Download The annual cycle of surface CO2 and O2 in the Ross Sea

The annual cycle of surface CO2 and O2 in the Ross
Sea: A model for gas exchange on the continental
shelves of Antarctica
Colm Sweeney1
1Lamont
Doherty Earth Observatory, Palisades, NY 10964.
Running Head: Annual cycle of surface CO2 and O2 in the Ross Sea
Draft: 10/15/01
Submitted to: Special volume on the Ross Sea of the Antarctic Research Series
Abstract
The annual cycle of NO3 + NO2 + NH4, CO2 and O2 in the surface waters of the
southwestern Ross Sea along 76.5oS is presented in this study. From the surface data and
sea ice concentrations annual sea-air fluxes of CO2 (-1.5±1.5 mol C m-2) and O2 (-3.7±3.0
mol C m-2) are calculated and confirmed by a mass balance approach which accounts for
the total flux of CO2 (0.16±0.13) and O2 (-5.2±0.2 mol C m-2) entering the Ross Sea from
off the shelf. The mass balance approach assumes that a negligible amount of carbon and
oxygen accumulates in the sediments and that all of the gas that ventilates to the
atmosphere must be replaced by fresher waters entering the Ross Sea. Based on this
study, a combination of winter sea ice cover and summer primary productivity prevent
any significant change in the CO2 inventory due to gas exchange despite the high partial
pressure of CO2 surface waters (425 atm) during the winter. Oxygen inventories in the
Ross Sea, on the other hand, are significantly increased as a result of gas exchange with
the atmosphere due to low O2 concentrations in the Ross Sea which are 90 mol kg-1
below saturation at sea surface temperatures of –1.89 C. The high flux associated with
large sea surface gradient in O2 is the source of high PO4* found in deep waters formed
along the Antarctic continental shelf.
Based on stability of wintertime CO2 concentrations and the “ice rectification”
hypothesis introduced by Yager et al. (1995), it is projected that with increases in
atmospheric pCO2 and greater seasonal ice cover, the Ross Sea will become a greater
CO2 sink with time. This analysis also supports the hypothesis that winter ice cover and
summer primary productivity at the polar front may have been an important factor in the
decrease in CO2 during the last glacial maximum.
Introduction
In the last decade the Ross Sea region has been extensively studied which has lead
to a rapid increase in our understanding of the annual variability of inorganic carbon and
nutrients (Gordon et al., 2000; Sweeney et al., 2000a; Sweeney et al., 2000b). In
particular, the Ross Sea has been noted for the extreme rates of carbon and nutrient
uptake through primary productivity of diatoms and Phaeocystis Antarctica (Smith and
Gordon, 1997; Arrigo et al., 1999; DiTullio et al., 2000; Smith et al., 2000b). Despite this
excessive biological uptake and export of nutrients and inorganic carbon at depths greater
than 200 m, there has been little effort to understand what the implications of the high
biological uptake and export are to the net air-sea exchange of CO2 and O2 in the Ross
Sea.
This study will show the annual cycle of CO2 and O2 in the surface of the Ross
Sea and the implied air-sea gas exchange that is derived from the observed saturation
level of CO2 and O2 in the surface waters and wind speed. In this analysis it is clear that
the effect of ice coverage on gas exchange in the Ross Sea is significant and can only be
constrained through a careful balance of the carbon and oxygen budgets in the Ross Sea.
To estimate of the carbon and oxygen that is entering and exiting the Ross Sea from off
the shelf, the balance of salt based on 18O of seawater measurements and an analysis of
the hydrological cycle in the Ross Sea made by Jacobs and Fairbanks (Jacobs et al.,
1985) will be used. By combining the above estimates with estimates of the annual burial
rates of organic matter in the sediments of the Ross Sea, the carbon and oxygen budget
can be closed in the Ross Sea.
Methods
3.1 Oceanographic cruises
The seasonal trends outlined in this study summarize observations made on
oceanographic cruises in the Ross Sea during 1996 and 1997 on the R.V.I. B Nathaniel B.
Palmer as part of the U.S. Southern Ocean Joint Global Ocean Flux Study
(JGOFS/AESOPS, or Antarctic Environment Southern Ocean Process Study, Smith et al.,
2000a). The first cruise (NBP96-04A) was designed to observe the early spring prebloom conditions and document the factors controlling the initiation of the bloom. The
second cruise (NBP97-01) was intended to investigate the CO2 and nutrient dynamics
during the austral summer, while the third cruise (NBP97-03), during autumn, was
intended to observe the heterotrophic portion of the seasonal cycle when surface water
nutrient and CO2 return to wintertime values. The fourth cruise (NBP97-08) was
conducted during the subsequent austral growing season (late spring) between the periods
bracketed by the first and second cruise and observed the progression of the bloom
through the peak of primary productivity and phytoplankton biomass accumulation. The
seasonal trends are based on only those measurements taken from a series stations
centered around 76o 30’S.
In order to identify major water masses in the Ross Sea, a larger dataset was used
which not only included data from the AESOPS cruises but also those from Research on
Ocean – Atmospheric Variability and Ecosystem Response in the Ross Sea (ROAVERRS –
NBP98-01) (Figure 1). To calculate the pre-phytoplankton bloom carbon and nutrient
concentrations of water masses off of the continental shelf, data from the Winter
Southern Ocean World Ocean Circulation Experiment (WOCE S-4P, NBP94-05) was
used.
3.2 Analytical Methods
3.2.1 Nutrients and Oxygen
In general, the methods employed for the bottle salinity, dissolved oxygen, and
nutrient analyses (PO43-, NO3-, NO2-, NH4+ and SiO(OH)3-) were similar to those
described in the JGOFS protocols (JGOFS, 1996; http://usjgofs.whoi.edu/protocols.html).
Minor differences in procedure are described in Gordon et al. (2000) and in the files that
accompany the archived data (http//usjgofs.whoi.edu). Particulars concerning nutrient
analysis during the NBP94-05 cruise are described in (Rubin et al., 1998) and do not
differ significantly from procedures used during the JGOFS cruises.
3.2.2 Carbonate System
Carbonate system measurements made on the JGOFS expeditions have been
described fully in Sweeney et al. (2000a). Briefly, total CO2 (TCO2) was measured by
coulometry in all cruises to precisions less than ±1.4 mol kg-1. All TCO2 values reported
in this study have been corrected using Certified Reference Material (http://wwwmpl.ucsd.edu/people/adickson/), whose values were determined manometrically by C.D.
Keeling.
