Chapter 6.2 Ocean transport of fresh water

475

CHAPTER

6 .2 Ocean Transport of Fresh Water Susan E . Wijffels

6 .2 .1 The importance of freshwater transport The majority of water cycling through the atmosphere derives from the ocean, causing either distillation or freshening of ocean surface waters . The steady-state portion of the salinity distribution in the ocean reflects a balance between the exchange of fresh water at the surface and compensating mixing and advective ocean processes . Hence, in the long-term, the transport and circulation of fresh water (and salinity) by the ocean mirrors the transport of latent heat by the atmosphere, which comprises as much as 1 .5 PW of the total of 4 PW of poleward atmospheric energy transport (Rosen, 1999) . Oceanic freshwater transport is therefore a fundamental parameter in the planetary energy budget (see Grassl, Chapter 1 .1 and Clarke et al ., Chapter 1 .2) . As both atmospheric, ocean and coupled climate models improve, the need for accurate estimates of the components of the planetary energy budget becomes increasingly acute . Without good estimates of the partition of energy transport between the ocean and the atmosphere, and between diabatic and adiabatic processes, diagnosing errors in the models or assessing the effects of new physical parameterizations is made more difficult . For oceanonly models in particular, the lack of accurate estimates of surface freshwater forcing has fuelled the use of unphysical model boundary conditions, the effect of which will be discussed below . Freshwater forcing imposes a vertical velocity at the ocean's surface, resulting in a Sverdruplike dynamical response that can be significant

OCEAN CIRCULATION AND CLIMATE ISBN 0-12-641351-7

(Huang and Schmitt, 1993) . In addition, the ocean thermohaline circulation, as modelled in coarseresolution climate models, is also sensitive to the strength of freshwater forcing, especially at high latitudes (Weaver et al., 1993 ; Rahmstorf, 1996) . The stability of the thermohaline circulation, in particular, is sensitive to both the freshwater forcing and the strength of ocean vertical mixing (Zhang et al ., 1999b) . Historically, as has been the case for ocean heat transport (see Bryden and Imawaki, Chapter 6 .1), most estimates of oceanic freshwater transport derive indirectly from measurements of atmospheric vapour transport (for a recent example, see Trenberth and Guillemot, 199 ) or from surface observations (ship and island) of rainfall rates and the variables required to estimate evaporation (e .g . da Silva et al., 1994) . Since the 19 0s, estimates based on in-situ ocean data started to appear . These direct estimates of oceanic fresh-water transport provide both an independent check of the atmospherically based numbers as well as insight into the oceans' response to the surface forcing.

6 .2 .2 Indirect estimates of oceanic freshwater transport In recent years, many new estimates of both atmospheric and surface fluxes of fresh water have appeared, based on the development of dataassimilating atmospheric general circulation models and the availability of new satellite data sets that measure variables used to deduce evaporation, the radiometric signature of precipitation or

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476

SECTION 6 LARGE-SCALE OCEAN TRANSPORTS

the net moisture content of the atmospheric column (Xie and Arkin,1997) . Estimates based on data-assimilating atmospheric modelling systems are difficult to assess for several reasons . First, the model output often doesn't obey total mass conservation (Trenberth and Guillemot, 199 ) . Second, the lack of atmospheric profile data over the oceans is a severe limitation for these estimates : the assimilation of scant island station data into these systems produces `bull's eyes' in the flux fields indicative of biases between the models and observations over the oceans (Trenberth and Guillemot, 199 ) . As the vapour exchange with the oceans dominates the air-sea fluxes, these flux biases are problematic for global budgets . Surface flux estimates of evaporation rely on empirical relations based on either radiometric parameters measured from satellites or marine meteorological measurements such as wind speed, relative humidity and sea surface temperature measured from ships and islands . Though constantly improving, these formulae are required to apply under a large range of conditions and can also suffer small biases which, when accumulated over large areas, can dominate basin totals . Shipbased measurements can also suffer biases . Precipitation is particularly challenging to estimate over the ocean, due to its sporadic nature in both time and space . Here, satellite estimates may be the only means of progress, but these also rely on empirical algorithms that require `tuning' . Lack of suitable open ocean time series for calibrating the satellite estimates is a limitation, as island stations likely suffer biases due to local small-scale orographic effects . It is not the intention here to catalogue or assess the presently available flux estimates exhaustively, but rather to highlight the need for suitable independent estimates of large areal averages that can be used to assess the new data sets being produced, and in particular identify the small but crucial biases that impact dramatically on ocean circulation models . To demonstrate the difficulties, a set of recent flux estimates has been assembled (Table 6 .2 .1) . Each member of the ensemble is assumed to be as valid as the next . As a group, these estimates and their variance can therefore be used as a gross quantitive measure of the present state of knowledge about the mean annual exchange of water

