two transatlantic sections: meridional circulation and heat flux in the subtropical north atlantic...

Deep-Sea Research, Vol. 32, No. 6, pp. 619 to 664, 1985. 0198-0149/85 $3.00 + 0.00 Printed in Great Britain. ~ 1985 Pergamon Press Ltd.

Two transatlantic sections: mcridional circulation and heat flux in the subtropical North Atlantic Ocean

DEAN ROEMMICH* and CARL WUNSCHt

(Received 23 April 1984; accepted 6 August 1984; in final revised form 25 October 1984)

Abstract--Transatlantic hydrographic sections were obtained in mid-1981 along latitudes 24.5° and 36.25°N. The tracks nearly duplicated sections made 23 years earlier as part of the Inter- national Geophysical Year. A total of 215 stations were occupied: data from a con- ductivity-temperature-depth (CTD) probe and water samples for analyses of oxygen, nutrient, and other tracer concentrations were collected from the ocean surface to near bottom. The 1981 sections are described and displayed, and the circulation is compared to that of the earlier survey.

Large-scale meridional velocity and basin-integrated transport are compared in the 1981 and IGY sections, using a hierarchy of geostrophic models. In the simplest model, a reference level is based on the gross thermohaline flow and consideration of water mass characteristics, The only transport constraint is that the geostrophic plus Ekman flows sum to zero. Subsequent models impose mass conservation in a set of layers and then conservation of potential vorticity in a single layer. Distribu- tions of salinity and potential vorticity on density surfaces are examined in order to identify layers in which an advective balance of tracer distribution is plausible and where gradients are strong enough to be of practical use.

It was found that the 1981 and IGY sections have similar features in their large-scale velocity fields and similar zonally averaged meridional transport. In each case, net northward transports of approximately 17 Sv of surface and intermediate water (above 0, = 36.82) were balanced by equal southward flow in the deep water. A significant shift toward greater depth occurred in the depth distribution of the deep southward flow in the 1981 sections. The wind-driven subtropical gyre is superimposed on this thermohaline overturning, and transport calculations in the gyre interior show the Sverdrup relation to be of questionable direct applicability.

Ocean heat transport in 1981 was found to be about 1.2 x 10 Is W across 24°N and 0.8 x 1015 W across 36°N. These values are indistinguishable from those obtained from the IGY data and from computations of air-sea heat exchange. The steadiness of the heat transport is attributed to the invariance of the zonally averaged meridional circulation.

INTRODUCTION

DURING the summer of 1981, the R.V. Atlantis H occupied two transatlantic sections along nominal latitudes of 24.5 °N and 36.25 °N (Fig. 1). A total of 215 casts were made with a con- ductivity-temperature--depth (CTD) probe and a 24-bottle rosette water sampler. In all cases, data were collected from the ocean surface to within a few meters of the bottom. The CTD work was conducted by a group from the Woods Hole Oceanographic Institution (Dr. M. McCartney supervised the technical work). Almost all water samples were analyzed for salinity, oxygen and nutrients (the latter by a team under the supervision of Dr. L. Gordon of

* Scripps Institution of Oceanography, University of California, San Diego, La Jolla, CA 92093, U.S.A. t Center for Meteorology and Physical Oceanography, Department of Earth, Atmospheric, and Planetary

Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, U.S.A.

619

620 D. ROEMMICH and C. WUNSCH

Oregon State University) and, on selected casts, for alkalinity (Dr. P. Brewer) and tritium/helium-3 (Dr. W. Jenkins). On the 36°N leg, a number of special neodymium samples were obtained by D. Piepgras (PIEPGRAS and WASSERBURG, 1983).

Except for some deviations near the coasts (dictated by logistics and by political problems--a war---on the eastern end of the 24°N line), the ship tracks duplicated those of the so-called International Geophysical Year (IGY) sections displayed by FUGLISTER (1960). (The 24°N IGY section was obtained in October 1957 by the R.V. Discovery H and the 36°N section in April to May 1959 by the R.V. Chain.) The objectives of the 1981 cruises, in part, included a study of the long time-scale variability of the North Atlantic circulation by comparison of the two sets of data (IGY with 1981), including the meridional fluxes of mass, heat, and salt, to obtain the first detailed zonal sections of nutrients for combination with the previous dominantly meridional GEOSECS data and an attempt to understand the benefits, if any, of having available much higher density data than has been hitherto possible.

We believe that the 1981 sections have the highest spatial resolution of any transoceanic sections that have ever been run, i.e., with a horizontal separation approaching mesoscale resolution and a vertical resolution on the sub-meter scale. A criterion for adequate horizontal resolution in surveys such as these is that the mesoscale features not be aliased into long wavelengths that one would define as components of the gyre interior general circulation. We computed horizontal wavenumber spectra of temperature in the 1981 data and found, away from western boundaries, a fairly broad mesoscale peak at wavelengths of 200 to 300 km. The nominal station spacing of 80km over abyssal plains and 50 km or less elsewhere probably resolved most but not all of the eddy energy. In contrast, the IGY survey at 24°N had a nominal spacing of 185 km and the resulting spectrum clearly had considerable eddy energy folded into wavelengths of 1000 km and longer.

The purpose of the present paper is to describe briefly the sections and how they were obtained and to display them in a preliminary form for immediate use. We will then use the data and the antecedent IGY data to examine the meridional fluxes of mass and heat in the subtropical Atlantic and their possible changes. Some preliminary results describing the changes in apparent water mass volumes have been published by ROEMMICH and WUNSCH (1984); a discussion of the nutrient distributions and fluxes will also be published elsewhere.

THE SECTIONS

The 36°N section was run as Leg 1 of Atlantis H Cruise 109 (C. Wunsch, Chief Scientist), with the first station being obtained on 11 June 1981 just off Cape May, New Jersey, in 100 m of water. The last station was obtained on 9 July 1981 in 100 m of water southeast of Cape St. Vincent (see Fig. 1). A two-day break exists between Stas 23 and 24 when the ship was diverted to Bermuda to land an ill member of the scientific party. Leg 2 of cruise 109 was devoted to a "0-spiral triangle" cruise (H. Stommel) although, as part of this leg, a section was made nominally along the crest of the Mid-Atlantic Ridge providing data nearly normal to the two zonal sections. For Leg 3 (D. Roemmich, Chief Scientist) the ship departed 11 August 1981 from the Canary Islands, with the first station obtained in 105 m of water off Cape Juby, Morocco, on 12 August 1981. The last station, just east of the Bahama Bank, was concluded on 4 September. The ship then made two sections across the Florida Current at latitude 26°N02 ' from Great Isaac Rock, Bahamas, to Hollywood, Florida, and at latitude 27°N23 ' from Fort Pierce, Florida, to Mantanilla Shoal, Bahamas (Fig. 1), the last station being concluded on 6 September 1981.

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Ship tracks along 1981 and [GY transatlantic sections. Solid lines indicate IGY tracks, dashed lines 1981 tracks.

The CTD instruments used were those of Nell Brown Instrument Systems, Inc. Basic editing of the data bank took place at the Woods Hole Oceanographic Institution by proce- dures described by MILLARD (1982; private communication, 1983). The data were then reduced to 2 dbar averages and sent to MIT for final editing that consisted primarily of replacement of a number of bad surface salinity values and the rejection of outliers from the section average T-S curves. The number of corrected points was very small (in case of doubt, points were not replaced).

A 24-bottle rosette water sampler was used for oxygen and nutrient determinations. Typically, a full 24 samples were run for each station. Oxygen titrations were done onboard ship, primarily by M. McCartney (36°N) and L. V. Worthington (24°N). Standard depths were 10, 50, 100, 150, 200 m and increases of 100 to I000, 1200, 1500 m and increases by 500 to 5500 m and near bottom. These data were used to calibrate the CTD oxygens as described by MILLARD (1982). Nutrients were analyzed by Autoanalyzer and then edited under the supervision of L. I. Gordon.

Figures 2 to 4 display the sections: temperature, potential temperature, salinity, nitrate, phosphate, oxygen, and silica along 24°N, 36°N, and in the Straits of Florida, respectively. (We have not displayed the nitrite data.) The sections were drawn by use of an objective analysis method described by ROEMMICH (1983). (We are publishing these sections at this time in black and white to make the results available as soon as possible. However, it is hoped that they can be reproduced in color, along with a number of transatlantic sections obtained subsequently, in an atlas form comparable to that of FUGUSTER, 1960.)

