zonally uniform. The results from Landschützer et
al. (2015) are consistent with a carbon sink influenced by the upwelling of
DIC-rich waters at high latitudes, and superimposed on this is the near-surface response to atmospheric forcing that modifies the sink efficiency and
could mask longer-term trends in the upwelling of DIC-rich waters at high
latitudes. A comparison of our results with those of Landschützer et
al. (2015) is problematic as their data are restricted to surface waters and
our analysis is on long-term trends in water mass properties below 50 m depth.
Both data sets do show continued uptake of CO2 throughout the period of
study and indicate the importance of the circulation in influencing the
regional carbon sink, which has also been established by recent model results
(DeVries et al., 2007).
Our results also agree with the conclusions from different model simulations
done by Matear and Lenton (2008), who established that intense wind regimes
(associated with a positive phase in the SAM) favour the uptake of
CANT and ventilation of the SAMW and AAIW layers. These authors
highlighted the complex response of the uptake of CO2 by the Southern
Ocean due to the diverse forcing acting on upper layers, which can be also
seen in our results (e.g. differences in the biogeochemical changes in the
AASW and AASWupw layers). Matear and Lenton (2008) also noticed the
complex relationship between the upwelling and subduction areas of the
Southern Ocean, with the same drivers acting in opposite directions for the
changes in non-anthropogenic DIC with respect to the changes in
CANT uptake.
Sensitivity of the results to underlying assumptions
This section considers the sensitivity of assumptions used to calculate
temporal changes in CANT, including errors associated with the
assumption of steady state in the oceanic circulation and remineralisation
processes, and the sensitivity to stoichiometric ratios for the biological
processes.
Comparison of CANT changes using other methods
We compared the changes in CANT obtained in our study with the
results from two regression-based methodologies (Table 7): the extended
multiple linear regression (eMLR) method (Friis et al., 2005) and the
two-regression method (Thacker, 2012). These methods use repeated hydrodynamic
sections to quantify the temporal change in CANT.
The eMLR method (Friis et al., 2005) estimates the change in CANT
between two repeats of a hydrodynamic section by establishing MLRs for each
section and relating the observed DIC for each observation to a set of other
measured oceanic variables:
DIC(t)=a0(t)+a1(t)P1(t)+…+an(t)Pn(t),
where ax(t) are the coefficients of the fit between DIC and the n
observed variables (P1, …Pn) chosen for the fit, all measured
at the time (t) of the survey.
Taking the difference between DIC at two times, t1 and t2, gives an equation
for the change in CANT over the time period between the two
hydrographic surveys (ΔCANT):
ΔCANT=a0t2-a0t1+(a1t2-a0t1)P1t2+…+(ant2-a0t1)Pn(t2)
The two-regression method was introduced by Thacker (2012) as an improvement
in regression-based methods. The region of study is first divided into
sub-regions since the empirical relationships between DIC and other
environmental variables vary spatially (Thacker, 2012). MLRs are investigated
between DIC and other measured variables (predictors) using a stepwise
technique. The procedure is applied for each sub-region using all data from
the repeated surveys within the period to be investigated, resulting in an
optimal MLR for each sub-region (similar to Eq. 6). A linear regression with
time is established for the residuals (observed DIC – predicted DIC) of the
regional fits, which directly gives ΔCANT averaged over
the space–time in each sub-region. The purpose of the first MLR is to remove
the natural variability of DIC, leaving the anthropogenic signal and noise
(random variability) in the residuals, and the second MLR is used to separate
the anthropogenic signal from the noise.
We applied both methodologies within the different water mass layers, which
were
separated by γn and used as sub-regions. The predictor variables
of θ, S, σ0, nitrate (NO3), SiO4 and AOU were
used for the MLR procedures. The three methodologies estimate similar rates
of increase in CANT for most water mass layers (Table 7). In the
STCW layer, the value of ΔCANT (eMLR) is higher than our
maximum estimate of ∂CANT∂t and the
value obtained from the two-regression method. The eMLR method is less
suitable for the upper layers of the ocean subject to high seasonal to
inter-annual variability, such as the STCW layer, resulting in large residuals
that bias the regression (Friis et al., 2005). For the AAIW layer, the
increase in CANT estimated by the two-regression method is half
the increase that is established by our method and the eMLR method. The lower
value of the trend estimated by the two-regression method is due to the fact
that the two-regression method uses stepwise MLR. This means that the
two-regression method only considers those predictors that give the best fit
while the eMLR method is forced to consider all the predictors in the fit.
This is also the cause of the low RMSE obtained with the two-regression
method compared to both our trends and those obtained by the eMLR method. For
deep to bottom layers, the two-regression method estimates a small increase in CANT similar to the one found in this study for the AABW layer
(Table 7) that can be considered negligible given the resolution of the
back-calculation method (±6 µmol kg-1). The eMLR method
finds increases in CANT (with relatively high uncertainties)
higher than the two-regression method that over the 16-year period also give
values of CANT change lower than the resolution of our
back-calculation method (although close to it for the LCDW layer, ∼ 5 µmol kg-1
for the 16-year period).
Circulation and biological processes at steady state
The back-calculation method assumes the circulation of the ocean and the
biological processes are at steady state. The contribution of non-linear
mixing is unknown, and some of the changes in DIC in the water mass layers
could be erroneously included in the estimates of CANT rather
than as a non-anthropogenic change in DIC. The non-steady state of the
circulation in our analysis is included to some extent through the changes in
CDISπ (Eqs. 3, 4), which is solved by the OMP analysis, which is
subjected to the limitations of quantifying the mixing mostly through
thermohaline changes in the water masses (Sect. S2).
For biological processes, remineralisation rates are usually considered to be
at steady state (Sarmiento et al., 1992). Climate change has been suggested
as potentially driving changes in carbon fixation and export that can
influence the uptake of CO2 by the oceans (Falkowski et al., 1998).
Pahlow and Riebesell (2000) first suggested that decadal changes in
remineralisation rates occurred in the deep waters of the Northern
Hemisphere, although this is still a matter of debate (e.g. Li and Peng,
2002; Najjar, 2009).