The partial pressure of CO2 (pCO2) exerted by 500 ml seawater samples during
the NBP94-05, NBP97-01 and NBP97-08 cruise was measured by a fully automated
equilibrator-gas chromatograph system (Chipman et al., 1993). This system equilibrates
500 ml samples of seawater with a head space of gas of known volume, pressure,
temperature and initial concentration of gas. Once the head space has been fully
equilibrated with the CO2 gas in the seawater sample a fixed volume of gas was than
passed through a gas chromatograph using a hydrogen gas carrier. The hydrogen gas
carrier also enabled the catalytic conversion of the CO2 to methane and quantification of
the CO2 concentration (mole fraction – ppm) with a flame ionization detector. The CO2
concentration was calibrated using three CO2-air gas mixtures calibrated against the
World Meteorological Organization (WMO) standards (Scripps Institution of
Oceanography, unpublished data, 1994) whose values bracketed values observed in the
water samples. Because the samples were not dried, the partial pressure of CO2 was
calculated in the following way:
pCO2 @ 4.0  C meas ( ppm) Total pressure of equilibrat ion (atm) 
(1)
where pCO2 @ 4.0 is the partial pressure of CO2 of a seawater sample measured
in a constant temperature bath of 4.00 C. The pCO2 was then converted to insitu
temperatures using (Chipman et al., 1993):
 T
0.0423 4.0
pCO2 @T insitu  pCO2 @ 4.0 o e

insitu
(2)
where pCO2 @ Tinsitu is the pCO2 corrected to insitu temperatures measured by a
conductivity temperature detector (CTD). Based on duplicate measurements, the
precision of our measurements was ±0.2 atm.
On two cruises (NBP96-04B and NBP97-03) alkalinity measurements were made
instead of pCO2 measurements. The alkalinity measurements (Millero et al., 1993) were
made with a precision of ±2.0 mol kg-1. From the alkalinity, TCO2, temperature, salinity
and nutrient measurements, the partial pressure of CO2 was calculated using formulations
of Roy et al. (1993). As described in Sweeney et al. (2000a), a comparison with other
methods and constants indicate that the calculated deep water alkalinities from the Roy et
al. (1993) formulation are closest to those measured in the deep water of the Ross Sea.
3.1.3 Wind and ice data
Sea ice concentrations presented in this study are derived from multi-channel
passive microwave SSM/I (1996-present) satellite data. This data is available from the
EOS Distributed Active Archive Center (DAAC) at the National Snow and Ice Data
Center (NSIDC), University of Colorado in Boulder, Colorado (http://nsidc.org). NSIDC
provides sea-ice concentrations derived from both the Goddard Space Flight Center
(GSFC) NASA Team (Cavalieri et al., 1991; Cavalieri et al., 1997; Gloersen et al., 1992)
and Bootstrap (Comiso, 1995) algorithms. The two algorithms use different channel
combinations, reference brightness temperatures (i.e., tie-points) and weather filters, so
they therefore have different sensitivities and biases (Comiso, 1997). The data used in
this study has been derived using the Bootstrap algorithm.
3.1.4 The annual climatology
It is clear that the annual variability physical, biological and chemical is extreme
in the Ross Sea (Jacobs and Giulivi, 1998; Orsi et al., 2001). However, in an effort to put
together a climatological average for the Ross Sea. it is necessary to merge data from a
few different years together. In particular, this paper will merge the four JGOFS cruises
into one year requiring data collected November and December of 1997 to fall between
data collected in late October of 1996 and data collected from January and February of
1997. Admittedly, other physical, chemical and biological data indicate that the
conditions of the November/December cruise made in 1997 differ significantly from the
previous year. One of the most significant differences is ice cover. To adjust for this
discrepancy, the ice coverage for both years will be shown to indicate how a major
physical constraint on the biological draw down or production of CO2 and O2 may have
affected the seasonal cycle. In the end, it should be pointed out that the relative timing of
the onset of the bloom makes little difference to the overall flux of CO2 and O2 in the
Ross Sea.
Results
4.1 Seasonal cycle of physical parameters in the Ross Sea
The seasonal cycle of temperature and salinity in Ross Sea is unique compared to
many other areas of the world oceans because while ice melting and freezing drive
dramatic changes in surface salinity, surface temperature does not vary more than 2 C
throughout the whole season cycle (Figure 2a-b). The combination of small seasonal
changes in temperature and large seasonal changes in salinity makes salinity changes and,
by inference, ice melt and freezing, the primary drivers of seasonal stratification in the
Ross Sea. Figure 2c shows a rapid increase in the mixed layer depth along 76.5oS in the
southwestern Ross Sea from late winter values of ~400 m (depth representing a 0.05 kg
m-3 increase in density from the surface, http://usjgofs.whoi.edu/) with the decrease in ice
concentration. By mid-January the mixed layer depths along 76.5oS are 40 m. With the
onset of ice in late March the pycnocline at the surface has been eroded concurrently with
an increase in salinity and sea ice concentration to depths of ~200 m.
4.2 Annual cycle of surface CO2, O2 and nutrients in the Ross Sea
4.2.1 Annual nitrogen and carbon cycle
From deep water measurements of pCO2 made in the late austral winter of the
1996, it is estimated that the surface pCO2 is 425 atm at temperatures of -1.89 C before
the phytoplankton bloom has started in the southwestern Ross Sea (Figure 3a). By the
end of October, a decrease in pCO2 of almost 25 atm and little change in temperature
indicate a draw down in CO2 through primary production despite heavy sea ice cover.
This draw down in CO2 accelerates as the sea ice concentration decreases. At the time of
the maximum draw down in CO2 in late January, pCO2 values are as low as 130 atm
along 76.5oS. Low values in pCO2 in the upper 20 m of the water column correspond to
TCO2 values of ~2075 mol kg-1 which are ~150 mol kg-1 below the pre- phytoplankton
bloom values of 2233 mol kg-1 (Figure 3b). By the time winter sea ice concentration
have reached a winter average of 80%, the pCO2 and TCO2 concentrations appear to be
returning to winter values. Measurements in late March indicate that mean TCO2 and
pCO2 in the upper 20 m of the water column were 2218 mol kg-1 and 377 mol,
respectively which is still below pre-phytoplankton bloom levels.
Similarly, the total inorganic nitrogen (NO3 + NO2 + NH4, TIN) decreases rapidly
due to the onset of a phytoplankton bloom starting at the end of October. This draw down
in TIN from values of 30 mol kg-1 continues through late January resulting in TIN
values as low as 5 mol kg-1 in the upper 20 m of the water column by the end of
January. As with the pCO2 and TCO2, surface TIN, values do not return to prephytoplankton bloom values by late March despite the fact that phytoplankton growth has
stopped (Smith et al., 2000b) and the mixed layer extends to depths of 200 m (Figure 3c)
– implying significant overturning of deep CO2 saturated waters.