The bottom topography displayed in these sections was read onboard ship at approximate 5-min (time) intervals. It compares well with that to be found in the IGY sections where the lines overlap (some depth-sounder-failure interpolations are readily identifiable in the IGY plates).

Because one of the goals of our work was to compare the 1981 sections with the IGY sections, it was decided to make an independent determination of the flow in the Florida Current to, at least, ascertain that during the 1981 period the flow was not abnormally large or small

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800i 80'~W 5(3' 40' 50' ' 79°'20'W


(e) 250 245 240

Om-

< 4 0 o ~

I00 &oo . ~

200.

300

400-

500

600 l

7001

8001_ 80°W 50' 40 30' 79"20'W

(f) 250

Om

I00

200

500

400.

5O0

600-

700-

800-

245 240

SILICATE

eo;w 5'0' 4b' ~b' 79020'w

Fig. 4. Sections of (a) temperature, (b) salinity, (c) dissolved nitrate-nirogen, (d) dissolved phosphate-phosphorous, (e) dissolved oxygen, and (f) dissolved silicate-silicon from Hollywood, Florida to Great Isaac Rock, Bahamas, along latitude 26 ° 02' in the Straits of Florida. Units are as in Fig. 2. Vertical exaggeration is 100:1. Station numbers are on the scale at the top and longitudes

on the scale at the bottom.


20

O v O O

"U

tt~ -20

-4C

Fig. 5.

Time -soon of

- 1 I I I l I I t 250 300 550 400 450 500 550

Days After I January 1981 00:00

Sea level difference between Bimini and Miami (Haulover Pier).

compared to its supposed true m e a n (NIILER and RICHARDSON, 1973). To do this inexpensively and conveniently, two pressure gauges (WuNSCH and DAHLEN, 1974) were placed on the Bahama Bank just south of Bimini for the period 4 June 1981 to l August 1982. Coupled with the existing tide gauges in the Miami area, it was believed, on the basis of previous work, that the sea level difference across the Straits would be a useful indicator of changes in the net transport (WUNSCH et al., 1969; ScHo'rr, private communication, 1983). With plans afoot to mount a full monitoring program in the period following the 1981 work, one could retrospectively calibrate the 1981 sea level difference measurement into a full transport value. In Fig. 5 wc display the Bimini-Miami (Haulover Pier) difference for these records (we predicted the tides for the raw records, subtracted them, and then filtered the residuals down to once/day values). The variability is known to be dominated by sea level fluctuations on the Miami side (WUNSCH et al., 1969). We see the expected (NIILER and I~CHARDSON, 1973) summer maximum and the autumn minimum. There is nothing in these records to suggest any anomalies in Florida Current transport although a definitive statement must await the backwards (in time) calibration effort underway.

VOLUME TRANSPORTS

In the remainder of this paper we will focus upon the zonally integrated meridional flow properties of the circulation, employing both the 1981 and IGY sections. Determination of absolute oceanic flows requires some form of model assumption. We will use a heirarchy of models of increasing complexity of assumption to analyze the flow field.

The simplest model assumes only that the ocean is in hydrostatic balance and that apart from the immediate vicinity of the sea surface where there is an Ekman layer, the flow is also in geostrophic equilibrium. The only additional assumption is that the net poleward flux of water across any zonal section is nearly zero. This single constraint does not permit estimates of the absolute velocity on small spatial scales (small being on the order of the station spacing). Nevertheless, the solutions are of interest because large-scale integral properties of the flow are already well-determined, particularly if the bottom is nearly flat. The near- rectangular shape of the 24°N section and the historical measurement program in the Straits of Florida were the basis of the heat flux estimate of BRYDEN and HALL (1980). Properties

640 D. ROEMMICHaBd C. WUNSCH

such as density and temperature are fairly strong functions of depth, so if the profile of transport per unit depth is well-determined, then good estimates may also be made of layer transports and fluxes of properties such as heat.

To proceed farther, e,g., to spatial scales smaller than that of the basin, additional constraints must be introduced. Later two additional sets of constraints will be considered. First, conservation of mass in a number of layers defined by isopycnal surfaces will be imposed. Here, the assumptions are that on large scales, transient storage and diapycnal advection and mixing are small. Then, we will examine the information content of additional conserved properties, such as potential vorticity, salinity, and nutrients. The addition of a second conserved property allows flow streamlines to be identified (as in the 13-spiral method of STOMMEL and SCHOTT, 1977) but requires an additional assumption of negligible mixing along isopycnal surfaces.

The sense of meridional geostrophic shear, taken over the entire width of the Atlantic, is such that the upper waters flow northward relative to the deep waters. This is a strong result demonstrated in many sections across both the North and South Atlantic (ROEMMICH, 1980; FU, 1981; BRYAN, 1982; WUNSCH and GRANT, 1982). Taking this observation together with the fact that little water flows in or out of the Atlantic at its northern boundary, one concludes that the absolute transport of surface and intermediate water is to the north and that of deep water is to the south. A near-zero total transport can be obtained with an average reference level in the vicinity of 1300 m. Salinity and nutrient distributions also indicate northward flow of intermediate water and southward flow of deep water (BROECKER et al., 1976, WUST, 1935). Low-salinity intermediate water extends northward in a tongue from its outcropping region in the Southern Ocean, and deep high-salinity water of North Atlantic origin is found in all oceans (REID and LYNN, 1971).

As a reference surface for geostrophic calculations, we chose the isopycnal 02 = 36.8, which approximately separates the fresh intermediate water of southern origin from the salty deep water of northern origin in the subtropical gyre at a mean depth of about 1300 m. In shallow areas where the bottom density is <(02 = 36.8, the ocean bottom was used as the reference surface. Changes in the reference surface did not greatly affect the results. Having chosen a reference surface, next a uniform velocity was applied to each section to satisfy the single constraint placed on total transport. In the Straits of Florida, the northward transport was required to be 30 Sv, in accord with direct measurements (NIILER and RICHARDSON, 1973) and with Fig. 5. At this latitude the transport in the interior, consisting of geostrophic transport plus Ekman transport, must balance the flow through the Straits. Northward Ekman transport across 24°N was estimated to be 6 +_ 2 Sv (with the error bar based on seasonal variations) from wind stress tabulations of HELLERMAN and ROSENSTEIN (1983). Thus, southward interior geostrophic transport must be about 36 Sv. At 36°N, Ekman transport was estimated to be 2 + 2 Sv to the south, and this amount was balanced by the total geostrophic transport. Thus in this lowest order model the only constraints used to compute the uniform reference level velocities were a simple balance between geostrophic and Ekman transports plus the measured transport through the Straits of Florida. The resulting zonal average profiles of geostrophic transport per unit depth are shown for 24°N (including the Florida Straits) and 36°N for both IGY and 1981 cruises (Fig. 6).

Profiles from different years are qualitatively similar though they differ in detail. Surface waters and intermediate waters flow northward above a zero-crossing at about 1300 m. The deep southward flows show a tendency to divide into two lobes, an upper lobe originating in the Labrador Sea and a deeper lobe in the Norwegian and Greenland seas. A second zero-


- 1 0 0 0

- 2 0 0 0 E v-3000

- 4 0 0 0

- 5 0 0 0

- 6 0 0 0

(a)

I I l I ~]

-0 .02 0.00 0. 2

Geoslrophic Transport per Unit Depth (106m2/s)

- 1 0 0 0

E - 2 0 0 0 .~[ 56ON

° - ooo ooo _ ooo _ _

I I 1 1

0.00 0.02

Geostrophic Transport per Unit Depth (106mZ/s)

Fig. 6. (a) Geostrophic transport per unit depth across 24*N including Straits of Florida, based on the reference level calculation with constraints on total transport only. The solid line is from 1981

data and the dashed line from IGY data. (b) Same as (a) but for 36"N.

crossing occurs in every section at depths from about 4500 to 5000 m above the poleward moving Antarctic Bottom Waters marked by low salinity and high silicate concentration.