Metzl et al. (1999) and Shadwick et al. (2015) observed that the uptake of
CO2 over the sub-Antarctic zone (SAZ, between the SAF and the STF) in
summer is mostly controlled by biological processes. If a change in
remineralisation rates has occurred, i.e. the changes in DICBIO
are due to biological effects, a change in nutrient concentrations of the
water masses would be expected. The detection of long-term trends of
nutrients in upper layers of the ocean can be masked by short-timescale
physical processes such as changes in the mixed layer depth, mesoscale
activity and advection (Sambrotto and Mace, 2000; Rintoul and Trull, 2001;
Sallée et al., 2010). We did not find a measurable trend in nutrient
concentrations for any of the layers in the period 1995–2011, except for an
increase in SiO4 over the Antarctic Divergence (Table 4) that is most
likely due to the upwelling of SiO4-rich deep waters. We cannot
confidently assign the changes in DICBIO and its impact on
∂CANT∂t in upper layers
(Table 5) to any particular process, and instead we provide a range of values
for ∂CANT∂t (Table 4). For the
scenario of intensification of the Antarctic upwelling, the increase in
low-O2 and DIC-rich waters would increase the content of DIC in
subsurface waters, leading (at least for the AASW and AASWupw
layers) to values of ∂CANT∂t
closer to the lower limit of the range (i.e. ∂DIC∂t≫∂CANT∂t, Table 4).
In deep layers of the section, the increase in DIC is not explained by the
long-term change in any of the terms in Eq. (1), which is another
implication of considering the circulation at steady state. The differences
found in the increase in DIC in the UCDW layer north and south of the SAF
(Sect. 4.3, Fig. 4) add consistency to the idea of the advection of older
waters to the section. Considering these differences, we can assign a change
in DIC of at least ∼ 0.20 ± 0.02 µmol kg-1 yr-1
(lower rate of increase in DIC found in deep–bottom layers, Table 4) due to the
upwelling intensification. Since the AASW and AASWupw
layers are the most affected by the upwelling, we can correct the values of
∂CANT∂t in these layers for this
effect (trends with ** in Table 4).
Changes in the rates of export of particulate organic carbon and
silicate from surface layers
We assume that the export of particulate organic carbon (POC), from upper
ocean layers (STCW, AASW and AASWupw) and remineralisation in the water
column, was constant between 1995 and 2011. The high latitudes are considered
important in terms of POC export, mostly because these areas are dominated
by large phytoplankton, in particular diatoms (Buesseler, 1998; Sambrotto
and Mace, 2000), and rapid export of carbon to deep waters from phytoplankton
blooms is possible (DiTullio et al., 2000; Lourey and Trull, 2001). The POC
exported is remineralised to DIC below the mixed layer (Wassman et al.,
1990; Asper and Smith Jr., 1999; Trull et al., 2001a; Fripiat et al., 2015).
Our results show that the increase in DIC in mode and intermediate waters is
fully explained by the uptake of atmospheric CO2, which could indicate
that there was no detectable change in the rate of export of POC over the
1995–2011 period. Estimates of POC export in the SAZ (∼ 3800 m), the
SAF (∼ 3100 m) and the PFZ (∼ 1500 m) using moored sediment
traps near the section (Trull et al., 2001b) were 0.5, 0.8 and 1.0 g C m-2 yr-1,
respectively. If all this POC is fully remineralised
in the UCDW layer (with a mean thickness of 2000 m), we obtain a range of 0.02–0.04 µmol kg-1 yr-1
for the maximum increase in DIC
due to the export of POC. This increase is close to the uncertainty of the
total DIC increase estimated for the UCDW layer, which means that in order to
generate an increase in DIC similar to that found in the UCDW layer, the rate
of POC export should be ∼ 10 times higher than the observed rates. This
change should certainly be noticeable in ∂DICBIO∂t in surface waters but most probably in deep waters
as well, which we do not see.
The observed increase in SiO4 in deep and bottom layers of the ocean is
consistent with the transport of SiO4-rich older waters to the section.
For UCDW, the increase in SiO4 north of the SAF is higher than at
southern latitudes (0.31 ± 0.08 µmol kg-1 yr-1
with respect to 0.19 ± 0.04 µmol kg-1 yr-1, not
shown). Nelson et al. (1995) observed a lower dissolution rate of diatom-dominated SiO4 exported in high-latitude regions compared to lower
latitudes. Nelson et al. (1995) estimated a mean silica production rate of
0.7–1.2 mol Si m-2 yr-1 for regions over diatomaceous
sediments and concluded that 15–25 % of the silica produced in the
upper ocean accumulates in the seabed. Of the silica produced in the mixed
layer of the upper ocean layers, at least 50 % is believed to dissolve in
the upper 100 m of the water column (e.g. Nelson et al., 1991; DeMaster et
al., 1992). For deep waters, the production rates of Nelson et al. (1995)
could result in a mean increase in SiO4 of 0.08–0.14 µmol kg-1 yr-1
in the water column (∼ 3400 m, mean depth). The maximum value of this increase could explain the trends
of SiO4 found for the AABW layer (Table 4), but 32–48 % of the
increase in SiO4 in deep bottom layers is not explained by the
remineralisation of exported silica and is most likely the result of the
advection of older SiO4-rich waters to the section.
Stoichiometric ratios for biological processes
The back-calculation method and the OMP analysis assume constant
stoichiometric ratios for remineralisation. The theoretical Redfield ratios
(Redfield, 1934, 1958) are usually considered as a mean for the whole ocean,
although they can vary from the theoretical value due to changes in
phytoplankton species composition, the food-web structure and nutrient
availability (Martiny et al., 2013).
We carried out a sensitivity analysis on the Redfield ratios following
Álvarez et al. (2014), to obtain values of stoichiometric ratios for the
section. A battery of OMP analyses were done with varying values of
RN between 9 and 10 in increments of 0.2, RP between
120 and 145 in increments of 5, and RSi between 0 and 8 in
increments of 2. For each variation in the stoichiometric ratios, an OMP
analysis was made for each section in order to determine best-fit R values
and whether there were differences in time for the stoichiometric ratios
along the sections. The smallest residuals (differences between the nutrients
measured and those estimated by the OMP analysis) were obtained for
RN= 9 and RP= 125. The residuals of SiO4
did not change significantly for any value of RSi and we consider
SiO4 to be a conservative variable. These results indicate
RPRN= 13.8, in agreement with values
obtained for the region RPRNϵ[8–15],LoureyandTrull,2001.
Conclusions
The results of our analysis south of Tasmania over the 1995–2011 period
support a scenario of intensification of upwelling in the vicinity of the
Antarctic Divergence due to an increase in the westerly winds at high
latitudes, most probably linked to the variability of the SAM. The
intensification of the upwelling favours the advection of older waters to
deep–bottom layers of the section where we found a net increase in DIC over the
16-year period. The enhanced upwelling causes the eventual entrainment of
low-O2 and DIC-rich waters into upper layers, explaining the trends of
decreasing O2 and increasing non-anthropogenic DIC found in surface
waters close to the Antarctic Divergence. This scenario also implies the
intensification of the convergence north of the SAF, implying a more
efficient ventilation of the SAMW and AAIW layers and thus an efficient
uptake of atmospheric CO2 by these layers. The enhanced upwelling lowers
the uptake of CANT in the AASW layer, but the effect of ventilation
more than compensates that of the upwelling, allowing the increase in
CANT in this layer. The atmospheric warming reduces the
dissolution of CO2 in upper layers north of the PF, presenting increases
in DIC lower than expected from the atmospheric CO2 increase.