To illustrate how the surface concentration of TCO2 would evolve from prephytoplankton bloom levels without gas exchange, the change in TIN and the Redfield
ratio based on Takahashi et al. (1985) are used (solid lines in Figure 3b). Using the trend
line from Figure 3c and the Redfield ratio for TCO2: TIN derived along deep ocean
isopycnals (116/16, Takahashi et al., 1985) an estimated concentration of TCO2 without
the effect of gas exchange (TCO2NOGAS) in the surface waters along 76.5oS is derived
using the following:
TCO2NOGAS=TIN(116/7)+TCO2DW
(3)
where TCO2DW pre-phytoplankton bloom value of TCO2 (2233 mol kg-1) and
TIN is the change in surface TIN concentration from pre-phytoplankton bloom
conditions. This estimate does not include the effects of gas exchange and assumes: 1)
that the biological uptake and remineralization of inorganic carbon and nitrogen occur at
a constant ratio and 2) that the ratio of TIN and TCO2 below the mixed layer stays
constant. By the end of March, when the mixed layer depths are ~200 m, the difference in
measured TCO2 and the TCO2NOGAS is 6.4 mol kg-1 indicating a net flux of less than 1.3
mol C m-2 into the Ross Sea for the period beginning in early October to the beginning of
March (Table 1).
4.2.2 The annual oxygen cycle
Concurrent with the dramatic draw down in CO2 and TIN, the surface oxygen
increases from pre-phytoplankton levels far below saturation (~280 mol kg-1, 100 mol
kg-1 below saturation) to values greater than 400 mol kg-1 through the middle of
December when there is a large data gap. Because of the small inventory of O2 in the
mixed layer and the large variability in sea-air flux in the Ross Sea, it is difficult to
estimate how supersaturated O2 concentrations became as a result of the biological
production in the period when no measurements have been made. In the same way that
the change in TCO2 concentration with no gas exchange is estimated over the growing
season, the change in oxygen concentration in the surface waters can be estimated. The
change in oxygen concentration with no gas exchange (O2NOGAS) is estimated using the
change in the TIN concentration and the Redfield ratio for O2:TIN ratio of 170/16
(Anderson and Sarmiento, 1994):
O2NOGAS=TIN(170/16)+O2DW
(4)
where O2DW pre-phytoplankton bloom value of O2 (280 mol kg-1). From the
change in TIN concentrations it is clear that without gas exchange O2 values could be as
high as 150 mmol kg-1 over saturation in late December (solid line Figure 3d). With gas
exchange and moderate gas exchange fluxes, it is likely that the oxygen concentration
continues to rise with the increases in chlorophyll concentrations which are observed in
late December (indicating increases in primary productivity, Arrigo et al., 2000). Based
on the relative saturation of oxygen in the surface waters and the low average monthly
wind speeds in late December and early January, it is likely that the O2 values will reach
450 mol kg-1. The peak in O2 rapidly lowers with decreases in primary productivity, and
increases in wind speed – effectively shutting down the source of new O2 and decreasing
the residence time of O2 in the mixed layer, respectively. Although there is another data
gap between the middle of February and late March, there is a significant downward
trend in oxygen indicating that oxygen fell below saturation by the end of February due
to remineralization of organic matter and overturning of deep oxygen depleted waters in
the Ross Sea.
4.2.3 The significance of 76.5oS in southwestern Ross Sea
It is clear from other studies of the Ross Sea that the biological draw down of CO2
and TIN and production of oxygen observed along the 76.5oS is not necessarily
representative of the average in the Ross Sea. Indeed, Sweeney et al. (2000a) point out
that the areal average net community production (NCP) over the whole Ross Sea (defined
by the area inside of 1000 m isopleth) is ~65% of the average NCP along the 76.5oS (7.5
mol CO2 m-2) by the end of January. Because sea ice is slow to disappear in other areas of
the Ross Sea, the draw down in nutrients and measured primary productivity is
significantly delayed and reduced due to the shorter interval of sea ice-free conditions
throughout most of the Ross Sea (Arrigo et al., 2000). Thus, while the seasonal trends in
the surface concentrations of nutrients, carbon and oxygen along 76.5oS are useful for
illustrating the relationship between ice concentration and chemical composition of the
surface waters, the nutrient and carbon draw down at this location should be considered
an end-member in the Ross Sea.
4.3 Annual air-sea fluxes of CO2 and O2
While fluxes of TCO2 and O2 can be estimated by using TIN over the bloom
period, these estimates makes many assumptions about the biological uptake and
remineralization ratios as well as the pre-phytoplankton bloom ratios of TCO2, O2 and
TIN in water masses that may move into the study area over the course of the bloom. It is
important, therefore, to look at other methods for estimating gas exchange fluxes. One
such method is the direct calculation of the gas transfer coefficient across the sea surface
using the relative saturation of CO2 and O2. The net sea-air flux of CO2 and O2 is
calculated using:
F = kC w - sC a 
(5)
where s is the inverse of Henry’s Law Constant (the ratio of concentration of gas
in air to the concentration of gas in water) or the solubility coefficient of the gas in water
and Cw and Ca are the concentration of gas in seawater and the overlying air, respectively.
The solubility for CO2 and O2 were calculated as a function of temperature and salinity
(Weiss, 1970;Weiss, 1974), k (cm hr-1) is gas transfer velocity which is taken to be a
function of wind speed (Wanninkhof, 1992). Local wind speeds have been compiled by
National Center for Environmental Prediction – National Center for Atmospheric
Research (NCEP-NCAR) (Lamont Climate Group Library –
http://rainbow.ldeo.columbia.edu). To calculate the gas transfer coefficient for average
monthly wind speeds (uav) the following is used:
k = 0.39 uav2 (Sc/660)-0.5
(6)
where k (cm hr-1) is the gas transfer velocity and Sc is the Schmidt number, which
is a dimensionless function of temperature and salinity. Although this relationship has
been found to be consistent with the direct measurements obtained during the recent Gas
Ex-98 field study using a eddy-correlation method (Wanninkhof and McGillis, 1999),
errors resulting from the non-linear relationship between wind speed and gas transfer
velocity can lead to underestimates of average transfer velocities in regions with highly
variable wind speeds. Because of highly variable wind speeds in both space and time, the
transfer velocity represented by Eq. (6) in the Southern Ocean may be underestimated by
as much as 30% (Boutin and Etcheto, 1991).