To make a quantitative comparison, the ocean was divided into 18 layers bounded by isopycnal surfaces. The layers are defined in Table 1, which lists the transport in each layer across each of the four sections for this lowest order model. (Hereafter, 2 4 ° N is taken to include the Straits of Florida as well as the interior.) Differences between sections in the estimates of layer transports have a number of plausible explanations. These include: (1) errors in the estimation o f relative geostrophic velocity, i.e., data errors or interpolation errors (in general, data errors, in comparison to interpolation errors are small enough to be ignored); (2) errors resulting from the imposition of a uniform velocity at the reference level; (3) true time variability in transport. The hydrographic sections were separated in time by weeks to


Table 1. Layer transports (Sv)

Reference level calculation Inverse calculation with total with mass conservation

transport constraints in layers

Upper Lower 240N 36°N 24°N 360N Layer interface interface 1981 IGY 1981 IGY 1981 IGY 1981 IGY

I Ocean 00=26.4 12.2 13.4 11.3 13 .3 12.1 12 .7 11 .2 13.4 surface

2 00=26.4 =26.8 0.3 --4).9 -1.1 -3.2 0.4 -1.2 -1.7 -2.6 3 =26.8 =27.1 2.7 1.6 1.2 -1.5 3.3 1.4 1.4 -0.6 4 =27. I =27.3 1.5 3.0 1.6 0.4 2. I 2.8 1.9 1.0 5 =27.3 =27.5 1.4 1.4 1.4 1.7 2.2 1.5 1.6 2.1 6 =27.5 =27.7 0.1 --0.3 2.7 1.4 1.4 0.2 2.9 1.8 7 =27.7 02=36.82 -0.9 - I . I 1.0 0.8 -0.1 -0.7 1.8 1.4 8 05=36.82 =36.89 -I .3 -1.9 -0.3 -0.4 -0.6 -1.7 -0.6 -1.0 9 =36.89 =36.94 -1.2 -2.4 -0.4 -1.6 -0.6 -2.4 -0.5 -2.5

10 =36.94 =36.98 -! .3 -2.4 -I .2 - I .4 -0.5 -2.3 -1.2 -2.7 11 =36.98 =37.02 -1.7 -3.2 -0.5 -I .3 -1.0 -3.0 -0.4 -2.8 12 =37.02 o4=45.81 -1.1 -3.2 -0.5 -I .3 -0.4 -2.5 -0.3 -2.4 13 o4=45.81 =45.85 -4.3 -2.9 -0.9 -0.8 -I .8 -I .9 -1.5 -1.7 14 =45.85 =45.87 -I .4 -2.4 -3.4 0.8 -1.5 --0.3 -I .5 -0.3 15 =45.87 =45.895 -7.7 -2.6 -7.5 -3.1 -8.5 -2.8 -8.3 -2.7 16 =45.895 =45.91 -4.4 1.3 -3.5 -5.4 -5.9 -2.8 -5.4 -2.4 17 =45.91 =45.925 3.4 1.1 0.3 1.6 0.2 1.4 0.3 1.8 18 =45.925 Ocean 3.7 1.4 * * -0.3 1.5 * *

bottom

* No water.

decades. Variability on any time-scale in this range could cause changes in basin-integrated layer transport. (4) Convergence along isopycnais of the geostrophically computed flow. Such convergence may be indicative of cross-isopycnal flow, storage in layers, or ageostrophic divergence.

The layers may be grouped as surface water (layers 1 to 2), intermediate water (layers 3 to 6), deep water (layers 7 to 16), and bottom water (layers 17 to 18). Then, if we consider the four sections (two at 24°N and two at 36°N) as estimates of the same meridional circulation pattern, the mean northward transports are 11.3 Sv + 3.2 for surface water, 5.8 + 0.7 for intermediate water, -19.9 + 4 . 7 for deep water, and 2.9 + 2.6 for bottom water. The meridional circulation is strong and the standard deviations are much smaller than the means for all except the bottom water. To make further progress, we must begin to sort out the reasons for differences between the two latitudes or between the two realizations.

First, consider the two 24°N realizations. In the upper 14 layers, the r.m.s, difference in transport of individual layers was only 1.2 Sv between the IGY and 1981 data. The degree of similarity here is remarkable. Are the remaining small differences due to estimation errors or to time variability of the transport? Note that the differences were systematic as best seen in Fig. 7, which shows the transport per unit depth on a layer by layer basis, and indicates the mean depth of the top and bottom of each layer. The 1981 data showed consistently less net southward flow in layers 6 to 12, totaling 7.1 Sv less than the IGY. Instead, the southward flow of deep water was much more concentrated in the lower layers, 13 to 16, where it was 8.6 Sv greater than the comparable IGY transport.


O! (o) q i

;

L

- 2 0 0 20

Transport Density

, _ - 2 0 0 0 E

-4000

-6000

24 ° N -- -TGY ~ 1 9 8 1

I

410 60

106 m3/km s )

8kO

._. . -2000 E

- 4000

-6000

( b ) L - -

M

~ _ ~

56°N -- - - TGY - - 1 9 8 1

126

I I ,

- 2 0 0 20 40 6LO 80

Transport Density (106m3/krn s)

Fig. 7. (a) Layer transport per unit depth across 24°N, including Ekman transport in the upper layer, based on the reference calculation with constraints on total transport only. The solid line is from 1981 data and the dashed line from IGY data. Depths are the mean depth of the layer

interface. (b) Same as (a) but for 360N.

This pattern of change was repeated at 36°N. There, the net southward flow in layers 8 to 12 was diminished by 3.1 Sv in 1981 (layers 6 to 7 showed northward transport in both realizations) and that of layers 13 to 16 increased by 6.8 Sv. One can also see in Fig. 7 that some layers had substantially different average thickness in 1981 than in the IGY. For example, layer 15 was much thicker in 1981, while overlying layers were thinner. Differences in water mass volumes were discussed by ROEMMICH and WUNSCH (1984).

From the list of plausible reasons for observed transport variations given above, any of the first three might account for changes at a single latitude. But estimation errors are unlikely to produce such similar patterns of change at two very different latitudes. Similarly, mesoscale eddies with size of order 100 km are also unlikely to cause similar changes in net layer transport across two latitude circles separated by many eddy diameters. It is therefore probable

644 D. ROEMMICH and C. WUNSC8

that the observed changes are either related to some large-scale change in the meridional movement of deep water or are an artifact of the imposition of a uniform reference level velocity. If the reference velocity is the source of the layer transport change, then there must have been a corresponding temporal change in it; thus in any case large-scale velocity changes are required if not transport changes.

Another way of stating this situation is that basin-wide shear was substantially greater at levels around 3000 to 4000m in 1981 than during the IGY both at 24 ° and 36°N. This change is illustrated in Fig. 8, a plot of o 4 along the 3000m isobath at 240N. In 1981 the values ranged from about 45.82 at the western boundary to 45.81 at the eastern boundary. In the IGY data, these boundary values were about 45.79 and 45.81. The overall gradients are such that the southward geostrophic transport decreased with depth in the IGY data and increased with depth in the 1981 data, consistent with Fig. 7. At greater depths, the picture is even more complicated as isopycnals impinge on the Mid-Atlantic Ridge, disappear, and reemerge at depths that differ by hundreds of meters. The high degree of variability on short spatial scales makes it difficult to guess the time scale for such changes. However, the consistent shift to greater southward transport in the deep layers of both the 24 ° and 36°N sections suggest a long-term change.

We have argued that the small (order I Sv) differences in transport of individual layers between the two 24°N realizations were unlikely to be due to estimation errors. The accuracy of estimates of geostrophic transport has often been questioned, particularly in light o f a d hoc

assumptions that are often made to vertically extrapolate shear into the deep ocean near sloping topography. The procedure used here does not make any special assumptions near boundaries. Rather, a consistent objective mapping technique was applied over the entire region (RoEMMICH, 1983), Because the layers had a typical thickness of a few hundred meters and the sections were about 6000 km long, a transport of 1 Sv represents an average velocity of about 0.05 cm s -l , and this we take as an upper bound on basin-averaged errors in relative geostrophic velocity at any given depth.

Consider now the latitude dependence of transport in the two realizations. Subtraction of layer transports across 24°N from those across 36°N gave estimates of divergence in the

Fig. 8.