between the ocean and the atmosphere . Based on the products in Table 6 .2 .1, ensembles of flux estimates have been formed : an eight-member ensemble of precipitation (P) estimates, a six-member ensemble of evaporation (E) estimates, and a 13member ensemble of E-P (net freshwater flux) estimates . After mapping to a common geographic grid, the pointwise means and standard deviation of the ensembles were formed . The ensemble mean E-P (Fig . 6 .2 .1a) shows the major source regions of atmospheric vapour : the subtropical oceans under the atmospheric highpressure centres . The major atmospheric water sinks are the tropical convergence zones and the subpolar oceans . Interestingly, while the northern hemisphere western boundary currents undergo strong cooling (Bryden and Imawaki, Chapter 6 .1), they are subject to only weak freshwater forcing . Over both the Kuroshio and Gulf Stream there occurs almost total cancellation between enhanced evaporation (Fig . 6 .2 .2a) and enhanced precipitation (Fig . 6 .2 .3a), both of which are likely associated with atmospheric storm tracks . Variability within the ensemble of net flux estimates (Fig . 6 .2 .1b) reveals a global background uncertainty of about 250 mm yr -1 , which when integrated over the Pacific down to 30°S adds up to 1 x 10 9 kg s -1 of freshwater transport, which is as large as the natural flux itself (Wijffels et al ., 1992) . The largest uncertainty occurs over the tropics and the mid-latitude storm tracks, as well as a region in the Southeast Pacific off Chile, which is an important water mass formation region (McCartney,1977) . Variability among evaporation products is much less (Fig . 6 .2 .2b), suggesting the bulk of the uncertainty in the total water flux derives from estimates of precipitation, which is confirmed by Figure 6 .2 .3b . Not surprisingly, the difficulty of measuring precipitation over the open ocean dominates errors in the net water flux to the ocean .

6 .2 .3 Impacts of uncertainties on model development The large uncertainties in estimates of freshwater fluxes over the oceans has retarded efforts to model the climate system . In the realm of atmospheric modelling, the lack of a reliable benchmark against which to compare models is a difficulty, with differences among models often smaller than



6 .2

Table 6 .2 .1

Ocean Transport of Fresh Water

477

Surface flux estimates used in this study. Surface latent heating rates were

converted to an equivalent evaporative water flux (E) using the latent heat of vaporization at the appropriate seasonal sea surface temperature taken from Levitus and Boyer (1994a) . Baumgartner and Reichel (1975) only provide zonal flux integrals and thus is not used in Figures 6 .2 .1-6 .2 .3 . In three data sets runoff R is also available . Source

Type

Parameters

Temporal coverage

Surface observations (SOC)

Josey et al. (1996)

E P

Climatology

Surface observations (LOADS)

da Silva et al. (1994)

E, P

Climatology

Surface observations

Baumgartner and Reichel,

E-P, R

Climatology

E

Climatology

(1975) [not gridded] Surface observations

Esbensen and Kushnir (19 1)

Surface observations

Oberhuber (19

E,P

Climatology

Blended satellite and surface

Legates and Wilmott (1990)

)

P

1979-96

Surface observations

Jaeger (19 3)

P

Climatology

Blended satellite and surface

Xie and Arkin (1997)

P

1979-9

P

1979-92 Climatology

observations

observations and model output Blended satellite and surface observations (Limb90/MSU)

Tod Mitchell, personal communication (1999)

Satellite observations

Jourdan et al. (1997)

Atmospheric re-analysis

Keith (1995)

E-P E-P, R

Atmospheric re-analysis (ECMWF)

Barnier et al . (1995)

E

19 6-

Atmospheric re-analysis (NCEP)

Trenberth and Guillemot

E-P

1979-95

E, P, R

Climatology

19 9

(ECMWF)

(199 ) Atmospheric re-analysis (NCEP)

Kalnay et al . (1996)

those among observational estimates . For instance, Gaffen et al . (1997) concluded that the majority of atmospheric models participating in the Atmospheric Model Intercomparison Project (AMIP) overestimate the poleward transport of moisture compared with estimates based on direct observations . However, such direct observations are based on mostly non-existent observations over the oceans (which cover the majority of the planet's surface area) and have unknown accuracy (Oort and Piexoto,19 3) . Ocean modelling efforts, however, suffer an even greater disadvantage . Ocean models require surface freshwater forcing . When forced with `observed' freshwater fluxes the models can drift off to unrealistic states (McWilliams, 1996), and it has been difficult to distinguish the cause : inaccurate model physics versus errors in the forcing field . To avoid this problem (and because freshwater forcing of surface bouyancy has been considered of secondary dynamical importance in comparison to thermal forcing), the practice of

relaxing to an observed salinity field at the surface (Haney condition for salinity) has become the convention . While a Haney condition for heat is plausible given both sensible and radiational cooling of warm temperature anomalies by the atmosphere, such a flux formulation for salinity has not been physically justified . In the realm of coupled climate modelling, the component ocean and atmospheric models are often first 'spun-up' separately before coupling . In order to maintain a stable `realistic' climate, flux adjustments are introduced into the coupled system, which are essentially the difference between the flux out of the atmosphere model spun-up over the observed sea surface temperature field, and the flux into the ocean model run under relaxation forcing (Gordon and O'Farrell, 1997) . This common expedient propagates freshwater flux errors from the unrealistic ocean spin-up under surface relaxation into the coupled system . Typical freshwater flux adjustments have magnitudes as large as the natural fluxes themselves, and the resulting