Table 1. Sea-air flux along the 76.5oS in the Southwestern Ross Sea. Fluxes are
calculated assuming no sea ice cover (without ice), with sea ice (with ice) and
based on the difference in measured and estimated oxygen and CO2 using the
change in TIN (TIN and equation 3 and 4). Flux is expressed as mol m-2 for the
time interval indicated and positive values represent a flux out of the Ross Sea.
Gas
Annual Flux
CO2 (without ice)
CO2 (with ice)
CO2 (TIN)
O2 (without ice)
O2 (with ice)
O2 (TIN)
-0.45±1.0
-1.5±1.0
-47.3±3.0
-3.7±3.0
Bloom period flux
(Mid Oct. –Mar.)
-2.0±0.45
-1.8±0.45
-1.3±0.8
-3.4±1.5
4.6±1.5
2.7±1.0
Winter Gas flux
(Mar. –Mid Oct.)
1.55±0.7
0.1±0.6
-44±1.5
-8.3±1.5
4.2.1 CO2 sea-air fluxes
For CO2 the large difference between atmospheric and surface water pCO2
(pCO2) results in a large flux of CO2 into the Ross Sea (1.8 mol m-2) from Mid October
to Late March when the Ross Sea is open (Figure 3a and 4a). This large flux into the
Ross Sea is driven by the biological draw down in CO2 from late October to the middle
of February. While it is unclear how sea ice will effect gas exchange, assuming that the
gas exchange is linearly proportional to surface concentration of sea ice would result in a
dramatic reduction of gas exchange after the beginning of March until the middle of
October. Because of the heavy sea ice cover, it is unlikely that the CO2 supersaturated
surface waters in the Ross Sea are well ventilated in the winter.
Yager et al. (1995) refer to the reduction in the sea-air exchange of CO2 during
winter due to sea ice as a “seasonal rectification”. In their model they suggest that sea ice
not only impedes gas exchange during the winter but also provides a habitat for ice-algae
to grow at the end of the winter. The ice algae and phytoplankton which take advantage
of the shallow mixed layers left by melting ice both act to significantly draw down
CO2(aq) concentrations in the water through photosynthesis. Thus, further inhibiting
ventilation of the CO2 rich winter waters as the sea-ice melts. Because the sea-air flux of
CO2 is small (average of 0-25 mmol m-2 day-1) compared to the large deficit (~6000
mmol m-2, assuming TCO2 concentration are 150 mmol kg-1 below saturation and 40 m
mixed layer depths), the period with no sea-ice is not long enough for the gas flux to
return the CO2 concentrations back to equilibrium with the atmospheric CO2
concentrations without the additional help of overturning of the deep waters. In this way
the biological draw down and sea ice act to suppress ventilation of highly super-saturated
waters with respect to CO2.
4.2. O2 sea-air fluxes
A similar cycle is seen in the annual flux of oxygen (Figure 4b). Because of the
highly super saturated surface waters in the early part of the bloom period (Figure 3d),
there is a large flux of O2 out of the Ross Sea. This degassing of O2 from the surface
waters continues through the middle of February when the mixed layer depth decrease
indicates that overturning and primary productivity have reached their minimums
(Gordon et al., 2000; Smith et al., 2000b). Accounting for sea ice in the daily O2 fluxes
reduced the annual sea-air flux from –47 to –4 mol m-2. It is also evident from the quick
rise in oxygen above saturation levels (Figure 3d) that biological production of oxygen
prevented the ventilation of O2 depleted waters as sea ice concentration decreased in the
early austral spring. Unlike CO2 the sea-air flux of oxygen is large (0-200 mmol m-2)
while the deficit of oxygen is small (~3000 mmol m-2, assuming the O2 concentration is
75 mmol kg-1 above saturation and 40 m mixed layer depths). Thus, while daily flux of
CO2 into the Ross Sea is too small to bring the CO2 back to equilibrium in the time lag
between the heavy net community production and the onset of sea ice at the end of the
austral summer, the daily flux of O2 is sufficient. With a large flux of O2 out of the Ross
Sea and deepening mixed layers, the O2 concentrations quickly become under-saturated
and provide a gradient for a strong flux of O2 into the Ross Sea both before and after the
onset sea ice cover in the Ross Sea. In this way the oxygen cycle is similar to the CO2
cycle in the Ross Sea because the sea ice “rectifies” the winter flux of oxygen into the
Ross Sea; however, biological productivity is not as effective at preventing ventilation of
O2 under-saturated deep waters. It is likely that the impact of primary production on the
sea-air flux of oxygen is even less pronounced throughout the rest of the Ross Sea where
primary production is reduced significantly due to a possible lack of micro-nutrients and
persistent sea-ice cover (Martin et al., 1990; Arrigo et al., 1998; Fitzwater et al., 2000).
The rapid equilibration time of O2 and the highly under-saturated waters observed
in the late winter (~80 mol kg-1 below saturation) at the surface are essential
mechanisms for providing excess oxygen to the deep waters formed on the shelf regions
of the Antarctic. Broecker et al. (1986) exhibits this phenomena in the distribution of
PO4* in the world oceans. PO4* illustrates the ratios of oxygen concentration ([O2]) to
phosphate ion concentration ([PO4]) in the following relationship:
PO4*=[PO4] + [O2]/175 - 1.95
(7)
The distribution of PO4* indicates that deep water in the North Atlantic may have
values as low as 0.73 mol kg-1 while deep water formed around the shelves of the
Antarctic have values as high as 1.95 mol kg-1. Assuming that 175 moles of O2 are used
up for every phosphate ion produced in biological remineralization (and the reverse for
biological production), variations in PO4* precludes the effect biology on the oxygen to
phosphate ratios. For this reason waters with a high PO4* found on the continental shelves
of the Antarctic are waters where O2 has been added through non – biological process
such as sea-air gas exchange.
4.4 Estimates of CO2 and O2 due to deep water formation
While estimates of the sea-air fluxes of CO2 and O2 based on direct measurements
of surface water concentrations and wind speed provide an initial estimate of the sea-air
flux, the high meso-scale variability and temporal gaps in data collection add large
amount of uncertainty to the estimates. To confirm the estimates made from surface
measurements in the southwestern Ross Sea, an attempt to approach the problem from the
bottom up is made by estimating the sea-air flux based on the difference in CO2 and O2
being buried and exported on and off the shelf. This problem has been somewhat
simplified by a 14C analysis done which shows negligible accumulation of organic carbon
(0.012 mol C m-2 yr-1) in sediment cores west of 175 W in the Ross Sea (DeMaster et al.,
1996). This would imply that if there is a flux of O2 or CO2 into or out of the Ross Sea
through sea-air gas exchange it would be transported off the shelf to insure a constant
annual mean concentration of both gases.
