Ocean heat transport across 24°N in the Pacific

Deep-Sea Research,Vol. 38, No. 3, pp. 297-324, 1991.

0198-0149/91 $3.00 + 0.00 © 1991 Pergamon Press plc

Printed in Great Britain.

Ocean heat transport across 24°N in the Pacific HARRY L. BRYDEN,* DEAN H. ROEMMICHt a n d JOHN A . CHURCH,:

(Received 7 June 1987; in revised form 31 August 1990; accepted 11 September 1990) Almtraet---4)ceanheat transport across 24°N in the North Pacific is estimated to be 0.76 x 1015W northward from the 1985 transpacific hydrographic section. This northward heat transport is due half to a zonally averaged, vertical meridional circulation cell and half to a horizontal circulation cell. The vertical meridional cell is a shallow one, in which the northward Ekman transport of warm surface waters returns southward only slightly deeper and colder, all within the upper 700 m of the water column. In terms of its meridional heat transport, the horizontal circulation cell is also shallow with effectivelyall of its northward heat transport in the upper 700 m of the water column. Previous estimates of North Pacific heat transport at subtropical latitudes had ranged between 1.14 x 10Is W northward and 1.17 x 1015W southward. The error in this new direct estimate of Pacific heat transport is approximately 0.3 x 1015W. In addition, it is suggested that the annual variation in poleward heat transport across 24"N in the Pacific is of order 0.2 x 1015W, as long as the deep circulation below 1000m exhibits little variation in water mass transport. Together, the Pacific and Atlantic transoceanic sections essentially close off the global ocean north of 24"N so that the total ocean heat transport across 24°N is estimated to be 2.0 x 1015W northward. This ocean heat transport is larger than the northward atmospheric energy transport across 24"N of 1.7 x 10Is W. The ocean and atmosphere together transport 3.7 x 1015W of heat across 24"N, which is in reasonable agreement with classic values of 4.0 x 1015W derived from consideration of the Earth's radiation budget but which is markedly less than the 5.3 x 1015W required by recent satellite radiation budget determinations.

INTRODUCTION THE a m o u n t of h e a t t r a n s p o r t e d b y the N o r t h Pacific O c e a n is c o n t r o v e r s i a l . T h e single direct e s t i m a t e by BRYAN (1962) y i e l d e d a s o u t h w a r d h e a t t r a n s p o r t across 32°N of 1.17 x 1015 W . Such e q u a t o r w a r d h e a t t r a n s p o r t is c o u n t e r i n t u i t i v e since the s u b p o l a r o c e a n is e x p e c t e d to lose h e a t to the a t m o s p h e r e , n o t t o g a i n h e a t . E a r l y e s t i m a t e s of a i r sea h e a t e x c h a n g e !ndicated a n e t heat loss of 0.41 x lOX5W b y t h e Pacific O c e a n n o r t h of 30°N (EMIG, 1967). T h e m o s t r e c e n t calculations o f a i r - s e a e x c h a n g e , which are p r e s u m ably m o s t r e l i a b l e , h o w e v e r , result in a n e t gain of h e a t by t h e o c e a n n o r t h of 300N of 0.19 x 1015 W ( C l a r k , as r e p o r t e d b y TALLEY, 1984). I n d e e d , t h e a i r - s e a h e a t e x c h a n g e s for the Pacific, as s u m m a r i z e d b y TALLEY (1984), are d i s c o u r a g i n g l y u n s e t t l e d , yielding o c e a n h e a t t r a n s p o r t s across 300N that vary f r o m 1.14 x 1015 W n o r t h w a r d (HASTENRATH, 1980) to C l a r k ' s s o u t h w a r d t r a n s p o r t of 0.19 x 1015 W . F r o m existing direct a n d i n d i r e c t

*Woods Hole Oceanographic Institution, Woods Hole, MA 02543, U.S.A. t Scripps Institution of Oceanography, La Jolla, CA 92093, U.S.A. :~CSIRO Division of Oceanography, Marine Laboratories, Hobart, Tasmania, Australia. 297