Wijffels



47

SECTION 6

LARGE-SCALE OCEAN TRANSPORTS

0°N 60° 40 20

20° 40° 60° 0 S (a) 0 N 60° 40° 20°

20° 40° 60° 0 S (b)

50° E

100°

150°

200°

250°

300°

350°

Fig . 6 .2.1 Average (a) and standard deviation (b) of an ensemble of 13 estimates of E-P over the oceans based on the products in Table 6 .2 .1 .The contour interval is 250 mm yr -1 , with the zero contour dotted .

total water flux to the ocean has spatial patterns that are highly unrealistic . McWilliams (1996) points out that with such unrealistic forcing, it is no surprise that the model's salinity field is also unrealistic . The result is that the use of salinity, the next best observationally known ocean quantity after temperature, has been greatly limited as a means of identifying and correcting physical errors in ocean models .

6.2 .4 Direct ocean estimates of freshwater transport Direct estimates of oceanic freshwater transport are an independent means of measuring the freshwater exchange between the ocean and atmosphere, and so have the potential to address some of the problems outlined above . In order to be of use, however, the strengths and limitations of these estimates need to be clearly examined .



479

0 N

60°

40°

20°

20°

40

60°

0 S (a)

0°N

60°

40°

20°

0°

20°

40°

60°

0°S (b)

50°E

100°

150°

200°

250°

300°

350°

Fig . 6 .2 .2 Average (a) and standard deviation (b) of an ensemble of six estimates of evaporation over the oceans based on the products in Table 6 .2.1 . Units are mmyr -1 and the contour interval is 250 mm yr -1 .

Direct estimates of oceanic freshwater fluxes are determined in the same way as those of heat (Bryden and Imawaki, Chapter 6 .1), using long hydrographic lines that enclose volumes of ocean, for which budget equations can be written . The technique relies on the assumptions that we can measure the steady-state portion of the velocity and the salinity field, and that geostrophic dynamics apply everywhere, except for an Ekman balance in

the mixed layer . Freshwater fluxes achieved by the temporal rectification of the salt and velocity fields at either the seasonal or eddy time scales are ignored in these estimates, although, as seen below, these fluxes have been found to be small over most of the ocean . For a volume of ocean enclosed by a hydrographic line, salt conservation applies in the steady state as the flux of salt through the atmosphere is

Wijffels



4 0

SECTION 6 LARGE-SCALE OCEAN TRANSPORTS

0°N

60°

40°

20°

0°

20°

40°

60°

0 S (a) 0°N

60°

40°

20°

0°

20°

40°

60°

0 S (b)

50°E

100°

150°

200°

250°

300°

350°

Fig . 6 .2 .3 Average (a) and standard deviation (b) of an ensemble of eight estimates of precipitation over the oceans based on the products in Table 6 .2 .1 . Units are mm yr -1 and the contour interval is 250 mm yr -1 .

negligible :

Mass conservation is written as : ffpsvdxdz=T ~(6 S .2 .1)

where (p is the in-situ density, S the salinity, v the cross-track velocity and x the along-track direction . T S represents the total salt transport associated with the interbasin exchange (e .g . Bering Strait or the Indonesian Throughflow) .

.2 .2) ffpvdxdz+[P_E+RI=T;(6 M where E, P and R are the net fluxes respectively into the surface of the ocean volume of evaporation, precipitation and runoff, and T M is the interbasin mass transport . In the haline ocean, the pure



6 .2

Ocean Transport of Fresh Water

water part of the above total mass transport is just : 55p(1_S)vdxdz+[P_E+RI=T 1 M _T S

(6 .2 .3)

Hence, the LHS of (6 .2 .3) is, strictly speaking, the oceanic freshwater transport . Generally, estimates of the mass flux across a section have uncertainties much larger than P - E + R, suggesting that (6 .2 .3 ) alone is not useful for calculating the surface forcing . We thus define a section areal average salinity and its deviation :

4 1

when the two are combined as in (6 .2 .5) . Ignoring small changes in p (which in practice can be carried with either v or S), and replacing v and S' in (6 .2 .5) with the observed quantities, the error in the freshwater flux due to measurement errors can be written as : error[p-E+R] ff p [S a v e + Sevo + Seve ] dx dz S

(6 .2 .7)

1 psu- 5 Sv+0 .1 psu-, 30 Sv+0 .1 psu- 5 Sv -

35 psu

ffSdxdz S' =S-S

(6 .2 .4) (6 .2 . )

ffdxdz For simplicity, we also assume that the interbasin flux of salt occurs at a known salinity, S ;, so that the salt flux is just T=S T; M . Combining (6 .2 .1), (6 .2 .2) and (6 .2 .4), the surface freshwater flux can now be written as a simple product of the salinity deviation and the velocity field : T1MS ;- ff

pS'vdx dz

[P-E+R]=-

( 6 .2 .5)