45.82 -

E 0 o o 4 5 . 8 0 - t

'* ~1 b

I

4 5 . 7 8 / I L

0

,'M /\ Ix /l~J .r~/),

24 ° N - - - TGY - - 1 9 8 1

i J 40'00 L • 2000 6000

D i s t a n c e ( k i n )

Values o f o 4 along the 3000-m isobath at 240N from 1981 data (solid line) and IGY data (dashed line).


bounded region. Both the 1981 and IGY realizations showed convergence in all the upper layers, divergence in intermediate and deep layers, and convergence in bottom layers. There was no water in layer 18 (04 > 45.925) at 36°N. The upper convergence was 3.8 Sv in layers 1 to 3 of the 1981 data and 8.1 Sv in layers 1 to 4 of the IGY data. The divergence was 11.7 Sv in layers 4 to 13 of the 1981 data and 16.3 Sv in layers 5 to 14 of the IGY data, and the lower convergence was 8.1 Sv in layers 14 to 18 of the 1981 data and 8.0 Sv in layers 15 to 18 of the IGY data.

Because of the high degree of similarity between the 1981 and IGY patterns of divergence, it is unlikely that they were produced by mapping errors. Instead, candidate explanations are that the divergence patterns are a spurious result of the uniform reference level velocities or that there is a significant imbalance in the geostrophic transport of the layers. The apparent divergences are large enough that storage or removal on the 23 year time scale of the surveys may be ruled out. Large ageostrophic divergences also seem unlikely in the relatively sluggish deep ocean. Perhaps along-isopycnal divergence is balanced by flow across isopycnals. But the implied vertical velocities, of order 10 -4 cm s -l, seem too high on the large spatial scale involved. Finally, it may be the case that the true reference level velocity was spatially distributed in such a way as to largely cancel the apparent divergences (Table I). Two simple ways of reducing the apparent divergences in the 1981 realization were examined. (1) The Florida Straits transport has considerable time variability (Fig. 5). Noting that the base of layer 4, o0 = 27.3, is close to the density of the densest water in the Florida Straits, the apparent divergences can be reduced by lowering the northward transport through the Florida Straits to 26 Sv with an equivalent reduction in southward interior transport across 24°N. This Florida Straits transport is within the measured range of transports (Fig. 5). In fact, the 1981 transport in the Florida Straits relative to a bottom reference level was about 6 Sv less than in the IGY section. A uniform reference level velocity is maintained in each section, and the upper convergence in layers 1 to 3 is thereby reduced to < lS v . The divergence in layers 4 to 13 is reduced from 11.7 to 7.7 Sv. (2) A single cold core ring or meander at 36°N reduced the total thickness of layers 1 to 3 by about 400 m over a distance of 100 km with a compensating 400 m increase in the thickness of layers 4 to 13. A 15 cm s southward velocity in this feature, with northward transport elsewhere in the section to maintain a balance of total mass, would eliminate the apparent upper convergence while halving the mid-level divergence.

A third alternative is to spread the correction out over large scales, avoiding large barotropic velocities in small regions. We have accomplished this balancing by means of an inverse method (WUNSCH, 1978) to yield a model of second order. Approximate mass conservation constraints were imposed in each of the lower 16 layers, which do not come into contact with the atmosphere within the area bounded by 24 ° and 36°N. The constraints on total transport used in the simple calculation were retained. Solutions trade off minimization of the sum of squares of reference level velocities against minimization of the sum of squares of residuals in the constraint equations. For details refer to WUNSCH (1978) or ROEMMICH (1981).

Layer transports from the inverse solutions are shown in Table 1, along with transports from the reference level calculation with total transport constraints. Initial imbalances in individual layers were small and the adjustments were likewise small (of order 1 Sv). However, the larger systematic imbalances over groups of layers were greatly reduced. For example, the 11.7 Sv divergence in layers 4 to 13 in 1981 data was reduced to 3.0 Sv. Also, the 3.7 Sv imbalance in 1981 layer 18 was reduced to 0.3 Sv. In the four sections, northward


transport of surface and intermediate waters ranged from 15.1 to 21.5 Sv and an equal amount of deep water flowed southward. As was found in the simple model, the main difference between the IGY and 1981 results was a shifting of the main southward transport from layers 6 to 12 in the IGY to layers 13 to 16 in 1981.

Even further reduction in the divergences is possible by pushing the layer residuals to smaller values. This would occur at the price of yet more structure in the velocity field at the original reference level. Because such structure cannot be ruled out, the magnitude of divergences remains highly uncertain.

GEOSTROPHIC VELOCITY

Here we examine the velocity fields that result from the basic model. Corrections to these fields in the inverse solution were of order 0.1 cm s -I and did not qualitatively change the structure. The velocities were smoothed to remove mesoscale eddies and display the large- scale velocity. The smoothing was done with a horizontal Gaussian low-pass filter, arbitrarily chosen to decay to e -~ of its central value at a distance of 500 kin. The filter was normalized to have unit area when integrated to infinity. Because the ocean has boundaries, the filter has less than unit area applied over the bounded region, and thus does not conserve transport. The distinction is slight in mid-ocean but reduces transport immediately adjacent to the boundary by 50%. We have made this choice rather than normalizing the finite filter to unit area, and thus conserving transport, to avoid an artificial appearance of boundary-trapped flow. Here we are concerned with the patterns of flow rather than the transport.

Figure 9 shows the 1981 smoothed geostrophic velocity at 24 ° and 36°N, respectively. At the western side of the 36°N section, the smoothed Gulf Stream flows northward in the upper 1400 m. Southward flow extends across the entire width of the ocean, though it is deepest and strongest in the west, near the Stream. At 24°N, most of the western boundary current is within the Straits of Florida, but a substantial Antilles Current extends well out into the western basin. Again at 24°N the surface southward flow is ocean-wide, but is deepest and strongest in the west. At about 1000 m depth, a weak northward flow is seen at both latitudes, in each case with a lobe on both sides of the Mid-Atlantic Ridge.

Deep flows in the two sections seem bound to topography and exhibit strong coherence between one section and the other. The southward-flowing deep western boundary current is seen in both sections but has a much stronger maximum at 36°N. Below and slightly to the east of the southward flow, a band of northward flow is seen at each latitude, with very similar structure in the two sections. The remainder of the deep circulation is weak, but note that, excluding the flows at the western boundary, the eastern and western basins both have anticyclonic circulations at 26 ° and at 36°N.

It was mentioned that the net meridionai flows were northward in layers with water having characteristics of southern origin and southward in layers of North Atlantic origin. On smaller scales, it is also the case that some of the boundary flows transport properties away from their source regions. In Fig. 10, salinity and geostrophic velocity at 24°N are mapped on an ordinate of 04, and the bottom values of 04 are greater by about 0.04 in the western basin than in the eastern basin, though the depths of the two basins are nearly equal. Again referring to Fig. 10, the ocean is shallower at the location of the deep northward-flowing western boundary current than at the location of the larger northward flow farther east: nevertheless, the northward-flowing water at the western boundary carries water as dense and as fresh as the flow farther east. The bottom, which is nearly flat, shows a pronounced depression in 04, with the northward velocities and low salinity in the bowl of this depression. The


O m .

IO00-

2000-

3000-

4000

5000

6000 ~

7 t W

(o) 230 220 210 200 190 180 170 160 1,55

..... I ' " ' I " " I ' " ' I . . . . I ' . . . . . . . . . . I . . . . I . . . . . . . I . . . . I ' ' I . . . . I . . . . I . . . . I"

/ \ . . . . . . . . I . . . . I . . . . I . . . . I . . . . ~ . . . . . . . . ~;__°N Velocity +

7 0 ° 6 0 ° 5 0 ° 4 0 * 3 0 * 2 0 ° I~W

(b)

5 I0 20 30 40 50 60 70 80 90 I00 " P " ' P " ' I ' " ' I " ~ ' I ' ' ' l ' ' ' I ~ I I ' l ' I ' ' l . . . . i . . . . I ' " ' l " ' ' l " " l . . . . I . . . . I . . . . I . . . . I . . . . l " r

Om. : .................... . . . .

1000 1

2000.

3000 .

4 0 0 0 -

5 0 0 0 .