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estimates then, the northward heat transport across 30°N in the Pacific Ocean might lie anywhere between 1.14 and - 1.17 × 1015 W. To understand the role of the ocean circulation in maintaining the global heat budget, the North Pacific heat transport is a key unknown. From the radiation budget at the top of the atmosphere based on recent satellite measurements, the atmosphere and ocean together must transport 5.3 × 1015 W of heat poleward across 25°N (STEPHENSe t a l . , 1981). Meteorologists seem certain that the atmosphere transports 1.7 × 1015 W poleward across 25°N (OoRT and VONDER HAAR, 1976; CARISSlUO et al., 1985). The meridional heat transport across 25°N in the Atlantic Ocean has settled at a value of 1.2 × 1015 W, based on good agreement between many direct estimates using old and new measurements (HALL and BRVDEN, 1982; ROEMMICH and WuNscn, 1985) and indirect estimates based on BtmKER'S (1976) air--sea heat exchange calculations. In the absence of a direct estimate of North Pacific heat transport, we would then be led to believe that the ocean currents must transport 2.3 x 1015W across 25°N in the Pacific, an amount far larger than any of the existing estimates and an amount incompatible with the absence of significant volumes of cold imermediate or deep waters formed in the North Pacific. Clearly, a modern direct estimate of ocean heat transport across 25°N in the Pacific is n e e d e d to help diagnose the mechanisms by which the Earth's heat balance is maintained. In spring 1985, a transpacific hydrographic section was carried out along 24°15'N aboard R.V. T h o m p s o n to provide the measurements necessary to make an estimate of ocean heat transport across 24°N in the Pacific. RoEr,tl~IcH et al. (1991) have described the water mass characteristics of the 24°N transpacific section, and ROEMMICrt and MCCALLISTER (1989) have used the 24°N section in a diagnostic inverse calculation of the North Pacific circulation. This hydrographic section was designed to complement the existing transatlantic sections and meridional heat transport estimates for the Atlantic Ocean at 24°N. Because there is essentially no Indian Ocean north of 24°N, the transpacific and transatlantic sections close off the world ocean north of 24°N. The sum of their poleward heat transports then yields the total ocean heat transport across 24°N which can be directly compared with estimates of the atmospheric energy transport into the polar cap north of 24°N in order to assess the roles of oceanic and atmospheric heat transport in maintaining the global radiation budget. DATA The transpacific section consists of 216 CTD stations taken between 30 March (San Diego) and 4 June 1985 (Nagasaki). Each cast was made to the bottom or to 6300 m, if the bottom was deeper than 6300 m, and generally, 36 water samples were collected throughout the water column and analysed for salinity, oxygen and nutrients. The cruise track was predominantly zonal along 24°15'N with diagonal legs at the eastern and western sides and across the Hawaiian Ridge in order to cross major topographic features perpendicular to the isobaths (Fig. 1). Stations 1-369" define the large-scale, mid-ocean circulation between San Diego and the Ryukyu Islands. They are conveniently divided into the

*While there are 216 CTD stations, they are numbered approximatelyalternately because expendablecurrent profiler (XCP) stations were generally made halfwaybetween CTD stations and the XCPs were assigned station numbers. NIILERet al. (1991) have described the XCP measurements along 24°N

