GEOPHYSICAL RESEARCH LETTERS, VOL. 31, L17305, doi:10.1029/2004GL020900, 2004
Is AOU a good measure of respiration in the oceans?
T. Ito, M. J. Follows, and E. A. Boyle
Program in Atmospheres Oceans and Climate, Massachusetts Institute of Technology, Cambridge, Massachusetts, USA
Received 1 July 2004; accepted 20 August 2004; published 14 September 2004.
[1] Apparent Oxygen Utilization (AOU) is widely used to
infer respiration in the oceans by assuming that surface
oxygen concentration is close to saturation with the
overlying atmosphere. However, significant disequilibrium
of oxygen has been observed in high latitude surface
oceans where the deep waters are formed. We explicitly
calculate True Oxygen Utilization (TOU) in a global ocean
physical-biogeochemical model to evaluate the ability
of AOU to represent respiration. We find significant
differences between AOU and TOU in the deep waters,
suggesting a systematic overestimation of respiration when
inferred from AOU. The surface heat flux and the
entrainment of thermocline waters together drive the
surface undersaturation of oxygen in the regions of water
mass formation, and their influences are significantly
INDEX
enhanced by sea ice cover at high latitudes.
T ERMS : 4805 Oceanography: Biological and Chemical:
Biogeochemical cycles (1615); 4820 Oceanography: Biological
and Chemical: Gases; 4255 Oceanography: General: Numerical
modeling. Citation: Ito, T., M. J. Follows, and E. A. Boyle
(2004), Is AOU a good measure of respiration in the oceans?,
Geophys. Res. Lett., 31, L17305, doi:10.1029/2004GL020900.
1. Introduction
[2] Dissolved oxygen is generally depleted in thermocline and abyssal oceans due to the respiration of organic
material. Air-sea equilibration of oxygen is generally rapid
relative to ocean circulation timescales, and observed
surface oxygen concentrations are indeed close to saturation
in many areas of the surface oceans [e.g., Broecker and
Peng, 1982; Chester, 2000]. Depletion of oxygen relative to
saturation has been used as a measure of respiration in the
interior ocean. AOU is defined as the difference between the
saturation oxygen concentration, O2,sat and the observed
oxygen concentration, O2 as defined in equation (1).
AOU ¼ O2;sat O2
ð1Þ
where O2,sat depends on potential temperature, T, and
salinity, S. Estimates of respiration inferred from AOU
could contain errors due to several processes as pointed out
by Shiller [1981], Broecker and Takahashi [1985] and
Broecker et al. [1991], which include air-sea disequilibrium
in the regions of water mass formation, non-linearity in the
solubility of oxygen, and respiration involving denitrification. The impact of denitrification can be diagnosed from
the distribution of N* [Gruber and Sarmiento, 1997; Sabine
et al., 1999] whereas the impact of the other effects have not
been well-understood nor quantified. In this study, we
Copyright 2004 by the American Geophysical Union.
0094-8276/04/2004GL020900$05.00
explicitly determine the decoupling of AOU from respiration in a 3-dimensional physical-biogeochemical model.
[3] True Oxygen Utilization (TOU) [Broecker and Peng,
1982] is defined as the difference between the preformed
oxygen content, O2,pre, and the observed oxygen content.
TOU ¼ O2;pre O2
ð2Þ
DO2 ¼ TOU AOU ¼ O2;pre O2;sat
ð3Þ
Preformed oxygen is set at the time of subduction and is
transported into the interior ocean by physical circulation.
Thus the preformed properties typically represent the mixed
layer properties during winter seasons. Differences between
AOU and TOU can arise from the surface disequilibrium at
the time of subduction and water mass formation, and from
mixing in the interior ocean combined with the nonlinearity
in the solubility of oxygen. DO2 is exactly equal to the
degree of saturation at the surface; see equation (3). DO2 is
mapped into the interior ocean through advection and
mixing. While O2,sat can be determined from observed
temperature and salinity, it is impossible to determine DO2
from direct observation since the true preformed oxygen
content remains unknown.
2. Decoupling of AOU From TOU
[4] Here, we explicitly simulate DO2 using the MIT
ocean general circulation and biogeochemistry model
[Marshall et al., 1997a, 1997b; P. Parekh et al., Decoupling
of phosphorus and iron in the global ocean, submitted to
Global Biogeochemical Cycles, 2004, hereinafter referred to
as Parekh et al., submitted manuscript, 2004] by calculating
the global distribution of preformed oxygen, O2,pre as an
additional tracer. The model also includes explicit representation of the biogeochemical cycles of DIC, Alkalinity, O2,
PO4, DOP and Fe. Horizontal resolution is 2.8 and there are
15 vertical layers. The model is forced dynamically with
monthly climatological surface boundary conditions of wind
stress and buoyancy flux [Trenberth et al., 1989; Conkright
et al., 2002]. The model employs explicit iron chemistry
forced with modeled monthly dust deposition field
[Mahowald et al., 2003; Parekh et al., submitted manuscript,
2004]. Air-sea gas transfer is parameterized following
[Wanninkhof, 1992]. Biological productivity can be limited
by the availability of Fe or PO4. All biological processes are
assumed to occur in Redfieldian stoichiometry [Anderson
and Sarmiento, 1994]. Preformed oxygen is set equal to
simulated O2 in the surface layer, and transported as a
conservative, passive tracer in the ocean interior.