700W 6 0 ° 50 ° 40* 30 ° 20 ° IOOW

Fig. 9. (a) Horizontal ly smoothed geostrophic velocity (cm s -~) at 2 4 ° N f rom 1981 data, based on the reference level calculation with constraints on total transport only. Southward velocities are

shaded. The western end is on the left. (b) Same as (a) for 36°N .

bottom topography of the Mid-Atlantic Ridge appears less rough in the a4 domain because the vertical gradient of a4 is small at the depth of the bottom features. The depression containing water of low salinity also has high values of silicate and low oxygen. In Fig. 11, values of potential temperature, salinity, silicate, and oxygen are shown for the deepest water samples at Stas 222 to 228; the depths have a total range of <80 m, and vertical differences in the properties over this depth range are small compared to the observed horizontal differences. Data from the 'bowl' in Fig. 10 appear at two stations, 225 and 226. The point is not that


(a) CT4= . . . . . . . . . i . . . . . . . . . i . . . . . . . . . i . . . . . . . . . I . . . . . . . . . , . . . . . . . . . ~ r ~

v ~ ~ ~ ~ _ v ' 3 4 . 9 B ~ . L . _ L . ~ I ] ~ ' - -34.96

.92 ' ~ 45 .85 ~ 1

45.90 I 4 5 9 5 n . . . . . . . . . I . . . . . . . . . 1 . . . . . . . . . I . . . . . . . . . I . . . . . . . . 2 ~1 . . . .

• OKm I 0 0 0 2000 3000 4000 5 0 0 0 6000

(b)

(74= " . . . . . . . . ' . . . . . . . . . ~ . . . . . . . . . ' . . . . . . . . . . ' . . . . . . . . . ' . . . . . . . . . . ' . . . . i

45.75 0 0 O -0.2 -0.2 -0.2

q2 / 45901 ~ 4595 JK; . . . . . . . ,0 o ' . . . . . . . . . ' . . . . . . . . . ' . . . . . . . . . ' 2000 3000 4000 5000 6000

FiB. 10. (a) Con(our map of deep salinity at 24°N (1981 data) with o 4 as ordinate. (b) Contour map of deep geostrophic velocity at 240N (1981 data) with a, as ordinate.

there is an anomaly in ®/S or other correlations but rather that water lower on the correlation curves appears at these stations.

As another example, consider the weak northward velocity at the eastern boundary at 900 m in the 24°N section (Fig. 9). A map of salinity near the eastern boundary (Fig. 12a) shows a salinity minimum at the same location. The IGY 24°N section intersected the eastern boundary several hundred kilometers farther south than the 1981 section (Fig. 1). Smoothed velocity and salinity near the eastern boundary from the IGY 24°N section are shown in Fig. 12. The nor thward current is again present, but much stronger, and the salinity minimum is more intense. The salinity of 35.03 is the lowest of any intermediate water east of the Mid- Atlantic Ridge at that latitude. The IGY section from 32°N also shows a salinity minimum at

Circulation and heat flux in the North Atlantic Ocean 6 4 9

1.8

W iI I---

u..l

~ 1.7 b- _/ < p- z W

o ¢1 1.6

- 34.89. i-

.88

- Z, .B7

if)

.86

- .B5 ~-

.84 228

o A x " +

o8 x S A $i + 0z

X

x '~' +

6 ×

+

226 224

STATION NUMBER

1 42 6O

- - 46 o - 'z / Jl x

50 m 5.9 °

+ -._j5 8

54

1 I 5 B . 222

Fig. I 1. Measured values of potential temperature, salinity, silicate, and oxygen from the deepest samples at Stas 222 to 228. The depth of samples varied by <80 m and differences are due mainly to

horizontal gradients.

(a)

t [ I ' I

STA. 165 160 0m

155

I000

200C

3 0 0 C

5 0 0 0 W m - 5 5 0 0 6 0 0 0

650 D. ROEMMICH and C, WUNSCH

(b)

Om

IOOC

200£

:~OOC

Om

I000

2000

3 0 0 0

5000 (c)

I

5500 6000 km

I I

, , I I 1 5000 5500 6 0 0 0 km

Fig. 12. (a) Contour map of salinity near the eastern boundary of the 24°N section in 1981. This section intersected the coast near 28°N (see Fig. 1). (b) Contour map of salinity near the eastern boundary of the IGY 24°N section, (c) Contour map of smoothed geostrophic velocity near the

eastern boundary of the I GY 24 o N section.

Circulation and heat flux in the North Atlantic Ocean 6 51

the eastern boundary at 800 m, though it does not appear at 16 ° or 36°N. One interpretation is that the eastward flow of Antarctic Intermediate Water in the tropics feeds an eastern boundary current that flows into the Mediterranean salt tongue.

DOES T H E S V E R D R U P R E L A T I O N A C C O U N T FOR THE M I D - A T L A N T I C C I R C U L A T I O N ?

The solutions discussed in the previous section were weakly constrained. The equations for conservation of mass in layers were imposed globally, that is in the form of integrals around the entire enclosed region. Here, and in the next section, we investigate a potentially more powerful set of constraints, obtained by appending a second conserved variable. In the absence of mixing, the second variable allows, in principle, fluid parcels in a density layer to be 'tagged' according to the value of the variable. If unique identification can be made with the available data, streamlines of flow may then be determined, as they represent the intersection of surfaces of constant density with surfaces of constant value of the second variable. We choose potential vorticity as the second variable, although other conserved quantities such as salinity may also be used. Conservation of potential vorticity together with density forms the basis of the 13-spiral method (STOMMEL and SCHOTT, 1977) and other theoretical studies of the wind-driven circulation (NEEDLER, 1972; RHINES and YOUNG, 1982; LurrEN et al., 1983), and a set of maps of potential vorticity on density surfaces in the North Atlantic was produced by McDOWELL et al. (1982). One must determine whether gradients of potential vorticity are large enough to make constraints of this type practical, and whether mixing is strong enough to preclude the identification of streamlines from values of potential vorticity.

In the model of LOYTEN et al. (1983) and other wind-driven ocean models, potential vorticity in the subtropical gyre is acquired at the ocean surface under the action of the wind stress curl. Ekman convergence forces water down along isopycnal surfaces into the interior. On leaving the surface, the potential vorticity of a fluid parcel remains constant as long as the parcel is in the interior region. Only layers so ventilated, which outcrop in the winter south of the line of zero wind-stress curl, are thus set in motion. The integral of transport over all layers in motion is specified by the Sverdrup relation. Motions driven by the large-scale thermohaline forces (convection) are ignored.

The reference level used in calculations in the previous section was chosen on the basis of global scale thermohaline constraints, but the Sverdrup relation was not built into the solution. It may be useful to think of the net transport across a line of latitude as being composed of three components: (1) the Ekman flux, (2) the wind-driven, or Sverdrup geostrophic interior, and (3) a thermohaline component of the interior geostrophic flow, driven by high- latitude convective processes, but confined to the western boundary by vorticity dynamics. The division between these two latter categories is arbitrary, but perhaps still useful.

Considerable confusion has come to be associated with the expression "Sverdrup transport", with many writers associating it with any quasi-geostrophic flow satisfying the linear vorticity balance

~w fDv=f D---~--' (1)

even when the source of divergence driving the flow may well be at the bottom of the ocean rather than the surface wind-stress curl as in Sverdrup's original formulation. The integral of


(1) from some level Zo to the surface gives

13 vdz f k . p V x - - f - fw 0 , (2)

where the vertical velocity, w, at the base of the Ekman layer is equal to the Ekman divergence given by the curl of the wind stress. The Sverdrup relationship, as originally defined, would required that there be some level z 0 where w 0 = 0. Unless such a zo(x, y) were a simple function of position, one would feel compelled to reject the hypothesis that the Sverdrup relationship dominates the overall flow field. LUrrEN et al. (1983) pointed out that density of outcropping water along the line of zero wind stress curl seems to be about a0 = 27.4 in winter. Does the observed transport of water above ae = 27.4 agree with a prediction based on measurements of wind stress? (The title of this section has been taken from a paper of that name by LEETMA et al., 1977.) In Fig. 13, we have plotted the meridional geostrophic transport of water above a0 = 27.4, using velocities from the simple model. Transport distribution in the inverse solutions was not substantially different. Transport was set to 0 at the eastern boundary and the integration proceeded westward. The total geostrophic transport should be equal to Sverdrup transport minus the ageostrophic Ekman transport. The Ekman transport was computed from the wind stress and the Sverdrup transport from the wind stress curl, using values of mean annual wind stress in 2 ° by 2 ° squares, compiled by HELLERMAN and ROSENSTEIN (1983). The difference, Sverdrup transport minus Ekman transport, is shown in Fig. 13.