S Here the first term on the right can be termed the `leakage' term and the second term is due to correlations of salinity and velocity across the section, which effect a freshwater transport . The above equation only applies to perfect measurements of the long-term mean salinity and velocity field, and neglects fluxes of fresh water (salt) due to horizontal mixing and temporal correlations of salt and velocity . Ignoring the latter for the moment, the velocity and salinity measured by a hydrographic section (and the associated circulation analysis) has an error such that the true long-term mean is given by the sum of the measured velocity v o and its error v e : vtrue = vo

+

ve

(6 .2 .6 )

The same is true for the salinity measurement . Note that for most studies using hydrographic sections, a circulation error of ff ve dx dz~2 Sv is usually assumed, which is much larger than P - E + R . However, this circulation error has an associated highly correlated error in the salt balance equation (55VS dx dz), which largely cancels

H0.lSSv

+ 10 .09 Sv + 0 .01 Svl

0 .17 Sv

Each of the terms in equation (6 .2 .7) can be grossly estimated using scale arguments to get an upper bound on the freshwater errors (6 .2 . ) . The first error term is the spatial correlation of the net error in the circulation field (conservatively -5 Sv) and the real salinity variation ('-1 psu) ; the second term is due to correlations between errors in measuring salinity ('-0 .1 psu) and the true velocity field ; and the last term is the correlation of errors in the velocity and salinity field . While the synoptic salinity field is well resolved by modern hydrographic sections, it is the longterm mean field that is needed for equation (6 .2 .5) . Temporal changes in salinity can be addressed in only a few regions where time series exist . Generally, comparison of historical and modern sections has revealed only small changes in the large-scale salinity field . Bindof f and Church (1992) estimate that the freshening of the Antarctic mode and intermediate waters they observe in the southern hemisphere subtropics between 1960s and 1990s only requires a change of about 40 mm yr -1 in freshwater flux at the latitudes where these waters are formed . Compared with the uncertainty of around 250mmyr -1 in present-day mean flux estimates (Fig . 6 .2 .1) these changes are small . Dobroliubov (1997) found differences of 0 .14 Sv or less between freshwater transports based on Atlantic sections made in the early 19 0s and 1990s . Larger differences were found with sections made in the 1950s and 1960s, though these are suspected of deriving from poor data resolution and might not be real . In the tropics, interannual changes in flux distribution

W+jffels



4 2

SECTION 6 LARGE-SCALE OCEAN TRANSPORTS

and sea surface salinity are as large as the seasonal cycle and, due to the faster advection time scales in the tropical ocean, are more likely to penetrate below the mixed layer and alter the subsurface salinity field . Thus tropical sections might be more likely to suffer interannual biases than those at higher latitudes . The possibility that temporal fluctuations in the velocity and salinity field rectify to produce a mean flux must also be examined, as this flux has been

ignored in equation (6 .2 .5) . In most of the ocean the dominant velocity variability is at mesoscale eddy time and space scales (Stammer, 1997b) and in the seasonal cycle . Using estimates of lateral mixing rates based on eddy length scales, energy levels from satellite altimeter data and mean climatological salt gradients, Stammer (199 ) estimates that the eddy freshwater fluxes are small (<0 .05 Sv) except along western boundary currents extensions ('-40° of latitude) .

60°N

30°

30°

60°S ~' 00 30°E 60° 90° 120°

~-' {-' I-' I-' 150° 1 0° 150° 120 0 900 60°

r-' 30°W

0°

0 .1

60°S

40°

20°

0°

20°

40°

60°N

Fig . 6 .2 .4 Seasonal rectification flux of the Ekman mass flux and sea surface salinity : (a) shown as a vector flux ; (b) integrated across the ocean basins (Sv) .



6.2

Ocean Transport of Fresh Water

A second source of temporal rectification between S and v is in the seasonal cycle in the Ekman layer . Seasonal changes in wind strength can coincide with changes in freshwater forcing and salinity . To examine how large these fluxes might be, the rectification flux was calculated for each ocean basin using the Hellerman and Rosenstein (19 3) wind-stress climatology and the Levitus and Boyer (1994b) sea surface salinity (Fig . 6 .2 .4) . Outside the tropics the freshwater flux is less than 0 .02 Sv, which is negligible . These estimates, along with those from Stammer (199 ), agree with those from high-resolution modelling studies, which find the temporal rectification terms (over all time scales) give fluxes of < 0 .1 Sv outside of the tropics (McCann et al ., 1994) . In total then, the expected uncertainty in the direct flux estimates might be as large as 0 .175v outside of the tropics (6 .2 . ), but may reach 0 .3 Sv in the tropics . Independently, Dobroliubov (1997) finds similar uncertainties . To go beyond the simple scaling argument in equation (6 .2 .7) requires simultaneous time series of both velocity and salinity overbasin scales, measurements that are not yet available .