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eastern section (Stas 1-152), the central section (Stas 154-305), and the western section (Stas 306--369) by the Hawaii Ridge and the Izu-Ogasawara Ridge (Fig. 2). Two sections were then made across the Kuroshio as it flows northward over the Okinawa Trough west of the Ryukyus (Stas 370-389) and as it flows eastward through the Tokara Strait north of the Ryukyus (Stas 390-408). In contrast to the Atlantic where the Gulf Stream transport through Florida Straits is known from independent direct current observations, the Kuroshio transport west of the Ryukyu Islands or through the Tokara Strait has not been measured. Because the warm Kuroshio flow makes an important contribution to the poleward ocean heat transport, acoustic Doppler velocity measurements were carried out along the transpacific section to help define the Kuroshio transport. These Doppler current measurements are also useful for defining the flows over relatively shallow topography associated with the two ridges and continental slopes, which the transpacific section crossed, since such flows may also contribute substantially to the estimate of ocean heat transport. Acoustic Doppler velocity profiling is a relatively new measurement technique. JoYCEet al. (1986) have demonstrated the efficacy of using acoustic Doppler velocities as reference level velocities in inverse calculations across short transects of the Gulf Stream. Because of the need for direct velocity measurements in the Kuroshio, the Doppler velocity measurements were included on board the R.V. T h o m p s o n for the 24°N transpacific section. CHURCH and JoYcE (personal communication) have shown that the Dopper velocities at 100 m depth are strongly correlated with geostrophic velocities at 100 m relative to 1000 m depth perpendicular to the section. Such comparisons are dominated by mesoscale eddy variability with alternating northward and southward currents of order 20 cm s-1. For the heat transport estimation, however, small biases in Doppler velocities can lead to large imbalances in the net transport across the 24°N section. For example, Church and Joyce made several tests for the relative horizontal orientation of the instrument alignment on board the T h o m p s o n , which indicated that the orientation was correct within +0.3 °. At a ship speed of 5 m s -1, an uncertainty of 0.3 ° in alignment results in an uncertainty in velocity perpendicular to the cruise track of 2.6 cm s-1. A bias of 2.6 cm s-1 in perpendicular velocity for the 12,000 km, 24°N section yields a bias in meridional transport of 1500 × ]06 m 3 s -1. Naively using the Doppler velocity measurements as reference level velocities might therefore cause substantial problems. To determine the bias in Doppler velocity measurements, geostrophic velocities for the transpacific section are referenced to the 100 m Doppler velocities and the geostrophic transport across 24°N is added to the Ekman transport. Under the assumption that this total transport should be zero, the bias in perpendicular (northward) Doppler velocity is calculated to be -1.44 cm s -1, which for a 5 m s -1 ship speed can result from an error in instrument alignment of 0.16 °, well within Church and Joyce's independent determination of +0.3 ° . It is not clear whether such a bias remains constant over the entire transpacific section. The ship's logging of satellite navigation data failed about halfway across the section at 153°E so that LORAN navigation is used for the western half of the section and satellite navigation for the eastern half. At the extreme western region west of 127°E near the Ryukyu Islands where the LORAN navigation coverage is very poor, Church and Joyce were compelled to examine each acoustic Doppler profile carefully for errors. It is common practice for ship's officers to make small corrections to the gyro compass during a long cruise, but no such adjustments were noted during the transpacific section. Thus, the

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Fig. 3. Scatter plot of Doppler velocities measured at 100 m depth vs geostrophic velocities at 100 m depth referenced to zero velocity at the bottom. The values have been averaged over approximately three stations pairs (200 km and 24 h) by applying a Hanning filter twice to the individual values. The line of one-to-one correspondence is drawn. The slope of the principal axis line is 0.99 and the correlation is 0.89.

resulting set of Doppler velocity measurements is uneven: east of 153°E, satellite navigation assures good data quality; from 153° to 127°E the LORAN navigation coverage is adequate; west of 127°E, the Doppler velocity measurements are carefully but subjectively edited. How to use the Doppler velocities for referencing geostrophic calculations is an active research area. The correlation between geostrophic velocities at 100 m referenced to zero velocity at the bottom and Doppler velocities at 100 m averaged over the distance between each CTD station pair is 0.68; and the root mean square (r.m.s.) difference between the two sets of velocities is 12.7 cm s-z. Using Doppler velocities as reference level velocities for geostrophic calculations would then result in large variability in the deep velocity field, with a variance of 160 cm 2 s -2 which is an order of magnitude greater than the variance of order 10 cm 2 s -2 measured by current meters at 4000 m depth in the North Pacific (ScHurrz, 1988). Thus, referencing geostrophic calculations directly to the Doppler velocities is not reasonable. Averaging the velocities over three station pairs, which is both a spatial average over about 200 km and a temporal average over about 24 h, increases the correlation between geostrophic and Doppler velocities to 0.89 and decreases their r.m.s. difference to 4.6 cm s -1 (Fig. 3). Averaging then improves the agreement between the velocities, but their difference still implies a deep velocity variance which is too large compared with direct obse~ations of the deep velocity field. Doppler velocities, however,