At 24°N, the observed geostrophic transport and the estimate from wind stress are seen to be in excellent agreement over the eastern 4000 km of the ocean. A similar conclusion, using only the IGY data, was reported by LEETMAA et al. (1977). However, one can see in Fig. 13 that, integrated all the way to the western boundary, the northward geostrophic transport above as = 27.4 was -19 Sv and was nearly identical in the IGY and 1981 data. To balance this flux, there was 6 Sv of northward Ekman flow in the interior and 30 Sv of northward geostrophic flow in the Florida Straits, giving a net poleward flow above a0 = 27.4 of 17 Sv, which might be defined as the upper thermohaline component. Since the flow in the ocean interior appears to be completely accounted for by the wind, the thermohaline flow occurs in the western boundary region (which is, however, considerably broader than the Gulf Stream).

This picture is quite different from that of LEETMAA et al. (1977) who balanced the poleward moving 30 Sv of the Florida Current against a net equatorward moving interior of equal magnitude and opposite sign above about 1000 m, yielding an ocean circulation that closed above that surface. Their balance does not account for the thermohaline contribution (convective overturning contribution) and would lead to a very small poleward heat flux if it were the totality of the flow field. It also leaves the movement of deep water unspecified.

At 36°N (Fig. 13b), except for a large eddy in the IGY data, the 1981 and IGY transports are again very similar. But the slope of a regression line over the eastern 4000 km of geostrophic transport is in each case less than the slope of the transport estimate from wind stress. The transport per unit width of the observed geostrophic circulation is only about -2.5 m 2 s -~ while the wind-driven estimate is about -4.5 m 2 s -~. It does not appear that this discrepancy can be eliminated by hypothesizing a southward velocity at the reference level. To enhance the southward transport above a 0 = 27.4 by the required amount, a reference velocity of about -0.3 cm s -] is necessary over the whole eastern basin. But this would produce a very large southward transport in the deep water east of the Mid-Atlantic Ridge, inconsistent with the sluggish deep transport in the eastern basin at 24°N.


- 1 0 O3

= - 2 0 0 Q.

c

- 3 0 I -

- 4 0

(o)

I

0

20

~"// \/'~t

/ o

J 1981

I I I

20~00 4000 60100

Distence (krn)

l:Ilbl I A

1 0 I r i

t, I 36 ° N

- 4 0 v V

- 5 0 I I I t

0 20~00 40~00 60JO0

Oistonce (km)

Fig. 13. (a) Geostrophic transport above a 0 = 27.4, integrated from the eastern boundary westward at 24°N, for 1981 data (thin solid line) and IGY data (dashed line), based on the reference level calculation with constraints on total transport only. The heavy solid line is wind-driven geostrophic transport computed from mean annual wind stress tabulated by HELLERMAN and ROSENSTEIN

0983). (b) Same as (a) but for 36*N.

Our present results, depicted in Fig. 13, perhaps suggest that the Sverdrup relationship does apply from 2000 to 6000 km east of the western boundary at 24°N. But unless it can be demonstrated that w 0 = 0 at the reference isopycnal, we may have no more than a coincidence Lthe LEETMAA et al. (1977) result may be only a coincidence]. Given the large bottom slopes present in the real ocean and the reality of deep flows derived from the thermohaline circulation that will lead to divergence terms, it might be thought implausible for the Sverdrup relation to apply. Charts of dynamic topography (100 db relative to 1500 db) and wind-driven geostrophic transport in the North Atlantic were prepared by LEETMAA and BUNKER (1978) . If these are overlain, one can see that in the eastern basin at the northern rim of the subtropical gyre, the contour lines intersect at steep angles. This is equivalent to the disparity between observed geostrophic transport and wind-driven transport noted above. At


about 36°N, in the eastern basin, the dynamic topography of the 100 dbar surface relative to 1500 dbar has isopleths that slant from northwest to southeast. Streamlines of Sverdrup transport, on the other hand, slant from northeast to southwest. This disparity is discussed further in the next section. As at 24°N, we identify the net northward transport of about 15 Sv across 36°N as the thermohaline component of the circulation.

D I S T R I B U T I O N OF P O T E N T I A L V O R T I C I T Y AND S A L I N I T Y

Vertical sections of smoothed potential vorticity in the 1981 data are presented in Fig. 14. The ordinate is a 0, ranging from 26.0 to 27.6. Relative vorticity was neglected, and potential vorticity was calculated over 2 grid points in the vertical (80 m) as

f Aa0 no Az"

We have used the same Gaussian filter as previously but applied the filter along lines of constant potential density and normalized it to unit area when integrated over the bounded ocean.

The prominent minimum in potential vorticity at about no = 26.5 is subtropical mode water, which outcrops on the seaward edge of the Gulf Stream in winter. This water and all water of lower density are exposed to atmospheric forcing within the area bounded by 24 ° and 36°N, so these fluid parcels are not expected to conserve their potential vorticity between these two latitudes. Parcels of water that are insulated from the ocean surface conserve their potential vorticity in the absence of mixing.

Between a o = 26.5 and 27.0 is a layer described by McDOWELL et aL (1982) as having uniform potential vorticity over most of the subtropical gyre. It is true that the gradient in potential vorticity in this layer is smaller within the gyre than at its edges. Nevertheless the gradient is distinguishable from zero. Figure 15 shows the unsmoothed potential vorticity along a 0 = 26.75 at 24°N, both from the 1981 and the IGY data. The 24°N section is in the interior of the gyre. The potential vorticity varies by about 2096 of its value over a distance of 3000 km with a minimum in mid-ocean. This variation is greater than the small-scale 'wiggliness' and is statistically non-zero. There remains the question of the dynamical sig- nificance of this zonal gradient. In magnitude, it is about 25% of the size of the ambient meridional gradient

13 A~0 ~0 Az"

The 36°N section (Fig. 15) has a small gradient in the western half and two very steep gradient regions in the east, again consistent with the smoothed Fig. 6 of McDOWELL et al.

(1982). The steep gradient regions, or fronts, are approximately at the same location as the subtropical front described by KASE and SIEDLER (1982).

Scaling assumptions required to obtain the Sverdrup balance do not hold in strong frontal regions. Furthermore, we cannot draw streamlines connecting the low values of potential vorticity found at the eastern end of 24°N with the much higher values at the eastern end of 36°N. Rathcr, the potential vorticity distribution in this layer suggests that streamlines must slant northwest to southeast in the eastern basin, as indicated by the dynamic topography (LEETMAA and BUNKER, 1978), and in contrast to streamlines of Sverdrup transport.

Proceeding still deeper, to a0 = 27.4 one can see in Fig. 14 that at 24°N the potential vortieity decreases monotonically from west to east with no minimum, while at 36°N there is


(a) o-o= 26.0 ~ ° o o ~ ~ o o

"" ~ 3 0 0 , 500

26.5 ~ 80

tD

27. 5 _ ~ / ~ ~ . - - - - - - - ~ ~ ao ~ 1 !

Okra I000 2000 3000 4000 5000 6000 km

60

(b) (--

Okrn

Fig. 14.

I000 ~000 5000 4000 5000 6000krn

(a) Contour map of smoothed potential vorticity at 24°N (1981 data) with o 0 as ordinate. (b) Same as (a) but for 36°N.


200i

T E

i~ 150

~ 100 E

0 EL

• ( a )

24 °N ~I cr e = 26.75

198 I(smoothed)~

I

198 V v ~ ~ r '

"Vn ~-~-G Y \I

I I I l

0 20100 40AO0 60100

Distance (km)

600 f i o I z t cre= 26.75 ~ t J

S~ 400 i981-

200 I ,, / -6 ,...~!/~ -' - ,',~TGY

0 I I I 1 1 1 I

0 2000 4000 6000

Distance (km)

Fig. 15. (a) Potential vorticity on o o = 26.75 at 24°N for unsmoothed IGY data (dashed line), unsmoothed 1981 data (,thin solid line), and smoothed 1981 data (heavy solid line). (b) Potential vorticity on o 0 = 26.75 at 36°N for unsmoothed IGY data (dashed line) and unsmoothed 1981 data

(solid line).

still a minimum, but it is shifted far east. At this level, the lines of constant potential vorticity slant from northeast to southwest, with the value at the western boundary at 24°N being found about 2500 km from the eastern boundary at 36°N. But geostrophic velocities at this level are much weaker than above and the transport here contributes little to the interior transport of the gyre.