[5] Typically, estimates of oxygen utilization, and regenerated carbon and nutrients, rely on the assumption that DO2
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surface undersaturation affects the distribution of oxygen
and AOU in the interior ocean. Figure 2 shows the
difference between AOU and TOU in meridional sections
of the Atlantic (23W) and the Pacific (165W). DO2 is
generally negative in the interior ocean implying systematic
overestimation of respiration by AOU due to the undersaturation at the high latitude surface outcrop which is
ventilated into the interior. The largest difference is found
in the Atlantic sector of the Southern Ocean indicating O(1)
errors in the regional estimates of the regenerated carbon
and nutrients inferred from AOU. The large errors, originating from the surface outcrop at the high latitudes, spread
into the deep basins along isopycnals.
Figure 1. Simulated, annual-mean surface disequilibrium,
DO2. Contour interval is 10 mM. Solid line represents
undersaturation, DO2 < 0. See color version of this figure in
the HTML.
is small and AOU TOU [e.g., Broecker and Peng, 1982;
Chester, 2000]). Here we quantitatively test this assumption
in the context of the numerical model. Figure 1 shows the
simulated, annual mean distribution of surface disequilibrium
of oxygen, DO2, at steady state. It is in a qualitative
agreement with the observational estimate of [Najjar and
Keeling, 1997], and is broadly consistent with the observed
seasonal variations. Tropical and subtropical oceans are
close to saturation with slight supersaturation due to the
photosynthesis and surface heating. High latitude surface
oceans are generally undersaturated in the model. The
largest disequilibrium is found in the region of deep water
formation of the model. DO2 reaches up to 73 mM in
the Weddell Sea, broadly consistent with previous observations [Weiss et al., 1979; Schlosser et al., 1991]. In the
North Atlantic, we find an undersaturation of 18 mM in
the Labrador Sea. Surface undersaturation is also found in
the western boundary currents of subtropical gyres. The
Figure 2. DO2 is shown in the Atlantic (23W) and the
Pacific (165W) of the control run. Contour interval is 10 mM.
Dashed line represents zero and negative values. These
results lead to local errors of up to 55 mM of regenerated
carbon. See color version of this figure in the HTML.
3. What Controls the Disequilibrium of
Oxygen in the Polar Oceans?
[6] What drives the large undersaturation in polar surface
oceans? Figure 3 schematically illustrates the processes
affecting the preformed oxygen near the region of deep water
formation. Heat loss and the associated solubility increase
could drive the surface condition toward undersaturation.
Convective mixing or entrainment of intermediate and deep
waters could also enhance undersaturation by bringing
oxygen depleted waters to the surface. The magnitude of
DO2 is set by the relative timescales of these processes with
respect to the air-sea gas exchange. A large DO2 is expected
when the rate of solubility increase due to heat loss or the rate
of entrainment flux are relatively rapid compared to that of
the air-sea gas exchange. Air-sea gas exchange can be
inhibited by sea ice cover effectively reducing the rate of
air-sea gas equilibration, significantly amplifying the effects
of heat flux and entrainment.
[7] In a suite of sensitivity studies, we repeat the numerical
simulation suppressing each of the three key processes in the
model: (A; No-ICE) the inhibition of air-sea gas transfer
by sea ice, (B; No-BIO) the effect of respiration on oxygen in
the interior ocean, and (C; No-HF) the effect of heat fluxes
on the solubility of oxygen in the surface ocean, keeping
all other parameters fixed. The results of the sensitivity
Figure 3. Schematic diagram of the physical processes
driving the surface oxygen away from saturation. Heat loss
at the surface increases the solubility of oxygen, which
leads toward undersaturation. Air-sea gas transfer is
inhibited by sea ice. Thermocline and deep waters are
depleted in oxygen relative to surface waters. Convective
mixing, over a background gradient of O2 can lead to
undersaturation by bringing up the O2-depleted intermediate
and deep waters. See color version of this figure in the
HTML.
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experiments are summarized in Table 1, and Figure 4
shows the distribution of DO2 in the Southern Ocean.
[8] In experiment (A), without the effect of sea ice on airsea gas transfer, we find that the magnitude of the surface
disequilibrium is significantly reduced by 64% in the
Weddell Sea, and by 86% in the Labrador Sea, suggesting
a substantial impact of sea ice on the gas exchange rates in
the polar oceans. This effect has little impact on DO2 in the
western boundary currents.