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provide valuable information where there is an uncertainty in the reference level, such as in regions of flow over relatively shallow bottom topography. In this work, we use the Doppler velocities to reference geostrophic velocities for all station pairs where the bottom depth is less than 1000 m. For such shallow stations, the estimated uncertainty of + 13 cm s-1 in Doppler velocity measurements is probably less than the uncertainty in assuming zero velocity at the bottom. Doppler velocities are used as reference level velocities for a total of 12 station pairs over the California shelf and slope, over the Hawaii and Izu-Ogasawara Ridges, near Okinawa and on the slope and shelf of the East China Sea. For all D o p p l e r velocities used here, the velocity bias of 1.44 cm s-1 determined above is subtracted from the perpendicular (northward) velocities processed by CHURCH and JOYCE (personal communication). COMPONENTS

OF PACIFIC OCEAN

HEAT TRANSPORT

As described by HALL and BRYDEN (1982), there are three principal components in an estimate of ocean heat transport for the North Pacific 24°N section: the wind-driven Ekman layer flow in the near-surface waters; the strong western boundary current flow of the Kuroshio west of the Ryukyu Islands; and the mid-ocean geostrophic flow between North America and the Ryukyu Islands. Each of these components is considered in sequence. To assure mass balance across the section which is essential for the heat transport estimate, the mid-ocean geostrophic flow is made to have a southward transport equal to the sum of the northward E k m a n and Kuroshio transports.

Ekman layerflow Wind stress values tabulated by HELLERMAN and ROSENSTEIN (1983) are used to determine the northward, wind-driven E k m a n layer transport, -frX/pfdx, where r x is the eastward wind stress, p is density and f i s the Coriolis parameter. Their zonally averaged wind stress at 24.25°N has an annual average of - 0 . 5 9 9 dyn cm -z and an April-May average of -0.689. Although the cruise track deviated from a zonal one, particularly at the eastern and western ends, wind stresses parallel to the track are virtually identical to the 24.25°N values. Wind stresses parallel to the cruise track from Hellerman and Rosenstein are integrated along the cruise track to determine the northward Ekman layer transport perpendicular to the track of 1 2 . 0 x l 0 6 m 3 s -1 for the annual average and 13.8 x 106 m 3 s -1 for the April-May values. To calculate an average temperature for the Ekman layer transport, the wind-driven flow is assumed to decrease linearly from its surface value to zero at 50 m depth so that its temperature, 0e, is an average of the temperatures at 0, 20 and 40 m depth: 0E = 00 × 0.35 × 020 x 0.48 + 040 X 0.17. The product of wind stress times 0e integrated along the cruise track yields an average temperature for the Ekman layer transport of 22.56°C (Table 1). Although this transpacific section was done only once, we treat the ocean heat transport estimate as an annual average value. Thus, we summarize the E k m a n layer component as a northward flow of 12.0 × 106 m 3 s -1 at an average temperature of 22.56°C. Compared with the Atlantic section at 24°N, the Pacific E k m a n layer transport is about 2.4 times larger than the Atlantic, reflecting primarily the greater width of the Pacific, but the Pacific surface layer temperatures are about 3°C colder. A discussion of the ocean heat transport estimate as an April-May average is presented in the section on annual variability.