It is interesting to compare the distribution of potential vorticity with that of salinity, For steady flow with no mixing, contours of the two tracers along density surfaces are parallel to each other and to streamlines. Figure 16 shows contours of salinity, with the same smoothing applied as in the field of potential vorticity. Features of the salinity distribution are shallow maxima at 3 6 ° N in the subtropical mode water and at 24°N in the subtropical underwater (WUST, 1964), a shallow water mass formed locally by high evaporation. Deeper, one sees the salinity minimum of Antarctic Intermediate water at the western boundary at 36°N and in


26.0

26.5

27.0

(a)

I .... I .... I ' " I .... I .... I .... I .... I .... I .... I .... I .... I .... I .... |

I 3 6 . 0 0 II

Okra 1000 L~O00 3000 4000 5000 6000 km

(b} C ~ = l . . . . I . . . . I . . . . . . I . . . . I . . . . I . . . . I . . . . . . I . . . . . . I . . . . I . . . . I . . . . I . . . .

' °Iylll///Ifl, ,. 2 6 , 5

3~00

27.0

2?.5

0 km

Fig. 16.

35.50 ~

" ~ ' - ' ~ - ~ - 3 5 . 2 0 - . ~ /

. I . . . . I . . . . I . . . . I . . . . I . . . . I . . . . 1 . . . . 1 . . . . I . . . . 1 . . . . d = ~ . ~ k _ ~ , ,

I000 2000 3000 4000 5000 6000 km

(a) Contour map of smoothed salinity at 24°N (1981 data) with c e as ordinate. (b) Same as (a) but for 36°N.

658 D. ROEMMtCH and C. WUSSCH

the interior in the 24°N section. At the eastern end of 36°N is the top of the Mediterranean salinity maximum. We have sketched isolines of salinity and potential vorticity between 24 ° and 36 o N for a range of densities, and found that they are roughly parallel for densities from o e -- 26.5 to 27.2. For example, on oe = 27.0, salinity contours in the range 35.7 to 35.8 slant from northeast to southwest as do the potential vorticity contours 100 and 120 that are located at about the same relative positions. Deeper, this parallelism is not found. On oe = 27.6 the salinity contours slant from northwest to southeast while those of potential vorticity slant from northeast to southwest. The implication is that at this level, below the wind-driven gyre, advection by the mean flow is very weak and mixing may be important.

If potential vorticity, or salinity, is to be used as a conservative tracer to constrain the geostrophic velocity fields, apparently the part of the water column where there is useful information is quite limited. Near the ocean surface, there are sources and sinks. Without good estimates of the strength of the sources and sinks, we cannot impose conservation constraints there. On the other hand, at depth the interior geostrophic flow seems to be so sluggish that a model should include mixing. At a range of densities, however, from about o e = 27.0 to 27.4, there are strong flows in the interior and the parallelism of the tracer fields suggest that a purely advective model is plausible.

A solution was therefore constrained as follows. First, the mass conservation equations were de-emphasized by reducing the number of layers to four. The bottoms of these layers were the surfaces o 0 --- 26.2, 27.0, 27.4, and the ocean bottom. In the third layer, four equations for conservation of potential vorticity were introduced. Values of smoothed potential vorticity in Fig. 14 were used to identify five hypothetical streamlines on the Oe = 27.2 surface at 24 ° and 36°N. It can be seen in Fig. 14 that there is a substantial zonal gradient in this layer and the variation is monotonic, decreasing from west to east, except for a minimum near the eastern boundary at 36°N. Because the contour lines are nearly vertical, the location of values on the density surface can be done fairly accurately, i.e., potential vorticity and density are nearly orthogonal so substantially independent information is contained in the two fields. The equations were then requirements that the transport in layer 3 between a particular pair of streamlines was the same at the two latitudes. Let xl(v) and x2~,) be the positions of two streamlines as a function of latitude. Then

V d x dz(24ON) ~- ¢LT dZt36ON). o0=27.0 x~ *'o0 = 27.0".v ~ p

Layer transports in the resulting solutions were nearly identical to solutions which used only the layer conservation equations. However, there was some horizontal redistribution of transport within layers due to the added constraints. This redistribution can best be seen in contour plots of the smoothed velocity. Figure 17 is of smoothed velocity from the potential vorticity conserving solution and should be compared with Fig. 9. The solutions which constrained total transport in layers do not appear appreciably different from Fig. 9. In the vorticity conserving solutions, the reference level has become much less discernible and contour lines are now seen to cross that level. The flow field appears more barotropic, particularly at 36°N where several features are seen to extend from surface to bottom. A particularly interesting example is the nearly barotropic flow in the same direction as the Gulf Stream located on the seaward edge of the Stream with a reversal farther offshore. SCHMIrz (1980) has described flows of weak depth-dependence at 55°W, where there was one such flow in the direction of the Gulf Stream but about 200 to 300 km seaward and a second with


(o) 250 ?_20 210 2 0 0 190 180 170 160 155 , I . . , I , , , , I JN , I , , , , i J * J , . , ~ , 1 , , , ~ 1 , ~ , ~ p , , , i n , , , I V , ~ v l , , , , I ~v~ , I ~ . , ] i , , , , I , ~ , , i , .

Om

I 0 0 0 1

ooo 2 \ 7 oooo L

6 0 0 0 ! . . . . . , . . . . , . . . . . . . . . . . . . . , . . . . . , , , , , . . . . , , , , , . . . . , . . . . , . . . . , ,

75*W 70 ° 6 0 ° 50 " 4 0 ° 30 ° 20 ° 15°W

(b) 40 50 ' ' i . . . . I .... I"

I 0 0 0 -

8 0 0 0 -

3 0 0 0

4000

5000

5 I0 20 50 .... P"'I'" ' I . . . . I . . . . I . . . . I '

,0.2 ~ 0 . 2

6 0 70 8 0 9 0 I O0 ' I . . . . I . . . . I . . . . I . . . . I . . . . I . . . . I . . . . I . . . . l '"r

3 6 ° N Ve loc i t y Ill I ~ ~

, , I , , , , l , , , , I , , , , I , , , , I , , , , I , , , , i , , , , I , , , , I , , , , ] , , , , I , , , , I , , r , l r

7 & W 6 0 ° 5 0 ° 4 0 o 5 0 ° 2 0 ° IO*W

Fig. 17. (a) Contour map of smoothed geostrophic velocity at 24°N from the inverse solution with mass conservation constraints in layers and potential vorticity conserved between a 0 = 27.0 and

o 0 = 27.4. (b) Same as (a) but for 36°N.

opposite direction farther offshore. A comparison of Figs 9 and 17 shows that some features have been strengthened and others reduced, but the major features described previously remain in both. The interior flow at 36°N still does not obey the Sverdrup balance. In fact, because the solution velocities were northward, the discrepancy was actually increased some- what.


OCEAN HEAT TRANSPORT

We have seen that the two realizations of subtropical gyre hydrography produced very similar estimates of net meridional transport of water masses. We will examine the implica- tions of such similarity for time variability of ocean heat transport and for our ability to measure mean heat transport from single, 'snapshot', realizations. An estimate of northern hemisphere heat transport (OORT and VON DER HAAR, 1976), made by taking the residual of the Earth's radiation budget and the heat budget of the atmosphere, together with measured heat storage in the ocean, showed a large annual variation of 4.9 x l0 ~5 W at 24°N. By examining the mechanisms of ocean heat transport, we will find that annual variation must be <20% of that amount.

Table 2 summarizes the components of the heat flux, as computed from the 1981 data at 24°N. The western boundary current is taken to contain the 30 Sv of flow in the Florida Straits plus the 10 Sv flowing northward in the upper layers just east of the Bahamas. The transport-weighted average temperature of this water is

v0 - - = 18.3oC.

We cannot distinguish between the temperature of wind-driven and thermohaline contribu- tions to the western boundary current, so we will assume that they are the same. This assumption adds a small uncertainty to the decomposition of Table 2. The northward Ekman transport is assigned the average surface temperature, about 25°C, across 24°N. The net heat transport by the wind-driven flow is only 0.1 x I0 I~ W. It is small because temperature con- trasts in the upper layers are small across the gyre, i.e., northward and southward flows have nearly the same temperature. On the other hand, the heat transport by the vertical meridional circulation is much greater (1.1 x 1015W) than for the horizontal recirculation because of the large temperature contrast (15.4°C) between the upper layer northward flow and the deep southward flow.