Our results rely on a limited number of sections that are studied every 3–7 years and
can only provide a long-term (decadal) average view of changes in water
masses. More surface observations and repeated deep-ocean sections are needed
to help resolve inter-annual changes in the Southern Ocean carbon sink and to
determine the main drivers and feedback to the carbon–climate system. The
effort to maintain hydrographic sections with CO2 system measurements
would also benefit from additional direct measurements of more variables of
the carbon system, e.g. pH, which has not been measured in the SR03 section.
The section data are available through the Global Ocean
Data Analysis Project
(https://www.nodc.noaa.gov/ocads/oceans/GLODAPv2/; Key et al., 2015;
Olsen et al., 2016). The original data for the different cruises were
corrected following the QC recommendations in GLODAPv2.
The Supplement related to this article is available online at https://doi.org/10.5194/bg-14-5217-2017-supplement.
The authors declare that they have no conflict of
interest.
This article is part of the special issue “The Ocean in a
High-CO2 World IV”. It is a result of the 4th International Symposium
on the Ocean in a High–CO2 World, Hobart, Australia, 3–6 May 2016.
Acknowledgements
The SR03 section was sampled as part of the World Ocean Circulation
Experiment CO2 Survey (WOCE, http://woceatlas.ucsd.edu), and more
recently the Global Ocean Ship-Based Hydrographic Investigation Program
(GO-SHIPS, http://www.go-ship.org). Carbon system parameters contribute
to the International Ocean Carbon Coordination Project of the United Nations
Intergovernmental Oceanographic Commission (IOCCP,
http://www.ioccp.org). Support for measurements on the section were
provided to Steve Rich Rintoul and Bronte Tilbrook by the Antarctic Climate
and Ecosystems Cooperative Research Centre (ACE CRC) and the Australian
Climate Change Science Program. Logistic support for the section and ship
time on the RSV Aurora Australis was provided by the Australian Antarctic
Division. The many scientific staff involved in the hydrographic sections and
the officers and crew of the ship were critical to obtaining good quality
data. We especially want to thank the work by Kate Berry and Mark Pretty for
high quality DIC and TA data, and for the hydrochemistry teams and CTD
watches of multiple cruises and especially Mark Rosenberg and Rebecca Cowley.
Paula Conde Pardo is a postdoctoral fellow supported by the ACE-CRC Project
R2.1: Carbon Uptake and Chemical Change.
Edited by: Kristy Kroeker Reviewed by: Marta Álvarez and
one anonymous referee
ReferencesÁlvarez, M., Brea, S., Mercier, H., and Álvarez-Salgado, X. A.:
Mineralization of biogenic materials in the water masses of the South
Atlantic Ocean. I: assessment and results of an optimum multiparameter
analysis, Prog. Oceanogr., 123, 1–23, 10.1016/j.pocean.2013.12.007, 2014.Amante, C. and Eakins, B. W.: ETOPO1 1 Arc-Minute Global Relief Model:
Procedures, Data Sources and Analysis, NOAA Technical Memorandum NESDIS
NGDC-24, National Geophysical Data Center, NOAA, 10.7289/V5C8276M,
2009.Anderson, L. A. and Sarmiento, J. L.: Redfield ratios of remineralization
determined by nutrient data analysis, Global Biogeochem. Cy., 8,
65–80, 10.1029/93GB03318, 1994.Aoki, S., Bindoff, N. L., and Church, J. A.: Interdecadal water mass changes in
the Southern Ocean between 30∘ E and 160∘ E, 32, L07607,
10.1029/2004GL022220, 2005.Aoki, S., Mizuta, G., Sasaki, H., Sasai, Y., Rintoul, S. R., and Bindoff, N.
L.: Atlantic–Pacific asymmetry of subsurface temperature change and frontal
response of the Antarctic Circumpolar Current for the recent three decades,
J. Oceanogr., 71, 623–636, 10.1007/s10872-015-0284-6, 2015.
Armour, K. and Bitz, C. M.: Observed and projected trends in Antarctic sea
ice, US Clivar Variations Newsletter, 13, 12–19, 2015.
Asper, V. L. and Smith Jr., W. O.: Particle fluxes during austral spring and
summer in the southern Ross Sea, Antarctica, J. Geophys. Res., 104,
5345–5359, 1999.
Baines, P. G., Edwards, R. J., and Fandry, C. B.: Observations of a new
baroclinic current along the western continental slope of Bass Strait, Aust.
J. Mar. Fresh. Res., 34, 155–157, 1983.Bates, N. R., Astor, Y. M., Church, M. J., Currie, K., Dore, J. E.,
González-Dávila, M., Lorenzoni, L., Muller-Karger, F., Olafsson, J.,
and Santana-Casiano, J. M.: A time-series view of changing ocean chemistry
due to ocean uptake of anthropogenic CO2 and ocean acidification,
Oceanography 27, 126–141, 10.5670/oceanog.2014.16, 2014.
Belkin, I. M. and Gordon, A. L.: Southern Ocean fronts from the Greenwich
meridian to Tasmania, J. Geophys. Res., 101, 3675–3696,
1996.Bender, M., Ellis, T., Tans, P., Francey, R., and Lowe, D.: Variability in
the O2/ N2 ratio of southern hemisphere air, 1991–1994:
Implications for the carbon cycle, Global Biogeochem. Cy., 10, 9–21, 1996.
Bindoff, N. L. and Church, J. A.: Warming of the Water Column in the Southwest
Pacific Ocean, Nature, 357, 59–62, 1992.
Bindoff, N. L. and McDougall, T. J.: Diagnosing Climate Change and Ocean
Ventilation using Hydrographic Data, J. Phys. Oceanogr., 24,
1137–1152, 1994.Boland, F. M. and Church, J. A.: The East Australian Current 1978, Deep-Sea
Res., 28, 937–957, 10.1016/0198-0149(81)90011-X, 1981.Bostock, H. C., Mikaloff Fletcher, S. E., and Williams, M. J. M.: Estimating
carbonate parameters from hydrographic data for the intermediate and deep
waters of the Southern Hemisphere oceans, Biogeosciences, 10, 6199–6213,
10.5194/bg-10-6199-2013, 2013.