[9] In experiment (B) where we suppress the effect of
respiration on dissolved oxygen, surface undersaturation is
reduced significantly in the high latitude Southern Ocean
(approximately 65%; see Table 1 and Figure 4b), though
in the Labrador Sea, the response is much weaker (approximately 24%). The air-sea disequilibrium is driven by the
mixing and entrainment of intermediate and deep waters
which have accumulated a large oxygen deficit due to
respiration. This effect is more pronounced in the polar
Southern Ocean where upwelling deep waters are older.
This effect has significant impacts on DO2 in the western
boundary currents, particularly in the Pacific.
[10] In experiment (C) where we suppress the temperature dependence of the solubility of oxygen, surface disequilibrium is reduced by 27% in the Weddell Sea and by
70% in the Labrador Sea (see Table 1 and Figure 4c). Heat
losses and associated solubility increase play a relatively
minor role in the high latitude Southern Ocean whereas this
effect dominates in the North Atlantic. This effect also has
impacts on DO2 in the western boundary currents.
[11] In the polar oceans, the effects of the heat loss and
the entrainment of old, deep waters together drive the
surface waters toward undersaturation, and they control
the magnitude of DO2 in the interior ocean. Combination
of these effects are largely additive, suggesting that the
simulated response of the surface disequilibrium is quasilinear in this particular model. Relative importance of these
processes varies over geographical regions in the model.
The effect of the ice cover amplifies the overall undersaturation by reducing the rate of air-sea gas transfer.
4. Discussion
[12] Surface disequilibrium of dissolved oxygen directly
affects the distribution of AOU in the interior ocean and
decouples AOU from TOU by a significant magnitude. The
decoupling is large and systematic, and its spatial pattern is
Table 1. The Summary of the Sensitivity Experimentsa
Weddell Sea
(55W, 70S)
Ross Sea
(180E, 70S)
Labrador Sea
(60W, 65N)
Kuroshio
(145E, 35N)
Gulf stream
(70W, 35N)
Control
-
Exp.(A)
No ICE
Exp.(B)
No BIO
Exp.(C)
No HF
73
(0)
51
(0)
18
(0)
21
(0)
4.8
(0)
26
(64%)
19
(63%)
2
(86%)
21
(0%)
4.9
(+2%)
25
(66%)
19
(63%)
13
(24%)
8
(60%)
3.0
(37%)
53
(27%)
40
(21%)
5
(70%)
16
(22%)
2.9
(41%)
a
Simulated DO2 in several geographical regions from each sensitivity
experiments are shown in units of mM. Numbers within the parenthesis
represent the fractional changes relative to control, defined as
(DO2/DO2(control) 1) 100 (%).
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Figure 4. Surface disequilibrium, DO2, in the Southern
Ocean from the control run and the sensitivity experiments
(A, B, C). Contour interval is 10 mM. Dash line represents
DO2 = 0. In the experiment (A), the inhibition of air-sea gas
transfer due to sea ice is suppressed. In the experiment (B),
the effect of respiration on the oxygen is suppressed. In the
experiment (C), the temperature dependence of the
solubility of oxygen is suppressed. See color version of
this figure in the HTML.
primarily oriented along isopycnals. Estimates of regenerated carbon and nutrients potentially include O(1) error
when inferred from AOU. Specific respiration rates inferred
from a time rate of change in AOU, so-called Oxygen
Utilization Rate (OUR), may be less affected by this error
[Jenkins, 1982]. The DC* method [Gruber et al., 1996;
Gruber, 1998], which is a method for estimating anthropogenic CO2 in the oceans, may also be affected by the effect
of oxygen disequilibrium. Overestimation of respiration will
result in an underestimation of preformed carbon and DC*.
However this may be partially accounted in the empirically
determined residual term, DCdiseq. More studies are needed
to access its impact on the uncertainty of anthropogenic
CO2 inventory estimates. Since the spatial distribution of
DO2 are systematic and primarily oriented along isopycnals,
it may be possible to develop a method to correct AOU
including the effect of surface disequilibrium by combining
numerical models and multiple tracer observations such as
noble gases [Hamme and Emerson, 2002]. We do not claim
that the numerical model used here can adequately capture
various physical processes controlling the disequilibrium of
surface oxygen, such as bubble formation and mesoscale
ocean circulation. Further theoretical and experimental
studies are required to understand and quantify these
physical processes controlling the saturation state of oxygen
in the polar surface oceans.
[13] Acknowledgments. We would like to acknowledge the fruitful
collaboration with Stephanie Dutkiewicz and Payal Parekh of MIT for the
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development of the physical-biogeochemical model. TI is grateful for the
support from the Office of Polar Program at NSF. MJF and EAB are
grateful for support from NSF grant OCE-0136609 and OCE-0350672.
Two anonymous reviewers provided useful comments for improving the
manuscript.
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