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Ekman layer transport and its temperature

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April-May average

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-0.599 -0.595

-0.689 -0.683

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Western boundary currentflow At 24°N in the Atlantic, the bottom velocity of the Gulf Stream flow through Florida Straits is about 11 cm s -x (BROOKS and NIILER, 1977). If there were such a bottom flow in the Kuroshio, geostrophic velocities referenced to zero at the bottom would underestimate the Kuroshio transport arid its contribution to meridional heat transport. For this reason, the acoustic Doppler velocity measurements were included on the 24°N section to reference the geostrophic velocities in the Kuroshio. Because the 24°N transpacific section crosses the Kuroshio where it flows over deep water (1700 m maximum depth), whereas the sill depth for the Kuroshio downstream in Tokara Strait is much shallower (of order 600 m), large velocities in the waters below 1000 m depth are not expected on this section. Indeed, the Doppler velocities and geostrophic velocities relative to the bottom are in good agreement, particularly over the deep water portions (Table 2). Such agreement gives added confidence in the Doppler velocity measurements. The differences between Stas 372 and 376 may be due to different positions recorded for the Doppler measurements and the CTD Sta. 374. It is notable that the average velocities over the distance between Stas 372 and 376 agree within 4.3 cm s-t. For the station pairs where the bottom depth is shallower than 1000 m, the Doppler Table 2.

Geostrophic velocities relative to the bottom and acoustic Doppler velocities in the Kuroshio

Stations

Bottom depth (m)

370-372 372-374 374-376 376-378 378-380 380-382 382-384 384-385 385-386 386-387 387-389

320 1480 1660 1550 1240 845 255 135 110 100 80

Oeostrophic velocity at 70 m depth (cm s -1) relative to zero velocity at the bottom - 100.1 0.8 80.9 89.5 86.1 107.1 38.5 18.5 -5.6 3.0 2.3

Acoustic Doppler velocity at 70 m depth (cm s-1) normal to the section -54.8 27.4 62.0 83.6 84.7 103.0 42.2 44.5 20.0 - 1.4 -4.7

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velocities are probably more reliable measures of the Kuroshio velocity since there is no reason to suppose the flow at the bottom is zero over such shallow regions. To determine the Kuroshio transport over Okinawa Trough, the D o p p l e r velocities perpendicular to the ship track at 70 m depth are used as reference level velocities for the geostrophic velocity profiles for station pairs where the bottom depth is less than 1000 m. For station pairs with bottom depths greater than 1000 m the bottom velocity is assumed to be zero. The resulting Kuroshio transport is 28.3 x 106 m 3 s -1 with an average temperature of 18.5°C. Such values are remarkably similar to the Gulf Stream flow through Florida Straits of 29.5 × 106 m 3 s -1 with an average temperature of 19.7°C. In the procedure of HALL and BRVOEN (1982), the contribution of the Kuroshio to the northward heat transport can be divided into a baroclinic heat flux of 0.85 x 1015 W and a barotropic flow of 28.3 x 106 m 3 s- 1 at an average temperature of 11.1 °C (Table 3). The colder barotropic temperature and larger baroclinic heat flux relative to the Gulf Stream values in the Atlantic are both due to the greater depth of the section across the Kuroshio. Because WORXmNGrON and KAWAI (1972) suggested there was a nearly equal contribution to the total Kuroshio flow from the eastern side of the R y u k y u Islands so that their total Kuroshio transport for the Ryukyu section was 59 x 106m 3 s -1, the geostrophic velocities are referenced to the D o p p l e r velocities for the continental slope and shelf region east of Okinawa where the bottom is shallower than 4000 m depth. The total transport for this short section between Stas 357 and 369 is only 4.4 x 106 m 3 s - 1 and has an average temperature of 4.930C, not very different from the mid-ocean average temperature of 3.560C. A strong northward flow further east between Stas 341 and 349 appears to

Table 3.