Broecker, W. S.: “NO” a conservative water mass tracer, Earth Planet. Sc.
Lett., 23, 8761–8776, 1974.
Buesseler, K. O.: The decoupling of production and particle export in the
surface ocean, Global Biogeochem. Cy., 12, 297–310, 1998.
Callahan, J. E.: The structure and circulation of Deep Water in the
Antarctic, Deep-Sea Res., 19, 563–575, 1972.Carter, B. R., Feely, R. A., Mecking, S., Cross, J. N., Macdonald, A. M.,
Siedlecki, S. A., Talley, L. D., Sabine, C. L., Millero, F. J., Swift, J. H.,
Dickson, A. G., and Rodgers, K. B.: Two decades of Pacific anthropogenic
carbon storage and ocean acidification along Global Ocean Ship-based
Hydrographic Investigations Program sections P16 and P02, Global Biogeochem.
Cy., 31, 306–327, 10.1002/2016GB005485, 2017.Chen, C.-T. A. and Millero, F. J.: Gradual increase of oceanic CO2,
Nature, 277, 205–206, 1979.
Chen, C.-T. A., Pytkowicz, M. R., and Olson, E. J.: Evaluation of the calcium
problem in the South Pacific, Geochem. J., 16, 1–10, 1982.
Davis, R.: Intermediate-depth circulation of the Indian and South Pacific
oceans measured by autonomous floats, J. Phys. Oceanogr., 35, 683–707,
2005.
Deacon, G. E. R.: The hydrology of the Southern Ocean, Cambridge University
Press, 15, 1–124, 1937.
DeMaster, D. J., Dunbar, R. B., Gordon, L. I., Leventer, A. R., Morrison, J.
M., Nelson, D. M., Nittrouer, C. A., and Smith Jr., W. O.: The cycling and
accumulation of organic matter and biogenic silica in high-latitude
environments: The Ross Sea, Oceanography, 5, 146–153, 1992.DeVries, T., Holzer, M., and Primeau, F.: Recent increase in oceanic carbon
uptake driven by weaker upper-ocean overturning, Nature, 542, 215–218,
10.1038/nature21068, 2017.Dickson, A. G.: Thermodynamics of the dissociation of boric acid in synthetic
seawater from 273.15 to 318.15 K, Deep-Sea Res. Pt. I, 37, 755–766,
10.1016/0198-0149(90)90004-F, 1990.Dickson, A. G. and Millero, F. J.: A comparison of the equilibrium constants
for the dissociation of carbonic acid in seawater media, Deep-Sea Res. Pt. I,
34, 1733–1743, 10.1016/0198-0149(87)90021-5, 1987.Dickson, A. G., Sabine, C. L., and Christian, J. R.: Guide to Best Practices
for Ocean CO2 Measurements, PICES Special Publication 3, 191 pp., 2007.Dickson, R. R. and Brown, J.: The production of North Atlantic Deep Water:
sources, rates, and pathways, J. Geophys. Res., 99, 12319–12341,
10.1029/94JC00530, 1984.
DiTullio, G. R., Grebmeier, J. M., Arrigo, K. R., Lizotte, M. P., Robinson,
D. H., Leventer, A., Barry, J. P., VanWoert, M. L., and Dunbar, R. B.: Rapid
and early export of Phaeocystis antarctica blooms in the Ross Sea,
Antarctica, Nature, 404, 595–598, 2000.Dlugokencky, E. J., Lang, P. M., Mund, J. W., Crotwell, A. M., Crotwell, M.
J., and Thoning, K. W.: Atmospheric Carbon Dioxide Dry Air Mole Fractions
from the NOAA ESRL Carbon Cycle Cooperative Global Air Sampling Network,
1968–2015, Version: 2016-08-30,
ftp://aftp.cmdl.noaa.gov/data/trace_gases/co2/flask/surface/ (last
access: 14 November 2017), 2016.Doney, S. C., Victoria, J. F., Feely, R. A., and Kleypas, J. A.: Ocean
Acidification: The other CO2 problem, Annu. Rev. Mar. Sci., 1, 169–92,
10.1146/annurev.marine.010908.163834, 2009.Falkowski, P. G., Barber, R. T., and Smetacek, V.: Biogeochemical Controls
and Feedbacks on Ocean Primary Production, Science, 281, 200–206,
10.1126/science.281.5374.200, 1998.Fay, A. R., McKinley, G. A., and Lovenduski, N. S.: Southern Ocean carbon
trends: Sensitivity to methods, Geophys. Res. Lett., 41, 6833–6840,
10.1002/2014GL061324, 2014.Feely, R. A., Sabine, C. L., Lee, K., Berelson, W., Kleypas, J., Fabry, V.
J., and Millero, F. J.: The impact of anthropogenic CO2 on the
CaCO3 system in the oceans, Sience, 305, 362–366, 2004.
Foster, T. D. and Carmack, E. C.: Frontal zone mixing and Antarctic Bottom
Water formation in the southern Weddell Sea, Deep-Sea Res., 23, 301–307,
1976.Friis, K., Körtzinger, A., J. Pätsch, J., and Wallace, D. W. R.: On
the temporal increase of anthropogenic CO2 in the subpolar North
Atlantic, Deep-Sea Res. Pt. I, 52, 681–698, 10.1016/j.dsr.2004.11.017,
2005.Fripiat, F., Elskens, M., Trull, T. W., Blain, S., Cavagna, A.-J., Fernandez,
C., Fonseca-Batista, D., Planchon, F., Raimbault, P., Roukaerts, A., and
Dehairs, F.: Significant mixed layer nitrification in a natural
iron-fertilized bloom of the Southern Ocean, Global Biogeochem. Cy., 29,
1929–1943, 10.1002/2014GB005051, 2015.
Fukamachi, Y., Rintoul, S. R., Church, J. A., Aoki, S., Sokolov, S.,
Rosenberg, M. A., and Wakatsuchi, M.: Strong export of Antarctic Bottom Water
east of the Kerguelen plateau, Nature, 3, 327–331, 2010.
Gordon, A. L. and Tchernia, P.: Waters of the continental margin off
Adélie Coast, Antarctica, in: Antarctic Oceanology II: The
Australian–New Zealand Sector, edited by: Hayes, D. E., Antarctic Research
Series 19, American Geophysical Union, Washington, DC, 59–69, 1972.Gruber, N.: Anthropogenic CO2 in the Atlantic Ocean, Global Biogeochem.