4 3

6.2 .5 Comparison of direct and indirect flux estimates In comparing direct ocean flux estimates with estimates of the surface flux, runoff from the continents must be taken into account (Fig . 6 .2 .5) . Despite attempts to catalogue the runoff of major rivers (Hils, personal communication, 1999) there appear to be few global estimates of runoff . Here we utilize Baumgartner and Riechel's (1975) compilation, which, globally, roughly agrees with the runoff estimated from the net E-P over land based on Keith's (1995) analysis of ECMWF output and the NCEP (Kalnay et al., 1996) long-term runoff (Fig. 6 .2 .5) . Note that in Keith's (1995) estimates there are areas of negative flux (where the accumulation decreases) expressing the non-conserving nature of the re-analysis moisture field . Given the limited availability of global runoff data sets, assessing the errors in the runoff fluxes is difficult and will remain an added uncertainty in these comparisons . Here, the Baumgartner and Reichel (1975) runoff is added to the surface fluxes integrated over the ocean basins to predict the ocean transport of fresh water .

1 .5 Baumgartner and Reichel(1975) Keith(1995) NCEP Largest 161 Rivers (GRDC)

44

z 0 rn

1

E O

m 44 44 4,

0 0°S

60°

40°

20°

0°

20°

40°

60°

0°N

Fig . 6 .2 .5 Total runoff into the oceans cumulated from north to south (Sv) from four sources : Baumgartner and Riechel (1975), Keith (1995), Kalnay et al. (1996) and the largest 161 rivers as catalogued at the Global Runoff Data Center (Hits, personal communication, 1997) . Wijffels

4 4

SECT ON 6 LARGE-SCALE OCEAN TRANSPORTS All three ocean basins exchange large amounts of water through linking passages : the Southern Ocean, the Indonesian Archipelago and the Bering Strait . As these interbasin fluxes are much larger than the exchange through the atmosphere, we choose instead to present only the divergent part of the oceanic freshwater flux (in contrast to Wijffels et al., 1992) . This is equivalent to removing an unknown constant equal to the Pacific-Indian Throughflow for the Indian and South Pacific Oceans, and the Bering Strait flow in the North Pacific and Atlantic . Hence in the plots to follow, the direct transport estimates shown will be the divergence relative to the entrances of the Bering and Throughflow straits (South of Mindanao, Philippines) . While the size and salinity of the Bering Strait flow is relatively well known (Coachman and Aagaard, 19 ), that of the Pacific-Indian Throughflow is not . Hence, investigators have had to make assumptions about both its size and salinity in order to generate an estimate of the basin divergence - that is to calculate the `leakage' term in equation (6 .2 .4) . Direct estimates for the longterm average Throughflow range between 5 and 10 Sv (Meyers,1996 ; Molcard et a!.,1996 ; Gordon et a! ., 1999a), though synoptically, higher variability (0-20 Sv) is observed (Fieux et al ., 1994 ; Bray et al., 1997) . For a 5-10 Sv transport range the resulting freshwater estimates between the Throughflow entrance channels and 30°5 WOCE section in the Pacific is only about 0 .2 Sv (Wijffels et al ., 2000), consonant with our estimates of the general uncertainty. Clearly, however, increased knowledge of the freshwater flux associated with the Throughflow is needed . The majority of direct flux estimates derive from single-section or regional analyses of long hydrographic lines, many of which were accomplished during WOCE (Table 6 .2 .2) . To date few truly global syntheses have been completed, with the exception of MacDonald (1995) . The latter study, however, specified freshwater fluxes using Baumgartner and Reichel (1975) and so is not included here . The flux divergences produced from the Southern Ocean analysis of Sloyan and Rintoul (2000b) are included in basin integral plots (Figs 6 .2 .6-6 .2 . , see Plates 6 .2 .6-6 .2 . , p . 492) by assuming zero freshwater flux from the Antarctic continent . Figures 6 .2 .6-6 .2 . (see Plates 6 .2 .6-6 .2 . , p . 492) present the divergent part of the freshwater transport for each ocean basin found from