Table 2 Average values of physical and chemical properties for water masses found in
the Ross Sea. Temperature, salinity and 18O are taken from Jacobs et al. (1985).
Nutrient and CO2 data for High Salinity Shelf Water (HSSW, S>34.6 and T<1.5C), Low Salinity Shelf Water (LSSW, S<34.6, T>-1.89C and T<-1.5C), Deep
Ice Shelf Water (DISW, S>34.6 and T<-1.89C), Shallow Ice Shelf Water (SISW,
DISW nutrients have been normalized by salinity) and Modified Circumpolar
Deep Water (MCDW, T>-1.5C, SiO4>78 mol kg-1) taken from JGOFS/AESOPS
data during October, 1996; February, 1997; April, 1997 and NovemberDecember, 1997. Nutrient and CO2 data for Circumpolar Deep Water (CDW,
Maximum temperatures below 150 m), Winter Water (Tmin, Minimum T in
upper 100 m) and Antarctic Surface Water (AASW, values based on salinity
verses concentrations trend), taken during the austral winter (September-October)
of 1994 in the Ross Gyre between 175oW and 140 oW (NBP94-5).
Water
Mass
CDW
AASW
WW
HSSW
LSSW
MCDW
SISW
DISW
Temp
(C)
1.17±0.25
-0.96±0.57
-1.64±0.18
-1.91±0.02
-1.56±0.21
-0.84±0.33
-2.04±0.04
-2.03±0.08
Salinity
34.70±0.02
34.11±0.18
34.30±0.05
34.84±0.05
34.53±0.04
34.54±0.02
34.36±0.02
34.68±0.04
18
O
TCO2
TIN
(o%)
(mol kg-1) (mol kg-1)
-0.07±0.05 2255±1
32.3±0.1
-0.31±0.08 2193±2
27.7±0.7
-0.44±0.15 2218±2
30.5±0.7
-0.42±0.05 2256±1
30.8±0.1
-0.45±0.04 2232±1
30.0±0.1
-0.34±0.09 2240±1
31.2±0.1
-0.58±0.04 2227±1
30.6±0.1
-0.54±0.06 2247±1
30.9±0.1
PO4
SiO4
Oxygen
(mol kg-1) (mol kg-1) (mol kg-1)
2.22±0.01
93.5±1.2
183±1
1.96±0.05
55.5±5.2
300±6
2.09±0.05
73.0±5.3
255±6
2.16±0.01
78.2±0.2
285±1
2.10±0.01
76.4±0.2
283±1
2.18±0.01
86.2±1.0
247±3
2.15±0.01
77.1±0.1
281±1
2.17±0.01
77.8±0.1
284±1
4.4.1 Background
In the most simplistic picture, the Ross Sea may be defined as a region which has
a total area of ~940,000 km2 bounded by the 1000 m isopleth to the northeast (~73°S)
between Cape Adare (170°E) and Cape Colbeck (158°W) (Figure 1). Almost 500,000
km-2 of the Ross Sea is covered by the Ross Ice Shelf which extends to a depth of greater
than 200 m and covers the regions to south of 78°S. To the north and east of the shelf
slope break is an area called the Ross Gyre. The Ross Gyre extends north to the Antarctic
Circumpolar current (ACC) from 160°E to 140°W. In this sector of the Southern Ocean
the southern limit of the ACC lies between 65°S and 60°S (Orsi et al., 1995). The waters
of the Ross Gyre are contained between the ACC and the Ross Sea and flow clockwise
(cyclonic) driven by westward winds located south of the ACC. At the southern extreme
of the Ross Gyre the clockwise circulation translates into a westward flow. A portion of
this westward flowing water moves onto the continental shelf of the Ross Sea while the
bulk of it moves along the continental slope (Gill, 1973; Ainley and Jacobs, 1981) and
mixes with HSSW, LSSW and ISW coming off the shelf at depths greater than 500 m.
From this simplified picture other investigators have assumed that only half of the
AABW formed on the continental slope of the Ross Sea is HSSW, LSSW and ISW
(Jacobs et al., 1985). The rest of the water making up AABW is Lower Circumpolar
Deep Water (LCDW) which originates to the north at a depth below 2000 m in the ACC
but reaches 200 m depth as it flows poleward (Locarnini, 1994). Because it is partially
formed from warmer intermediate waters to the north of the ACC, the LCDW plays an
important role supplying the heat needed to maintain the polynyas at the shelf break of
the Ross Sea through most of the winter (Jacobs and Comiso, 1989). Once the LCDW
has moved onto the shelf it is observed to have temperatures and salinities lower than
those observed off the shelf due to mixing with fresher and colder surface waters and was
first distinguished as Modified Circumpolar Deep Water (MCDW, Jacobs et al., 1970,
Figure 5a and 6, Table 2). It is the LCDW and colder Antarctic Surface Waters (AASW)
and Winter Water (also known as temperature minimum water, Tmin) which replace the
colder and more saline HSSW, LSSW and ISW water leaving the Ross Sea. To keep this
circulation going, it is important to not only extract heat but also add salt to the waters
going off the shelf. While heat can easily be extracted by the relatively high winds, weak
pycnoclines and deep mixed layers, the addition of salinity requires a more complicated
process.
One of the primary mechanisms suggested for generating the salinity needed to
create HSSW, LSSW and ISW in the Ross Sea is through brine rejection (Gill, 1973;
Killworth, 1974; Jacobs et al., 1985; Zwally et al., 1985; Broecker et al., 1986; Locarnini,
1994; Toggweiler and Samuels, 1995). Through this mechanism AASW in the Ross Sea
are frozen and pushed northwest off the shelf by prevailing currents and winds (Jacobs
and Comiso, 1989). The brine left behind by the exported sea ice is added to the water of
the Ross Sea to form LSSW and HSSW. From the relatively high salinity, low 18O and
supercooled temperature of ISW (Table 2) it is likely Deep Ice Shelf Water (DISW) is
formed from HSSW that has been subducted under the Ross Ice Shelf to a depth
sufficient to allow the temperatures to decrease below the surface freezing point due to
basal melting of the Ross Ice Shelf.