Cy., 12, 165–191, 1998.Gruber, N., Gloor, M., Mikaloff Fletcher, S. E., Doney, S. C., Dutkiewicz,
S., Follows, M. J., Gerber, M., Jacobson, A. R., Joos, F., Lindsay, K.,
Menemenlis, D., Mouchet, A., Müller, S. A., Sarmiento, J. L., and
Takahashi, T.: Oceanic sources, sinks, and transport of atmospheric CO2,
Global Biogeochem. Cy., 23, GB1005, 10.1029/2008GB003349, 2009.
Hanawa, K. and Talley, L. D.: Mode waters, in: Ocean Circulation and Climate,
edited by: Siedler, G., Church, J., and Gould, J., International Geophysics
Series, Academic Press, New York, 373–386, 2001.Hauri, C., Doney, S. C., Takahashi, T., Erickson, M., Jiang, G., and Ducklow,
H. W.: Two decades of inorganic carbon dynamics along the West Antarctic
Peninsula, Biogeosciences, 12, 6761–6779,
10.5194/bg-12-6761-2015, 2015.Helm, K. P., Bindoff, N. L., and Church, J. A.: Observed decreases in oxygen
content of the global ocean, Geophys. Res. Lett., 38, L23602,
10.1029/2011GL049513, 2011.Herraiz-Borreguero, L. and Rintoul, S. R.: Regional circulation and its
impact on upper ocean variability south of Tasmania, Deep-Sea Res. Pt. II,
58, 2071–2081, 10.1016/j.dsr2.2011.05.022, 2011.Hill, K. L., Rintoul, S. R., Ridgway, K. R., and Oke, P. R.: Decadal changes
in the South Pacific western boundary current system revealed in observations
and ocean state estimates, J. Geophys. Res., 116, C01009,
10.1029/2009JC005926, 2011.Hood, E. M., Sabine, C. L., and Sloyan, B. M.: The GO-SHIP Repeat Hydrography
Manual: a Collection of Expert Reports and Guidelines, IOCCP Report Number
14, OCPO Publication Series Number 134,
http://www.go-ship.org/HydroMan.html (last access: 14 November 2017),
2010.
Ikegami, H. and Kanamori, S.: Calcium-Alkalinity-Nitrate Relationship in the
North Pacific and the Japan Sea, Journal of the Oceanographical Society of
Japan, 39, 9–14, 1983.Iudicone, D., Speich, S., Madec, G., and Blanke, B.: The Global Conveyor Belt
from a Southern Ocean perspective, J. Phys. Oceanogr., 38, 1401–1425,
10.1175/2007JPO3525.1, 2008.
Jackett, D. R. and McDougall, T. J.: A Neutral Density Variable for the
World's Oceans, J. Phys. Oceanogr., 27, 237–263, 1997.Jacobs, S.: Observations of change in the Southern Ocean, Philos. T. R. Soc.
A, 364, 1657–1681, 10.1098/rsta.2006.1794, 2006.
Jacobs, S. S.: On the nature and significance of the Antarctic Slope Front,
Mar. Chem., 35, 9–24, 1991.Johnson, G. C.: Quantifying Antarctic Bottom Water and North Atlantic Deep
Water volumes, J. Geophys. Res., 113, C05027, 10.1029/2007JC004477, 2008.
Joyce, T. and Corry, C.: Requirements for WOCE Hydrographic Programme Data
Reporting, WHPO Publication 90-1 Revision 2, WOCE Report 67/91, Woods Hole,
Mass., USA, 1994.Key, R. M., Olsen, A., van Heuven, S., Lauvset, S. K., Velo, A., Lin, X.,
Schirnick, C., Kozyr, A., Tanhua, T., Hoppema, M., Jutterström, S.,
Steinfeldt, R., Jeansson, E., Ishi, M., Perez, F. F., and Suzuki, T.: Global
Ocean Data Analysis Project, Version 2 (GLODAPv2), ORNL/CDIAC-162, ND-P093,
Carbon Dioxide Information Analysis Center, Oak Ridge National Laboratory, US
Department of Energy, Oak Ridge, Tennessee,
10.3334/CDIAC/OTG.NDP093_GLODAPv2, 2015.Khatiwala, S., Primeau, F., and Hall, T.: Reconstruction of the history of
anthropogenic CO2 concentrations in the ocean, Nature, 462, 346–350,
10.1038/nature08526, 2009.Kouketsu, S. and Murata, A. M.: Detecting decadal scale increases in
anthropogenic CO2 in the ocean, Geophys. Res. Lett., 41, 4594–4600,
10.1002/2014GL060516, 2014.Lacarra, M., Houssais, M.-N., Sultan, E., Rintoul, S. R., and Herbaut, C.:
Summer hydrography on the shelf off Terre Adélie/George V Land based on
the ALBION and CEAMARC observations during the IPY, Polar Sci., 5, 88–103,
10.1016/j.polar.2011.04.008, 2011.Landschützer, P., Gruber, N., Haumann, A., Rödenbeck, C., Bakker, D.
C. E., van Heuven, S., Hoppema, M., Metzl, N., Sweeney, C., Takahashi, T.,
Tilbrook, B., and Wanninkhof, R.: The reinvigoration of the SouthernOcean
carbon sink, Science, 349, 1221–1224, 10.1126/science.aab2620, 2015.Lauvset, S. K., Gruber, N., Landschützer, P., Olsen, A., and Tjiputra, J.:
Trends and drivers in global surface ocean pH over the past 3 decades,
Biogeosciences, 12, 1285–1298, 10.5194/bg-12-1285-2015,
2015.Lenton, A. and Matear, R. J.: Role of the Southern Annular Mode (SAM) in
Southern Ocean CO2 uptake, Global Biogeochem. Cy., 21, GB2016,
10.1029/2006GB002714, 2007.Lenton, A., Metzl, N., Takahashi, T., Kuchinke, M., Matear, R. J., Roy, T.,
Sutherland, S. C., Sweeney, C., and Tilbrook, B.: Global Biogeochem. Cy., 26,
GB2021, 10.1029/2011GB004095, 2012.Lenton, A., Tilbrook, B., Law, R. M., Bakker, D., Doney, S. C., Gruber, N.,
Ishii, M., Hoppema, M., Lovenduski, N. S., Matear, R. J., McNeil, B. I.,
Metzl, N., Mikaloff Fletcher, S. E., Monteiro, P. M. S., Rödenbeck, C.,
Sweeney, C., and Takahashi, T.: Sea-air CO2 fluxes in the Southern Ocean
for the period 1990–2009, Biogeosciences, 10, 4037–4054,
10.5194/bg-10-4037-2013, 2013.Le Quéré, C., Rödenbeck, C., Buitenhuis, E. T., Conway, T. J.,
Langenfelds, R., Gomez, A., Labuschagne, C., Ramonet, M., Nakazawa, T., Metz,
N., Gillett, N., and Heimann, M.: Saturation of the Southern Ocean CO2
Sink Due to Recent Climate Change, Science, 316, 1735–1738,
10.1126/science.1136188, 2007.Lewis, E. and Wallace, D. W. R.: Program Developed for CO2 System
Calculations. ORNL/CDIAC-105. Carbon Dioxide Information Analysis Center, Oak
Ridge National Laboratory, U.S. Department of Energy, Oak Ridge, Tennessee,
1998.Li, Y.-H. and Peng, T.-H.: Latitudinal change of remineralization ratios in
the oceans and its implication for nutrient cycles, Global Biogeochem. Cy.,
16, 1130, 10.1029/2001GB001828, 2002.