accumulating the surface and runoff fluxes southwards from basin reference points . In the Indian Ocean transports are relative to a zero flux condition at the Through- flow channels (Pacific side) ; for the Pacific, zero flux is applied at Bering Strait and the Throughflow channels (Pacific side) ; while the Atlantic curve is relative to a zero flux at Bering Strait, and thus includes the freshwater fluxes into the Arctic Oceans . In the Indian Ocean, only two latitudes are currently constrained by direct flux estimates : 32 and 1 °S (Fig. 6 .2 .6, see Plate 6 .2 .6, p . 492) . Over the evaporative zone between latitudes 15 and 40°5, the indirect flux estimates are fairly consistent (similar slopes) and are also in reasonable agreement with the direct estimates . It is in the regions of high precipitation north of 10°S and south of 40°S that the curves diverge, confirming again the large scatter among available estimates of precipitation over the ocean . Note in particular the 0 .5 Sv variability in the net freshwater divergence north of 1 °5 . The Atlantic Ocean is best covered by direct flux estimates (Fig . 6 .2 .7, see Plate 6 .2 .7, p . 492), which are remarkably consistent, except for the estimates of Dobroliubov (1997) at 24 ° N . Nearly all of the major transport maxima are delineated by the direct estimates . Again, the indirect estimates diverge most strongly over regions of high precipitation . When integrated over the Atlantic between Bering Strait and 40°S, the indirect transport divergences range between 1 .0 and - 0 .1 Sv, while the direct estimates indicate very little net fresh-water divergence over the basin . Problems with biases in the indirect transport estimates are even more pronounced in the Pacific Ocean due to its huge size (Fig . 6 .2 . , see Plate 6 .2 . , p . 492) . Not surprisingly the indirect estimates vary wildly over the South Pacific where little in-situ atmospheric data is available . Here again, despite the variations in assumptions made to close the ocean mass balance, Throughflow sizes and different data sets, the direct estimates are quite consistent, and show much less scatter than the indirect estimates . It is remarkable that despite 25 years between sections, the estimates of Wunsch et al . (19 3) and Wijffels et al. (2000) are very similar . Though a more consistent and larger set of direct ocean transports estimates can be anticipated from the WOCE synthesis, the ensemble available to date can already be used to reassess ideas about the interocean exchange of fresh water . Using



6 .2

Table 6 .2 .2

Ocean Transport o Fresh Water

4 5

Direct estimates of oceanic freshwater convergence north of a given latitude .

Units are in kg 9 s - ~ . Negative numbers indicates the ocean to the north receives a net excess of precipitation and runoff over evaporation . Atlantic convergences are shown relative to the Bering Strait and so include the Artic Ocean . Pacific convergences are relative to Bering Strait and the Indonesian Throughflow passages south of the Philippines Latitude

Pacific

nd an

Atlant c Tota

5 °N

Source Dobroliubov (1997) [19 0s data Dobroliubov (1997) [1990s data

0.17

Bacon (1997) Roemmich and McCallister ( 19 9)

0.26

Dobroliubov (1997) [19 0s data]

0.25

0. 2

0. 6

Dobroliubov (1997) [1990s data] Raemmich and McCallister ( 19 9)

-0.49

Dobroliubov (1997) [19 0s data] 0. 3

24° N

Dobroliubov (1997) [1990s data] Roemmich and McCallister (19 9) Bryden et al. (1991)

-0 .23

Hall and 'Bryden (19 2) Dobroliubov (1997) [19 0s data] Dobroliubov (1997) [1990s data]

11 ° N 10°N

Friedrichs and Hall (1993) 0.03

0 .15

11 °S 12°S 17°S

Wijffels et a!. (1996a)

-0.23

Holfort and Siedler (1997)

-0.15

Holfort and Siedler (1997)

-0.33

Sloyan and Rintoul (2000b)

-0.30

Tsimplis et al. (199 )

1 °S

-0.14*

Sloyan and Rintoul (2000b)

19°S

-0.03T

Sloyan and Rintoul (2000)/Robbins and To©le (1997) -0.10

23 ° S 2 °S

-0.43

0 .12

Holfort and Siedler (1997) Holfort and Siedler (1997)

-0.06

Wunsch et al. (19 3)

27 ° S

0 .16

Holfort and Siedler (1997)

30 ° S

0.24

Holfort and Siedler (1997)

32 ° S

0.09

Tsimplis et al. (199 )

0.06 0 .33*

Wijffels et al. (2000) 0.20*

Sloyan and Rintoul (2000b)

0.31 40 ° S 43 ° S

0.61

Robbins and Toole (1997)

0 .01*

Sloyan and Rintoul (2000b)

0.02

Saunders and King (1995b)

-0 .0

Wunsch et al. (19 3)

Note : * Indicates a convergence for the ocean volume south of the designated latitude and not the ocean volume to the north, as for the other estimates ; an estimate derived from the sum of the 32 °S flux number and the divergences between 32° S and 1 °S . Numbers in the column marked `total' are a sum of the italicized estimates in the various basins for that latitude band .

the Baumgartner and Reichel (1975) climatology, Wijffels et al . (1992) deduced that the Pacific received an excess of roughly 0 .5 Sv of fresh water, which was then redistributed by the interocean circulation to the evaporative basins of the Atlantic and Indian Oceans via exchange through the Indonesian Throughflow, Bering Strait and

Southern Ocean . However, the new direct ocean estimates indicate a different scenario . Figure 6 .2 . (see Plate 6 .2 . , p . 492) shows that between Bering Strait and 30°S the freshwater divergence over the Pacific is small, indicating a net balance of evaporation and precipitation over that basin . Direct estimates for the Atlantic/Arctic also