4.4.2 The Ross Sea salinity balance
Using a mass balance approach and estimates of salinity, temperature and 18O
specified by Jacobs et al., (Jacobs et al., 1985) (Figure 5a-b, 6 and Table 2) a net flux of
45 cm yr-1 of ice are exported off the shelf which translates into the formation of 1.4
Sof deep water on the continental shelf defining the Ross Sea. This calculation is based
on the following assumptions made by Jacobs et al. (1985) including proportional
mixture of waters flowing on and off the shelf (Table 3):
Assumption #1: The average 18O of water in Ross Sea stays constant (at steady
state) despite a large input of 18O depleted melt water and precipitation. Based on Jacobs
et al. (1985), glacial melt water, with a 18O of -36o% is added to the average depth of the
Ross Sea at the rate of about 0.36 m yr-1 thus decreasing the 18O by 0.025o% yr-1. In
addition, the estimated precipitation and the freezing of sea ice will decrease the 18O by
0.005o% yr-1 and 0.002o% yr-1, respectively. The cumulative decrease in the 18O
(0.0332o% yr-1 *3.2 x 105 km3) will be countered by an equivalent increase in the 18O
due to the difference in the 18O coming on and going off the shelf (0.24o%, Figure 5b, 6
and Table 2). Thus,
Annual 18O decrease (precip., ice melt, etc.) = (0.0332 o%)(3.2 x 1014 m3) =1.06 x 1013 o% m3 yr-1
18O flux increase from deep water formation = (0.24 o%)(Deep water formation rate)
Assuming no inter annual variability in 18O (steady state):
18O flux increase from deep water formation =Annual 18O decrease (precip., ice melt, etc.)
Deep water formation rate = 4.39 1013 m3 yr-1 (or 1.4 S)
Table 3 Average values of water flowing onto the continental shelf of the Ross Sea based
on water mass recipes from (Jacobs et al., 1985).
Temp
(C)
Inflow Water* -0.065
Salinity
34.453
18
O
TCO2
TIN
PO4
SiO4
Oxygen
(o%) (mol kg-1) (mol kg-1) (mol kg-1) (mol kg-1) (mol kg-1)
-0.2225
2230±1
30.7±0.2
2.12±0.02
78.8±1.9
230±2
34.54
-0.42
2232±1
30.3±0.2
2.12±0.01
76.5±1.1
276±2
Average Shelf -1.45
Water**
-1.689 34.5876 -0.465
2237±1
30.2±0.1
2.12±0.00
76.8±0.2
283±1
Outflow
Water***
*Water coming onto shelf=50% CDW + 25% AASW + 25%WW
**Average Shelf water = 15% AASW +20% MCDW +35% LSSW + 20% HSSW + 5% DISW+ 5% SISW
***Average going off = 70% LSSW + 10% HSSW + 20% DISW
Assumption #2: Ice contains ~4100 g salt m-3 (Jacobs et al., 1985) which is about
10% of normal seawater. Using this assumption, the quantity of ice that is be exported off
the shelf can be calculated based on the salinity difference between water coming onto
and off of the shelf (0.135) given the formation rate of deep water.
Salt flux onto the shelf = Salt flux off the shelf
Salt flux off of the shelf = (134.6 g m-3)(4.39 1013 m3 yr-1) = 5.91 x 1015 g salt yr-1
Salt flux onto the shelf = (Depth of net ice export)(Area of RS)(Salt released by ice)
Depth of net ice export = 5.91 x 1015 g salt yr-1/30000 g salt m-3/4.41 x 1011 m2 = 0.45 m yr-1
Assumption #3: Using a mass balance approach Jacobs et al. (1985) suggest that
precipitation and glacial melt add 0.52 m yr-1 of freshwater to Ross Sea per year and only
half of the ice formed is exported off the continental shelf. From this, the total amount of
ice formed each year can be calculated.
Salt flux to balance fresh water input = (Salt content of Ross Sea)(Depth of freshwater)
1.8 x 104 g salt m-2 yr-1 = (34540 g salt m-3)(.52 m yr-1)
Ice export to balance fresh water = (Salt flux to balance fresh water input)/(Salt released by ice)
0.60 m yr-1 = 1.8 x 104 g salt m-2 yr-1/(34110 g salt m-3-4100 g salt m-3)
Annual ice accumulation = 2[(Net ice export) + (Ice export to balance fresh water)]
= 2(0.45 m yr-1 + 0.60 m yr-1) = 2.10 m yr-1
Assumption #4: Along with salt rejection there will also be an enrichment in
nutrients and CO2 due to the brine left behind as a result of exported ice. Thus, the total
flux of CO2, nutrients and oxygen onto the shelf can be calculated as sum of net flux,
incoming minus outgoing flux of nutrients, and the flux of nutrients derived from brine
left by the sea ice exported off the shelf (brine flux).
Table 4: Fluxes of Nutrients and CO2 coming into the Ross Sea. The brine flux into the
Ross Sea is calculated assuming that the source of ice is AASW and that the same
fraction of nutrient and CO2 will be removed from the AASW as salt (the salinity
of the ice is assumed to be 4.1). The net flux into the Ross Sea is based on the
difference in nutrient and CO2 flux coming into and going out of the Ross Sea
(Table 1.2). Uncertainties reflect the propagation of errors from estimating
average nutrient and CO2 concentrations of incoming and outgoing water.
Brine Flux:
Nutrient
(109 . mol yr-1)
TCO2
386.5±0.2
NO3+ NO2+NH4
5.2±0.0
PO4
0.37±0.00
SiO4
13.3±0.2
O2
47.9±0.3
Net Flux:
(109 . mol yr-1)
-315.2±58.6
19.9±8.7
0.05±0.83
88.1±83.4
-2325±82
Total Flux
Total Flux
(109 . mol yr-1) (mol m-2 yr-1)
71.3±58.6
0.16±0.13
25.1±8.7
0.06±0.02
0.001±0.002
0.42±0.83
101.4±83.4
0.2±0.2
-2276.9±81.8
-5.2±0.2
Figure 7a-c shows the salinity normalized concentrations of TCO2, O2 and SiO4 as
a function of salinity. By normalizing the each water mass concentration to the mean
salinity of the Ross Sea one can see the effect of ice melt and precipitation or freezing
and evaporation have on each water masses. While not as rigorous as the calculations
described above it helps to illustrate difference between the concentrations of all the
water masses in addition to accounting for concentration enrichment due to brine
rejection. It is clear from the error bars that while oxygen in the waters leaving the Ross
Sea are enriched by 50 mol kg-1 and accounts for a total flux off the Ross Sea shelf of
5.2 mol O2 m-2 yr-1, the total CO2 and silicate in the waters leaving the Ross Sea is not
significantly depleted.