Lourey, K. J. and Trull, T. W.: Seasonal nutrient depletion and carbon export
in the Subantarctic and Polar Frontal Zones of the Southern Ocean south of
Australia, J. Geophys. Res., 106, 31463–31487, 2001.Lovenduski, N. S., Gruber, N., and Doney, S. C.: Towards a mechanistic
understanding of the decadal trends in the Southern Ocean carbon sink, Global
Biogeochem. Cy., 22, GB3016, 10.1029/2007GB003139, 2008.Lumpkin, R. and Speer, K.: Global Ocean Meridional Overturning, J. Phys.
Oceanogr., 37, 2550–2562, 10.1175/JPO3130.1, 2007.
Mantyla, A. W. and Reid, J. L.: On the origins of deep and bottom waters of
the Indian Ocean, J. Geophys. Res., 100, 2417–2439, 1995.Marshall, G. J.: Analysis of recent circulation and thermal advection change
in the northern Antarctic Peninsula, Int. J. Climatol., 22, 1557–1567,
10.1002/joc.814, 2002.
Marshall, G. J.: Trends in the Southern Annular Mode from observations and
reanalyses, J. Climate, 16, 4134–4143, 2003.Marsland, S. J., Bindoff, N. L., Williams, G. D., and Budd, W. F.: Modeling
water mass formation in the Mertz Glacier Polynya and Adélie Depression,
East Antarctica, J. Geophys. Res., 109, C11003, 10.1029/2004JC002441,
2004.Martiny, A. C., Pham, C. T. A., Primeau, F. W., Vrugt, J. A., Moore, J. K.,
Levin, S. A., and Lomas, M. W.: Strong latitudinal patterns in the elemental
ratios of marine plankton and organic matter, Nature, 6, 279–283,
10.1038/NGEO1757, 2013.Matear, R. and Lenton, A.: Impact of historical climate change on the
Southern Ocean carbon cycle, J. Climate, 21, 5820–5834,
10.1175/2008JCLI2194.1, 2008.
Matear, R. J., Hirst, A. C., and McNeil, B. I.: Changes in dissolved oxygen
in the Southern Ocean with climate change, Geochem. Geophy. Geosy., 1,
1525–2027, 2000.
McCartney, M. S.: The subtropical recirculation of mode waters, J. Mar. Res.,
40 (Suppl.), 427–464, 1977.
McDougall, T. J.: Neutral surfaces in the ocean: implications for modelling,
Geophys. Res. Lett., 14, 97–800, 1987.McNeil, B. I., Tilbrook, B., and Matear, R. J.: Accumulation and uptake of
anthropogenic CO2 in the Southern Ocean, south of Australia between 1968
and 1996, J. Geophys. Res., 106, 31431–31445, 2001.Mehrbach, C., Culberson, C. H., Hawley, J. E., and Pytkowicz, R. M.:
Measurement of the apparent dissociation constants of carbonic acid in
seawater at atmospheric pressure, Limnol. Oceanogr., 18, 897–907,
10.4319/lo.1973.18.6.0897, 1973.Metzl, N., Tilbrook, B., and Poisson, A.: The annual fCO2 cycle and the
air–sea CO2 flux in the sub-Antarctic Ocean, Tellus, 51, 849–861,
1999.
Moore, J. K. and Abbott, M. R.: Phytoplankton chlorophyll distributions and
primary production in the Southern Ocean, J. Geophys. Res., 105,
28709–28722, 2000.
Moore, J. K., Abbott, M. R., and Richman, J. R.: Location and dynamics of the
Antarctic Polar Front from satellite sea surface temperature data, J.
Geophys. Res., 104, 3059–3073, 1999.
Mosby, H.: The waters of the Atlantic Antarctic Ocean, Scientific Results of
the Norwegian Antarctic Expeditions 1927–1928, 1, 11, 131 pp., 1934.Murata, A., Kumamoto, Y., Watanabe, S., and Fukasawa, M.: Decadal increases
of anthropogenic CO2 in the South Pacific subtropical ocean along
32∘ S, J. Geophys. Res., 112, C05033, 10.1029/2005JC003405,
2007.Najjar, R.: The dark side of marine carbon, Nat. Geosci., 2, 603–604,
10.1038/NGEO812, 2009.
Nelson, D. M., Ahern, J. A., and Herlihy, L. J.: Cycling of biogenic silica
within the upper water column of the Ross Sea, Mar. Chem, 35, 461–476, 1991.
Nelson, D. M., Tréguer, P., Brzezinsk, M. A., Leynaert, A., and
Quéguiner, B.: Production and dissolution of biogenic silica in the
ocean: Revised global estimates, comparison with regional data and
relationship to biogenic sedimentation, Global Biogeochem. Cy., 9, 359–372,
1995.Olsen, A., Key, R. M., van Heuven, S., Lauvset, S. K., Velo, A., Lin, X.,
Schirnick, C., Kozyr, A., Tanhua, T., Hoppema, M., Jutterström, S.,
Steinfeldt, R., Jeansson, E., Ishii, M., Pérez, F. F., and Suzuki, T.:
The Global Ocean Data Analysis Project version 2 (GLODAPv2) – an internally
consistent data product for the world ocean, Earth Syst. Sci. Data, 8,
297–323, 10.5194/essd-8-297-2016, 2016.Orr, J. C., Fabry, V. J., Aumont, O., Bopp, L., Doney, S. C., Feely, R. A.,
Gnanadesikan, A., Gruber, N., Ishida, A., Joos, F., Key, R. M., Lindsay, K.,
Maier-Reimer, E., Matear, R., Monfray, P., Mouchet, A., Najjar, R. J.,
Plattner, G.-K., Rodgers, K. B., Sabine, C. L., Sarmiento, J. L., Schlitzer,
R., Slater, R. D., Totterdell, I. J., Weirig, M.-F., Yamanaka, Y., and Yool,
A.: Anthropogenic ocean acidification over the twenty-first century and its
impact on calcifying organisms, Nature, 437, 681–686,
10.1038/nature04095, 2005.