Wijffels

486

SECTION 6

LARGE-SCALE OCEAN TRANSPORTS

suggest a net divergence of fresh water much smaller than previously thought: Holfort and Siedler's (1997) estimates of a 0.24 Sv convergence between Bering Strait and 30°S is roughly half that predicted by Baumgartner and Reichel (1975), while the estimates at 40°S indicate almost no net divergence over the Atlantic/Arctic. The Indian Ocean direct estimates, however, are consonant with a large net evaporation, as predicted previously (Fig. 6.2.6, see Plate 6.2.6, p. 492). Hence, if the Pacific Ocean is not the source of the excess oceanic fresh water required to supply the Indian deficit, only one possibility remains: excess precipitation and ice melt over the Southern Ocean. Indeed Sloyan and Rintoul estimate that 0.54 Sv of excess fresh water is removed by the Southern Ocean circulation south of 30-40°S, highlighting the importance of the Southern Ocean in the global oceanic freshwater balance. Based on the independent estimates in Table 6.2.2, the excess evaporation occurring north of Sloyan and Rintoul's lines add up to 0.39 Sv. The freshwater global balance is thus 0 . 5 4 - 0 . 3 9 = 0.15±0.29Sv: indistinguishable from zero. 6.2.6 Mechanisms of oceanic freshwater transport Stommel and Csanady (1980) pointed out that ocean heat and freshwater transport is related to the rates of conversion of water from one part of temperature-salinity (T/S) space to another. They went further and attempted to model this process assuming salinity was a simple function of temperature. Recent analyses of freshwater fluxes across ocean sections and in general circulation models (Rahmstorf, 1996) reveal that this assumption is incorrect, though the underlying idea remains a powerful one, as it links forcing fluxes to water mass volume and exchange in T/S space. Speer and Tziperman (1992) recast this approach in terms of density classes and express the competition between density class conversion by surface fluxes and that by interior diapycnal mixing, an idea employed in Sloyan and Rintoul's (2000b) analysis. The challenge in analysing ocean sections will be in distinguishing the water mass conversion occurring at the surface and that due to internal mixing. Use of freshwater fluxes will be a key element in this work. Definition of a tracer transport mechanism across an ocean section is somewhat ad hoc. Hall

and Bryden (1982) chose to form zonal averages (and deviations) of velocity and properties on pressure surfaces and termed the resulting products the 'overturning' component of tracer transport, while the residual, associated with the correlation of velocity and tracer at a pressure level, was termed the 'horizontal' or gyre component. A similar decomposition can be carried out within density layers (e.g. Wijffels et al., 1996a; Robbins and Toole, 1997), and as density is largely determined by temperature, more closely relates back to Stommel and Csanady's (1980) suggestion. Freshwater divergence is also achieved by the interbasin flows, the 'leakage' term in equation (6.2.4). Unfortunately, few detailed decompositions of the oceanic freshwater flux across hydrographic lines are available, but those that are reveal interesting mechanisms. Wijffels et al. (1996a) find that at 10°N in the Pacific, the small net freshwater divergence over the Pacific relative to Bering Strait is achieved by a balance between three major flux mechanisms: net export of very fresh water through Bering Strait; northward freshwater transport by the shallow meridional circulation where the northward Ekman flux is fresh and the compensating southward thermocline flow more saline; and a southward freshwater flux by the subthermocline horizontal circulation where salty South Pacific waters flow north in the eastern Pacific and fresh North Pacific Intermediate water flows south in the western Pacific. The deep and bottom water circulations in the low-latitude Pacific achieve little net freshwater transport. In the Indian Ocean, Robbins and Toole (1997) diagnose a large net evaporation north of the section at 32°S of 0.31 ±0.09Sv. They find three mechanisms acting to import fresh water to supply this net evaporation: a leakage term associated with the inflow of fresh Indonesian Archipelago waters that are distilled and leave across 32°S as salty thermocline waters; upwelling of deep and intermediate waters and their export as salty thermocline waters; and horizontal inflow of fresh Antarctic Intermediate Water in the east that leaves saltier in the west. Wijffels et al. (2000) have similarly found the import of fresh water by the recirculation and modification of Antarctic intermediate waters to be an important transport mechanism at 32°S in the Pacific. It is likely that the production and northward export of fresh

6.2

487

Ocean Transport of Fresh Water

mode and intermediate waters from the Southern Ocean to the southern subtropical gyres may be the single most important mechanism for balancing the large net flux (0.61 Sv; see Table 6.2.2) received by the oceans south of 30°S from the atmosphere and through ice flows. The ability of ocean general circulation models to reproduce the estimated freshwater fluxes and their mechanisms will be a stringent test of model realism. In ocean-only models, surface flux forcing will determine the net equilibrium transports (unless the problematic relaxation boundary conditions are used), but internal model physics will determine how this flux is achieved. 6.2.7 G l o b a l budgets