The fact that only 0.2±0.2 mol of SiO4 m-2 is added to the Ross Sea annually is an
important confirmation of the recipes for incoming and outgoing waters suggested by
Jacobs et al. (1985) because it confirms estimates which suggested that 0.11 mol SiO4 m-2
yr-1 accumulates on the shelf every year (DeMaster et al., 1996). If the amount of CDW
entering the Ross Sea was 60% of incoming water masses instead of 50% as Jacobs et al.
(1985) suggest the accumulation SiO4 would increase to 0.5±0.2 mol m-2yr-1 – an
estimate significantly different from DeMaster et al. (1996).
The salinity balance and the recipes of incoming and outgoing water masses of
Jacobs et al. (1985) confirms that there is a significant sea-air flux of O2 into the Ross Sea
but does not confirm the air-sea flux of CO2 into the Ross Sea calculated from the annual
climatology of sea surface pCO2. The lack of convergence in estimates may be explained
in two ways. First, it is important to acknowledge that in addition to the many assumption
made to for the salinity balance approach no accounting has been done of particulate
matter that may be exported out of the highly productive Ross Sea. This is especially
relevant in light of the disparity between estimates of exported particulate carbon in the
Ross Sea. Sweeney et al. (2000b) suggested that 3.8 mol m-2 of organic carbon was
exported past 200 m based on the difference between NCP and total organic carbon in the
upper 200 meters while only 0.29 mol C m-2 was caught in the sediment traps at 200 m
(Collier et al., 2000) during that same period. If lateral advection of this material were
responsible for some of the disparity in export estimates, accounting for particulate
organic matter might significantly decrease the net flux of carbon into the Ross Sea from
waters off the shelf. If enough organic matter was exported off the shelf the deficit in the
mean annual total CO2 concentration which would have to be made up with flux of
atmospheric CO2 into the Ross Sea. Presumably an export of organic material would also
imply a greater flux of O2 through gas exchange into the Ross Sea as well. It is clear,
based on the need to maintain a net flux of nitrate and phosphate into the Ross Sea near
zero (assuming negligible sedimentary accumulation), that the full 3.5 mol C m-2
discrepancy between estimates of Sweeney et al. (2000b) and Collier et al. (2000) can be
accounted for. However, if 1 mol C m-2 yr-1 was exported off the continental shelf of the
Ross Sea as particulate organic matter this would also imply a net export 0.14 mol N m-2
yr-1 and 0.0086 mol P m-2 yr-1 as organic matter using Redfield ratio (Anderson and
Sarmiento, 1994).
The second reason for the disparity between the two estimates of the CO2 sea-air
flux is the anomalously high productivity found in the southwestern Ross Sea along
76.5oS. While it is clear that phytoplankton growth does occur throughout the Ross Sea,
Sweeney et al., (2000a) show that even in late January during the height of the
phytoplankton bloom that some areas of the Ross Sea surface CO2(aq) are at equilibrium
with atmospheric CO2. This would imply that while the phytoplankton bloom observed
along 76.5oS in the southwestern Ross Sea provided a gradient which causes a net flux of
CO2 into the Ross Sea, the phytoplankton bloom did not draw down CO2 to the same
extend throughout the Ross Sea. As pointed out previously, Sweeney et al. (2000a)
indicate that the mean NCP along 76.5oS was almost 35% higher than the mean for the
entire Ross Sea. It is, therefore, likely that the estimated sea-air flux base on surface
measurements along 76.5oS is bias by the excessive biological draw down in CO2 that
was observed in that region.
Discussion
Although the Ross Sea may be considered somewhat anomalous because of the
high primary productivity, it may provide a very good model for understanding the major
mechanisms that control gas exchange in high latitudes.
5.1 Implications for gas exchange on the continental shelves of Antarctic
It is clear from both the analysis of surface water measurements and analysis of
incoming and outgoing water masses from the Ross Sea that sea ice plays a major role in
the prevention of gas exchange as observed by Yager et al. (1995) and Gibson and Trull
(Gibson and Trull, 1999). The major piece of evidence for the lack of gas flux during
wintertime when the sea ice cover is extensive can be found in the oxygen concentrations
throughout the winter in the Ross Sea (Figure 3d). Despite the large daily fluxes of
oxygen that are estimated during the austral winter without ice cover (Figure 4), there is
no evidence from the analysis of water masses moving onto or off of the shelf that would
support O2 fluxes as high as 47 mol m-2 yr-1 which is estimated in conditions with no sea
ice. Gibson and Trull (Gibson and Trull, 1999) have made similar observation in Prydz
Bay, East Antarctica where oxygen remained under-saturated throughout the winter.
Despite the presences of winter sea ice and high biological production in the Ross
Sea (Smith and Gordon, 1997; Smith et al., 2000b) and elsewhere in the Antarctic shelf
regions, there is a large flux of oxygen into waters on the continental shelves due to the
contribution of CDW which contains waters that are extremely under-saturated in O2
coming onto the shelf. As discussed earlier, it is this large flux of O2 into the Antarctic
surface waters that provides the high end-member PO4* in the deep waters that is formed
on the shelves of Antarctica.
5.2 Implications for gas exchange in the past and future
From the interaction of sea ice and primary production and the resulting gas
exchange that is observed in the Ross Sea, it is possible to confirm “rectification”
hypothesis suggested by Yager et al. (1995). Specifically, this theory points out the
importance of the phytoplankton bloom both before the sea ice breaks up and after it has
disappeared. Because there is very little gas exchange during the winter, the CO2
saturated waters are never able to degas. It is only after CO2 has been depleted in the
surface waters by primary productivity that the sea ice retreats providing a large sink, not
source, for atmospheric CO2. Yager et al. (1995) suggest that with increased warming due
to increases in atmospheric CO2 levels and consequential increase in seasonal ice
coverage that there may be a negative feedback provided by the interaction between sea
ice and primary production.
While primary production plays a significant role in creating a sea-air gradient for
CO2 gas exchange, so does the concentration of CO2 in the atmosphere. If the rate of
increase in atmospheric CO2 were to stay constant, the atmospheric pCO2 will be equal to
that of the pre-phytoplankton bloom pCO2 of 425 atm in less than 40 years. This
increase in atmospheric pCO2 as a result of anthropogenic inputs may also provide an
important negative feedback for gas exchange in the future as the air-sea gradient
increases. Because the unventilated lower CDW is the largest source of water on the
continental shelves of Antarctic, the wintertime concentrations of TCO2 on the
continental shelves of Antarctica are unlikely to change much over the next 40 years.