Orsi, A. H., Whitworth III, T., and Nowlin, W. D.: On the meridional extent
and fronts of the Antarctic Circumpolar Current, Deep-Sea Res. Pt. I, 42,
641–673, 1995.
Pahlow, M. and Riebesell, U.: Temporal trends in deep ocean Redfield ratios,
Science, 287, 831–833, 2000.Pardo, P. C., Vázquez-Rodríguez, M., Pérez, F. F., and
Ríos, A. F.: CO2 air-sea disequilibrium and preformed alkalinity in
the Pacific and Indian Oceans calculated from subsurface layer data, J.
Marine Syst., 84, 67–77, 10.1016/j.jmarsys.2010.08.006, 2011.Pardo, P. C., Pérez, F. F., Khatiwala, S., and Ríos, A. F.:
Anthropogenic CO2 estimates in the Southern Ocean: Storage partitioning
in the different water masses, Prog. Oceanogr., 120, 230–242,
10.1016/j.pocean.2013.09.005, 2014.Peña-Molino, B., Rintoul, S. R., and Mazloff, M. R.: Barotropic and
baroclinic contributions to along-stream and across-stream transport in the
Antarctic Circumpolar Current, J. Geophys. Res.-Oceans, 119, 8011–8028,
10.1002/2014JC010020, 2014.Purkey, S. G. and Johnson, G. C.: Global Contraction of Antarctic Bottom
Water between the 1980s and 2000s, J. Climate, 25, 5830–5844,
10.1175/JCLI-D-11-00612.1, 2012.
Redfield, A.: On the proportions of organic derivatives in sea water and
their relation to the composition of plankton, in: James Johnstone Memorial
Volume, edited by: Daniel, R. J., University Press of Liverpool, 177–192,
1934.
Redfield, A.: The biological control of chemical factors in the environment,
Am. Sci., 46, 205–221, 1958.Ridgway, K. R.: Seasonal circulation around Tasmania: An interface between
eastern and western boundary dynamics, J. Geophys. Res., 112, C10016,
10.1029/2006JC003898, 2007.
Rintoul, S. R.: On the origin and influence of Adelie Land Bottom Water, in:
Ocean, Ice, and Atmosphere: Interactions at the Antarctic continental margin,
edited by: Jacobs. S. and Weiss, R., Antarct. Res. Ser., 75, 151–171, 1998.
Rintoul, S. R. and Bullister, J. L.: A late winter hydrographic section from
Tasmania to Antarctica, Deep–Sea Res. Pt. I, 46, 1417–1454, 1999.
Rintoul, S. R. and England, M. H.: Ekman Transport Dominates Local Air–Sea
Fluxes in Driving Variability of Subantarctic Mode Water, J. Phys. Oceanogr.,
32, 1308–1321, 2002.
Rintoul, S. R. and Sokolov, S.: Baroclinic transport variability of the
Antarctic Circumpolar Current south of Australia (WOCE repeat section SR3),
J. Gephys. Res., 106, 2815–2832, 2001.
Rintoul, S. R. and Trull, T. W.: Seasonal evolution of the mixed layer in the
Subantarctic Zone south of Australia, J. Geophys. Res., 106, 31447–31462,
2001.
Rintoul, S. R., Donguy, J. R., and Roemmich, D. H.: Seasonal evolution of
upper ocean thermal structure between Tasmania and Antarctica, Deep-Sea Res.
Pt. I, 44, 1185–1202, 1997.Sabine, C. L., Feely, R. A., Gruber, N., Key, R. M., Lee, K., Bullister, J.
L., Wanninkhof, R., Wong, C. S., Wallace, D. W. R., Tilbrook, B., Millero, F.
J., Peng, T.-H., Kozyr, A., Ono, T., and Rios, A. F.: The Oceanic Sink for
Anthropogenic CO2, Science, 305, 367–371, 2004.Sabine, C. L., Feely, R. A., Millero, F. J., Dickson, A. G., Langdon, C.,
Mecking, S., and Greeley, D.: Decadal changes in Pacific carbon, J. Geophys.
Res., 113, C07021, 10.1029/2007JC004577, 2008.Sallée, J.-B., Wienders, N., Speer, K., and Morow, R.: Formation of
subantarctic mode water in the southeastern Indian Ocean, Ocean Dynam., 56,
525–542, 10.1007/s10236-005-0054-x, 2006.Sallée, J. B., Speer, K., and Morrow, R.: Response of the Antarctic
Circumpolar Current to Atmospheric Variability, J. Climate, 21, 3020–3039,
10.1175/2007JCLI1702.1, 2008.
Sallée, L. B., Speer, K. G., and Rintoul, S. R.: Zonally asymmetric
response of the Southern Ocean mixed-layer depth to the Southern Annular
Mode, Nat. Geosci., 3, 273–279, 2010.Sallée, J. B., Matear, R. J., Rintoul, S. R., and Lenton, A.: Localized
subduction of anthropogenic carbon dioxide in the Southern Hemisphere oceans,
Nature, 5, 579–584, 10.1038/NGEO1523, 2012.
Sambrotto, R. N. and Mace, B. J.: Coupling of biological and physical regimes
across the Antarctic Polar Front as reflected by nitrogen production and
recycling, Deep-Sea Res. Pt. II, 47, 3339–3367, 2000.
Sarmiento, J. L. and Sundquist, E. T.: Revised budget for the oceanic uptake
of anthropogenic carbon dioxide, Nature, 356, 589–593, 1992.
Sarmiento, J. L., Hughes, T. M. C., Stouffer, R. J., and Manabe, S.:
Simulated response of the ocean carbon cycle to anthropogenic climate
warming, Nature, 393, 245–249, 1998.
Sarmiento, J. L., Gruber, N., Brzezinski, M. A., and Dunne, J. P.:
High-latitude controls of thermocline nutrients and low latitude biological
productivity, Nature, 427, 56–60, 2004.Shadwick, E. H., Trull, T. W., Tilbrook, B., Sutton, A. J., Schulz, E., and
Sabine, C. L.: Seasonality of biological and physical controls on surface
ocean CO2 from hourly observations at the Southern Ocean Time Series
site south of Australia, Global Biogeochem. Cy., 29 223–238,
10.1002/2014GB004906, 2015.