Direct oceanic freshwater transport estimates are presently available in all basins at six latitudes (Table 6.2.2), allowing the global meridional freshwater transport to be examined. Since few of the major rivers flow meridionally, the zonally integrated meridional ocean transport of fresh water

is largely equal and opposite to transport in the atmosphere (Wijffels et al., 1992). These estimates can be compared with direct estimates of atmospheric moisture transport or those produced by atmospheric general circulation models (Fig. 6.2.9). Gaffen et al. (1997) examined the moisture budget in atmospheric models in the AMIP and concluded that most models overestimate the poleward flux of moisture, based on a comparison with Oort and Piexoto's (1983) observations. The available direct ocean measurements are still too sparse for any conclusions to be drawn on this issue, though generally the direct estimates agree more with the AMIP ensemble and less with Oort and Piexoto's estimates, except at higher latitudes (30°S and 50°N) where the opposite is true. It is worth noting that Oort and Piexoto's estimates are reliant on scarce in-situ atmospheric data over the large tropical and southern hemisphere oceans, and so are likely less reliable there. To constrain the total meridional moisture transport in the atmosphere more usefully, a larger number of direct ocean transport estimates are required as well as a better

soc COADS Direct AMIP Interquartile Range Oort and Piexoto (1983)

80°N 9

1

Fig. 6.2.9 Total ocean meridional freshwater transport (10 kgs~ ): estimates based on ocean data are shown using the sources in Table 6.2.2; indirect estimates are based on indicated sources plus the Baumgartner and Riechel (1975) continental runoff. Shaded is the interquartile range for the atmospheric models participating in AMIP as reported by Gaffen et al (1997). Oort and Piexoto's (1983) direct atmospheric estimate is marked as x-x. Wijffels

488

SECTION 6

LARGE-SCALE OCEAN TRANSPORTS

estimate of their errors - those shown in Fig. 6.2.9 are based on the scale arguments discussed earlier and so are rather conservative. 6.2.8 Summary Direct estimates of oceanic freshwater transport can be determined to useful accuracy from the analysis of ocean sections. Most significant is that, in contrast to values based on integrating surface fluxes, the error in the direct ocean estimates are largely independent of the size of the area of ocean enclosed. While they can only shed light on the long-term mean flux at locations where sections have been made, they provide estimates of large areal integrals that are not reliably gained from any other source. It is the combination of these estimates with those that do resolve both the true spatial and temporal variability (e.g. satellite data sets) that holds much potential. Until surface flux products or atmospheric models improve, the direct estimates are likely to remain the most accurate of large areal averages of the water flux between the ocean and the atmosphere. When synthesized globally, direct oceanic freshwater transports can be used to assess the performance of atmospheric and coupled models. Estimates available to date show more agreement with atmospheric model output than with direct estimates of atmospheric moisture fluxes. Presently, direct oceanic freshwater fluxes are not as well reported or analysed as are the companion heat fluxes. While most estimates are fairly consonant with each other and error estimates made here based on simple scaling arguments, others are quite anomalous (such as Dobroliubov, 1997). Tracking down the source of these differences requires a detailed breakdown of the mechanisms making up these fluxes. In respect of freshwater flux estimates, the full potential of the WOCE data set is yet to be realized. We can anticipate global, better-documented and more accurate freshwater transport estimates to be produced. In addition there may be great potential in the idea of 'tuning' surface flux products using direct ocean estimates to remove flux biases (e.g. Isemer et aL, 1989; da Silva et ai, 1994). This may lead to flux products that are accurate enough to force ocean climate models directly with confidence

and thus remove the need to use the unphysical surface relaxation boundary condition. This in turn will open the door for the much more meaningful use of salinity as a model diagnostic. The limited number of direct oceanic freshwater transport estimates presently available points to the need to reassess our view of basin-wide balances and interbasin exchanges. Already, the idea that the Atlantic Ocean is highly evaporative and must import large amounts of oceanic fresh water must be revised: the new estimates suggest nearzero gain down to 40°S (Saunders and King, 1995b). We can also deduce a larger freshwater exchange than previously estimated between the Southern Ocean and the basins to the north via the injection of Antarctic Intermediate Waters into the southern hemisphere subtropical gyres. It is noteworthy that this horizontal v' S' flux effected by the subtropical gyres is not accounted for in simple box models of the global thermohaline circulation (e.g. Broecker et aL, 1990). Such models must fold this upper ocean transport into the deep water component of the global thermohaline circulation, confusing the role of freshwater transport as a control on the circulation (Rahmstorf, 1996). More careful and detailed studies of how the ocean transports fresh water are required to pick apart the relative role of the shallow wind-driven gyres relative to that of the deep circulation in balancing the surface forcing. The data collected during WOCE will make such analyses possible.

Acknowledgements Thanks is given to the following: Dr Hils from the Global Runoff Data Center for providing the Center's runoff data; Sergie Dobroliubov for providing his results; Steve Covey for access to the AMIP numbers; Todd Mitchell for advice and access to some of the new climatologies; Siobhan O'Farrel and Tony Hirst for use of the CSIRO coupled model output; Stefan Rhamstorf for bringing some of the new modelling work to my attention; Jim Mansbridge for help with data processing; and the International WOCE Office for sponsoring my attendance at the WOCE Conference in Halifax. This article was supported by the Australian Climate Change Research Project funded by Environment Australia.