However, as the atmospheric CO2 increases over the next 40 years, the sea-air gradient
will become increasingly negative and provide a greater flux of CO2 into the waters on
the continental margins of Antarctica.
It has also been suggested that the rectification hypothesis explains the large
decrease in atmospheric CO2 during glacial time (Stephens and Keeling, 2000). Stephens
and Keeling (2000) hypothesis that with the large decreases in temperature suggested in
the high latitudes during the last glacial maximum (Petit et al., 1999) that winter sea ice
coverage extended beyond the present day polar front (Crosta et al., 1998) and prevented
exchange between atmosphere and deep CO2 saturated waters. In addition, they
suggested that when the ice melted, it left well stratified and biologically productive
waters which prevented further exchange with the atmosphere. While CO2 was not well
ventilated, Stephens and Keeling (2000) found that O2 was quicker to equilibrate and was
less affected by primary productivity as it has been pointed out in this study.
Summary
The annual change in the surface concentrations of CO2 and O2 are estimated to
be more than ~150 mol kg-1 and ~160 mol kg-1, respectively, along 76.5oS in
southwestern Ross Sea. The large changes in CO2 and O2 are also accompanied by ~25
mol kg-1 changes in TIN. These changes in nutrients, CO2 and O2 are primarily driven
by a large phytoplankton bloom centered in this area of the Ross Sea polynya where sea
ice concentrations decrease quickly in the late austral winter.
The large amplitude in the annual cycle of both CO2 and O2 surface
concentrations along 76.5oS in the southwestern Ross Sea creates large sea-air gradients
which in regions without sea ice would be responsible for both gas fluxes into and out of
ocean depending on the season. Based on estimates using a linear relationship between
sea ice concentration and gas flux, the CO2 flux increases by 1.0 mol m-2 yr-1 and the O2
flux decreases by 44 mol m-2 yr-1 into the Ross Sea when sea ice concentrations are taken
into account. The annual flux of CO2 (-1.5±1.5 mol C m-2) and O2 (-3.7±3.0 mol C m-2)
calculated from measurements of surface concentrations through one annual cycle are
confirmed by a mass balance approach which accounts for the total flux of CO2
(0.16±0.13) and O2 (-5.2±0.2 mol C m-2) entering the Ross Sea from off the shelf. In
order to compare results, it is assumed that a negligible amount of carbon and oxygen
accumulate in the sediments and that all of the gas that ventilates to the atmosphere must
be supplied from fresh waters entering the Ross Sea. In comparing the two methods it is
important to realize that the mass balance approach does not take into account possible
exchange of particle matter with areas off the shelf. It also must be pointed out that data
collected along 76.5oS from which sea-air gas exchange fluxes were calculated may over
estimate the impact of biological draw down of CO2 and production O2 by as much as
35%.
The role of sea ice and primary productivity in the annual cycle of gas exchange
in the Ross Sea strongly supports the “rectification” hypothesis (Yager et al., 1995) and
provides a working model for scenarios hypothesized by Stephens and Keeling (1999). In
this scenario they suggest that a large portion of the decrease in atmospheric CO2
measured in ice cores can be accounted for by increases in sea ice coverage during the
last glacial maximum. In addition, this study supprts Yager et al. (1995) who suggest that
an increase in seasonal sea ice coverage in areas like the Ross Sea with global warming
may promote a larger draw down in CO2 during the summer months and, hence, a greater
sink for CO2 with global warming.
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Figure 1. Station locations. Hydrographic station location for cruises on the R/V
Nathaniel B. Palmer. Black dots represent cruises done for the Antarctic Environment
Southern Ocean Process Study (AESOPS – NBP96-04B, NBP97-01, NBP97-03 and
NBP97-08) and (ROAVERRS – NBP98-01). The white dots represent data from the
World Ocean Circulation Experiment (WOCE – NBP94-05).
a)
b)
c)
Figure 2. The seasonal cycle of a) temperature (C), b) salinity and c) mixed layer depth
(m) in the Ross Sea at 76.5oS between 170oW and 180o. Average values were measured in
the upper 20 m of the water column (●), a smoothed average of the measurements (--).
The light gray background line refers to the % ice coverage (right axis) for the 1996-1997
season (solid line) and 1997-1998 season (dashed line).
a)
b)
c)
d)
Figure 3. The seasonal cycle of surface a) partial pressure of CO2 (pCO2), b) total CO2
(TCO2), c) total inorganic nitrogen (TIN) and d) oxygen in the Ross Sea at 76.5oS
between 170oW and 180o. Average values were measured in the upper 20 m of the water
column (●), a smoothed average of the measurements (--). Solid line represents estimates
of surface TCO2 and O2 concentrations based on changes in surface TIN using Redfield
ratios of 116/7 and 175/7, respectively. The light gray background line refers to the % ice
coverage (right axis) for the 1996-1997 (solid line) and 1997-1998 (dashed line) seasons.
a)
b)
Figure 4. The daily sea-air flux of a) CO2 and b) oxygen in the Ross Sea at 76.5oS
between 170oW and 180o. Daily flux of gas based on the average sea ice concentration (-)
and no sea ice conditions (--) assuming the flux is proportional to the concentration of sea
ice, wind speed and sea-air gradient. The fraction of sea ice coverage used was for the
1996-1997 season.
a)
b)
Figure 5. Temperature/Salinity and 18O/Salinity relationship between water masses on (o)
and off (●) of the Ross Shelf. Incoming water (x) is 50% CDW, 25% AASW and
25%Tmin. Outgoing water (●) is 70% LSSW + 10% HSSW + 20% DISW. Error bars
represent the standard error of the mean.
Figure 6 18O and salt budget of the Ross Sea according to Jacobs et al. (1985). Based on
Jacobs et al. (1985), water masses coming into the Ross Sea (Circumpolar Deep Water
(CDW), Antarctic Surface Water (AASW) and Winter Water (WW)) have a 18O of
0.2225 and salinity of 34.453 while waters moving off of the continental shelf that
defines the Ross Sea have a 18O of 0.465 and salinity of 34.588.
a)
b)
c)
Figure 7. Salinity verses a) total CO2 (TCO2) b) oxygen and c) Silicate. Each value
represents the mean concentration of dissolved inorganic carbon, silicate and oxygen
normalized to the average salinity of the Ross Sea (34.54, =[C]*salinity/34.54). Water
masses on (o) and off (●) of the Ross Shelf have been seperated. Incoming water (x) is
50% CDW, 25% AASW and 25%Tmin. Outgoing water (●) is 70% LSSW + 10% HSSW
+ 20% DISW. Error bars represent the standard error of the mean.