Sloyan, B. M. and Rintoul, S. R.: The Southern Ocean Limb of the Global Deep
Overturning Circulation, J. Phys. Oceanogr., 31, 143–173, 2001.Sloyan, B. M., Ridgway, K., and Cowley, R.: The East Australian Current and
Property Transport at 27∘ S from 2012 to 2013, J. Phys. Oceanogr., 46, 993–1008, 10.1175/JPO-D-15-0052.1, 2016.Sokolov, S. and Rintoul, S. R.: Structure of Southern Ocean fronts at
140∘ E, J. Marine Syst., 37, 151-184, 2002.Sokolov, S. and Rintoul, S. R.: Multiple jets of the Antarctic Circumpolar
Current South of Australia, J. Phys. Oceanog. 37, 1394-1412, 10.1175/JPO3111.1, 2007.Sokolov, S. and Rintoul, S. R.: Circumpolar structure and distribution of the
Antarctic Circumpolar Current fronts: 1. Mean circumpolar paths, J. Geophys.
Res., 114, C11018, 10.1029/2008JC005108, 2009.
Speer, K., Rintoul, S. R., and Sloyan, B.: The Diabatic Deacon Cell, J. Phys.
Oceanogr., 30, 3212–3222, 2000.Speich, S., Blanke, B., de Vries, P., Drijfhout, S., Döös, K.,
Ganachaud, A., and Marsh, R.: Tasman leakage: A new route in the global ocean
conveyor belt, Geophys. Res. Lett., 29, 1416, 10.1029/2001GL014586,
2002.Talley, L. D.: Closure of the global overturning circulation through the
Indian Pacific, and Southern Oceans: Schematics and transports, Oceanography,
26, 80–97, 10.5670/oceanog.2013.07, 2013.Thacker, W. C.: Regression-based estimates of the rate of accumulation of
anthropogenic CO2 in the ocean: A fresh look, Mar. Chem., 132–133,
44–55, 10.1016/j.marchem.2012.02.004, 2012.Thompson, D. W. J. and Solomon, S.: Interpretation of Recent Southern
Hemisphere Climate Change, Science, 296, 895–899,
10.1126/science.1069270, 2002.
Tomczak, M.: A multi-parameter extension of temperature/salinity diagram
techniques for the analysis of non-isopycnal mixing, Prog. Oceanogr., 10, 147–171, 1981.Tréguer, P. J.: The Southern Ocean silica cycle, Comptes Rendus
Geoscience, 346, 279–286, 10.1016/j.crte.2014.07.003, 2014.
Trull, T., Rintoul, S. R., Hadfield, M., and Abraham, E. R.: Circulation and
seasonal evolution of polar waters south of Australia: Implications for iron
fertilization of the Southern Ocean, Deep-Sea Res. Pt. II, 48, 2439–2466,
2001a.
Trull, T. W., Bray, S. G., Manganini, S. J., Honjo, S., and Frangois, R.:
Moored sediment trap measurements of carbon export in the Subantarctic and
Polar Frontal Zones of the Southern Ocean, south of Australia, J. Geophys.
Res., 106, 31489–31509, 2001b.
Uppstrom, L. R.: The boron/chloronity ratio of deep-sea water from the
Pacific Ocean, Deep-Sea Res., 21, 161–162, 1974.van Heuven, S., Pierrot, D., Rae, J. W. B., Lewis, E., and Wallace, D. W. R.:
MATLAB Program Developed for CO2 System Calculations, ORNL/CDIAC-105b,
Carbon Dioxide Information Analysis Center, Oak Ridge National Laboratory,
U.S. Department of Energy, Oak Ridge, Tennessee,
10.3334/CDIAC/otg.CO2SYS_MATLAB_v1.1, 2011.van Heuven, S., Hoppema, M., Jones, E. M., and de Baar, H. J. W.: Rapid
invasion of anthropogenic CO2 into the deep circulation of theWeddell
Gyre, Philos. T. R. Soc. A, 372, 20130056, 10.1098/rsta.2013.0056, 2014.van Wijk, E. M. and Rintoul, S. R.: Freshening drives contraction of
Antarctic Bottom Water in the Australian Antarctic Basin, Geophys. Res.
Lett., 41, 1657–1664, 10.1002/2013GL058921, 2014.Vázquez-Rodríguez, M., Padín, X. A., Pardo, P. C., Ríos,
A. F., and Pérez, F. F.: The subsurface layer reference to calculate
preformed alkalinity and air-sea CO2 disequilibrium in the Atlantic
Ocean, J. Marine Syst., 94, 52–63, 10.1016/j.jmarsys.2011.10.008, 2012.Verdy, A., Dutkiewicz, S., Follows, M. J., Marshall, J., and Czaja, A.:
Carbon dioxide and oxygen fluxes in the Southern Ocean: Mechanisms of
Interannual variability, Global Biogeochem. Cy., 21, GB2020,
10.1029/2006GB002916, 2007.Wassmann, P., Vernet, M., Mitchell, B. G., and Rey, F.: Mass sedimentation of
Phaeocystis pouchetii in the Barents Sea, Mar. Ecol.-Prog. Ser., 66,
183–195, 1990.
Waters, J. F., Millero, F. J., and Sabine, C. L.: Changes in South Pacific
anthropogenic carbon, Global Biogeochem. Cy., 25, GB4011,
10.1029/2010GB003988, 2011.Weiss, R.: Carbon dioxide in water and seawater: the solubility of a
non-ideal gas, Mar. Chem., 2, 203–215, 10.1016/0304-4203(74)90015-2,
1974.
Whitworth III, T. and Nowlin Jr., W. D.: Water masses and currents of the
Southern Ocean at the Greenwich meridian, J. Geophys. Res., 92, 6462–6476,
1987.Williams, N. L., Feely, R. A., Sabine, C. L., Dickson, A. G., Swift, J. H.,
Talley, L. D., and Russell, J. L.: Quantifying anthropogenic carbon inventory
changes in the Pacific sector of the Southern Ocean, Mar. Chem., 174,
147–160, 10.1016/j.marchem.2015.06.015, 2015.
Wong, A. P. S., Bindoff, N. L., and Church, J. A.: Large-scale freshening of
intermediate waters in the Pacific and Indian oceans, Nature, 400, 440–443,
1999.Zickfeld, K., Fyfe, J. C., Eby, M., and Weaver, A. J.: Comment on
“Saturation of the Southern Ocean CO2 Sink Due to Recent Climate
Change”, Science 319, 570, 10.1126/science.1146886, 2008.