Given that most of the heat is transported by the vertical meridional circulation, how large can the annual variability be? The deep western boundary current at 24°N is some 6000 km downstream from the region where deep water is formed. With an advection speed of 10 cm s -z this translates to a time lag of about two years from the source to 24°N. It seems unlikely that a seasonal signal at the source regions could persist so far downstream. Indeed, the sections span three seasons (spring, summer, and autumn), and the lack of significant differences in the total southward flow of deep water indicates that there is no seasonal signal in transport greater than the noise level. The other possibility for a change in heat flux by the vertical meridional circulation is if the northward-flowing surface component of the flow (which we have said is in the western boundary current) changed temperature seasonally. There are seasonal changes in the temperature structure in the Florida Straits (NIILER and RICHARDSON, 1973), but the transport weighted average temperature, as deduced from Niiller and Richardson's Fig. 14, remains nearly constant.

We also can estimate the magnitude of heat flux variations in the horizontal recirculation. By varying the magnitude of recirculation by 4 Sv, in accord with the seasonal signal in the Florida Straits (Fig. 5), or the heat content of the upper 200 m in the ocean interior by about 1.3 x 109 J m -2, as reported by BRYAN and SCHROEDER (1960), a seasonal cycle of 0.1 to 0.2 × 10 J5 W is calculated, depending on the assumed temperature balancing the Florida Straits variations and the vertical distribution of the changes in heat content. In the former case the


Table 2. Heat flux components across 24°N

Transport vO,tF .f.f Cp vOdxdz (Sv) (°C) (x 10t~W)

Wind-driven components Northward geostrophic flow in

western boundary current

Northward Ekman flow in interior

Southward geostrophic flow in interior above o 0 = 27.4

Heat flux in wind-driven circulation

Thermohaline components Northward flow in western

boundary current

Southward deep flow

Heat flux in thermohaline circulation

Total heat flow

23 18.3 1.8

6 25 0.6

-29 18.7 -2.3

0.1

17 18.3 1.3

-17 2.9 -0.2

1.1

1.2

maximum heat transport would occur in early summer, in the latter in winter. However, the heat flux error in a single realization was estimated as 0.3 x 1015 W by BRYDEN and HALL (1980) and WtmSCH (1980), therefore, realizations in different seasons are not expected to produce significantly different results, and this was the case.

Global ocean heat flux at 24°N is the sum of Atlantic and Pacific heat fluxes. There is no extant transpacific section at this latitude. However, it is known that there is no deep water formed in the North Pacific (WARREN, 1981), SO the vertical meridional heat flux is probably much less there than in the Atlantic. This is supported by TALLEY's (1984) finding, based on integration of air-sea heat transfer, that the meridional heat flux at this latitude in the Pacific is not significantly different from zero. Any seasonal signal is once again probably due to the gyre circulation. It again can be argued, on the basis of the strength of the circulation and changes in heat content, that the seasonal variations at 24°N are <0(10 ~5) W. GmL and N.LER (1973) showed in a scaling argument that on large spatial scales at annual period, advection of heat is much smaller than storage, which is equivalent to our conclusion. PORT and VAN DER HAAR'S (1976) residual calculation implied that advection of heat by ocean currents at annual period was of the same order as storage, but advection in the ocean interior is simply not strong enough to carry water parcels a long distance in a season.

Table 3 lists the computed heat fluxes across both 24 ° and 36°N for the IGY and 1981 data for the simple reference level calculation and the inverse solution. The solution with potential vorticity conservation was not different from the one in which only total mass was conserved in layers. The heat fluxes at each latitude are indistinguishable betweeen realizations because the vertical meridional cell did not change strength in 23 years. The values at both latitudes are consistent with integrated air-sea fluxes computed by BUNKER (1984). The decrease in northward heat flux between 24 ° and 36°N is not due to a decrease in volume transport of the vertical meridional cell between those two latitudes. Rather it is caused by the cooling of the Gulf Stream waters and the consequent decrease in the temperature difference between the northward-flowing surface water and the southward-flowing deep water.


Table 3. Heat flux estimates

Reference level calculation with transport constraints

Inverse calculation with mass conservation in layers

240N 36~N 240N 360N 1981 IGY 1981 IGY 1981 IGY 1981 IGY

Oceanic heat flux 1.2 1.2 0.8 0.5 1.2 1.1 0.8 0.7 (10 a5 W)

SUMMARY

We have analyzed two realizations of the hydrography of the subtropical North Atlantic Ocean, separated in time by 23 years. A sequence of calculations was carried out, beginning with a simple reference level model containing constraints on total transport. Subsequent calculations included more assumptions and resulted in a more refined determination of the flow. An inverse method was used to construct a solution obeying mass conservation in density layers, and then a solution with mass conservation in density layers plus conservation of potential vorticity in one layer was found. The mass conservation constraints served to determine net layer transports across each section while the potential vorticity constraints added information on the horizontal distribution of transport within the layers.

Gross features of the zonally averaged meridional circulation were very similar in each of the calculations for both the 1981 and IGY realizations. About 17 Sv of surface and intermediate water flowed northward above a2 = 36.82, with an equal southward transport below this surface. However, there was a substantial change in shear between 3000 and 4000 m. As a result, the 1981 realization had approximately 6 Sv less southward flow betweeen o2 = 36.82 and o4 = 45.81 (roughly 1300 to 3000 m) and correspondingly more southward flow between o4 = 45.81 and o4 = 45.91 (roughly 3000 to 5000 m).

Contoured sections of geostrophic velocity, horizontally smoothed to suppress features smaller than 1000 km, were qualitatively similar in the 1981 and IGY realizations at 24 ° and 36°N. Southward recirculation of the western boundary current occurred in an ocean-wide band extending approximately to the isopycnal surface o 0 = 27.4. The recirculation was deepest and strongest in the west, near the boundary current. The 1981 data showed at both latitudes two deep western boundary currents, a strong southward flow at the western boundary over a depth range of about 1600 to 4400 m at 36°N and a weaker flow at about 3000 to 5200 m at 24°N. In each case, the flow direction was reversed just below and offshore of the southward current. This structure was also present in the IGY 360N section but not at 24°N. Other deep flows appeared bound to topography and there was a tendency for deep anticyclonic flow in both basins in both realizations. Examples of boundary flows carrying tracer signals away from their source regions were the deep northward-flowing western boundary current and an eastern boundary current carrying Antarctic Intermediate Water across 24°N.

Horizontal structure of the transport above o 0 = 27.4 was compared with predictions of wind-driven circulation models. A western boundary region with high variability and intensified reeirculation was found to extend about 2000 km into each section. The interior region then consisted of the eastern 4000 km where the variability was much less. The Sverdrup relationship appeared to be of doubtful direct applicability.


In preparation for the potential vorticity conserving model, the distribution of salinity and

potential vorticity on density surfaces was studied. No layer with ocean-scale uniform poten-

tial vorticity was found. Parallelism in the distribution of two tracers identified layers in which

a model neglecting horizontal mixing is plausible, Conversely, it was seen that at tt 0 = 27.6 intersections of the salinity and potential vorticity fields imply that mixing is of the same order as advective terms.

Northward transport of heat in the ocean was found to be about 1.2 x 1015W across 24°N and 0.8 x 1015W across 36°N. These values were in agreement with previous studies and

those of the 1957 to 1959 data and the 1981 data were indistinguishable. The high degree

of similarity was attributed to the time invariance of the gross zonally averaged meridional circulation. A decomposition of the heat flux into thermohaline and wind-driven components

at 2 4 ° N showed that only about- 10% could be attributed to the horizontally recirculating

wind gyre. Transatlantic sections from different years and from both hemispheres have

produced very consistent descriptions of the meridional heat flux because of the time invariance of the meridional cell, together with the fact that the meridional cell is responsible

for the bulk of the heat transport.

Acknowledgements--CTD data were collected by the Woods Hole CTD group under the supervision of Dr. M. McCartney. Nutrient analyses were supervised by Dr. L. Gordon of Oregon State University. We greatly appreciate the cooperation of Captain Jorgensen and the crew of the R.V. Atlantis I1. The work was supported by the National Science Foundation under Grant OCE-80-185140 to MIT and Grant OCE-81-21262 to the University of California, San Diego.

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two transatlantic sections: meridional circulation and heat flux in the subtropical north atlantic...

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