Net heat transport across 24°N in the Pacific Northward transport ( x 106 m 3 s - l )

E k m a n layer 12.0 Kuroshio 28.3 Mid-ocean geostrophic - 40.3 Net northward heat transport = 0.76 x 1015 W

Method 1

Method 2

T e m p e r a t u r e (°C) 22.56 18.46 15.11

Component heat flux

Transport ( x 106 m 3 s -1)

Temperature difference (°C)

Northward heat transport (1015 W)

E k m a n layer Kuroshio

12.0 28.3

22.56 - 15.11 = 7.45 18.46 - 15.11 = 3.35

0.37 0.39 Net = 0.76 x 1015 W

Baroclinic heat flux (× 1015W)

Transport ( x l 0 6 m 3 s -1)

Barotropic temperature difference (*C)

Barotropic heat flux (x1015W)

E k m a n layer 12.0 22.56 - 3.56 = 19.00 0.93 Kuroshio 0.85 28.3 11.12 - 3.56 = 7.56 0.88 Mid-ocean geostrophic - 1.91 40.3 Net baroclinic -- - 1 . 0 6 x 1015 W Net barotropic = 1.81 x 1015 W Net northward heat transport = 0.76 x 1015 W

Ocean heat transport in the Pacific

307

be half an eddy circulation, with a nearly equal southward flow between Stas 328 and 341 (Fig. 2). Thus, the supposed Kuroshio flow along the eastern side of the Ryukyus appears to be as elusive as the fabled Antilles Current east of the Bahamas in the Atlantic (OLSON et al., 1984). The second section across the Kuroshio as it flows through Tokara Strait north of the Ryukyu Islands is difficult to interpret. The baroclinic shear in the CTD section is weak and the Doppler velocity measurements are wildly varying, perhaps because the L O R A N navigation was nearly unusable there. Did the Tokara Strait section fail to measure a sizeable branch of the Kuroshio in the Osumi Strait just south of Kyushu? Or was there a substantial flow into the Sea of Japan in the Tsushima Current at the time? Because of such unresolved questions, this section is not used for the heat tranport estimate.

Mid-ocean geostrophic flow Following the analysis of ROEMt,tlCH and MCCALLISrER(1989), who chose an initial reference surface along the e2 = 36.9 isopycnal which is near 2000 m depth, we initially calculated geostrophic velocity profiles for each station pair using a reference level of 2000 dbar where the velocity is assumed to be zero. For station pairs where the depth is less than 1000 m, the Doppler velocity at 100 m is used to reference the geostrophic profiles. For station pairs where the depth is greater than 1000 m but less than the reference level, the bottom velocity is assumed to be zero. Mid-ocean geostrophic transport for Stas 1-369 (from California to Okinawa) referenced to 2000 dbar in this way is 22.0 x 106 m 3 s -1 southward. Adding a southward velocity of -0.032 cm s -1 uniformly over the entire section generates a total southward transport of 40.3 x 106 m3 s-1 to compensate exactly for the combined northward Ekman and Kuroshio transports. This southward mid-ocean geostrophic transport occurs at an average temperature of 11.67°C. There are several difficulties with the resulting mid-ocean circulation, however. First, there is a net southward transport of 15 x 106m 3 s -1 of deep water colder than 2.5°C across the entire section. For the North Pacific where deep and bottom waters are not formed (WARREN, 1983), such a large net southward transport of deep water is unreasonable. Secondly, in the western portion of the 24°N section across the Philippine Basin between Sta. 306 at the Izu-Ogasawara Ridge and Sta. 369 at Okinawa, there is a substantial southward transport of deep water relative to the reference level of 2000 dbar. ROE~tMICHet al. (1990) argued that the Philippine Basin is essentially closed off north of the 24°N section for depths greater than 2500 m. Even before the barotropic southward current is added, the southward transport referenced to 2000 dbar for Stas 306-369 is 8 × 106 m 3 s-1 below 2500 m depth. Such a large transport is untenable. Since it results from the penetration of northward geostrophic shear below 2000 dbar in the Philippine Basin, a deeper reference level is called for. Because of the difficulties in the circulation implied by the 2000 dbar reference level, we decided to choose a deeper basic reference level of 3000 dbar since: (a) it is a standard reference level; (b) the geostrophic transport referenced to 3000 dbar in the Philippine Basin is only 0.4 × 106m3s -1 southward below 2500m depth; and (c) 3000dbar is approximately in the middle of the silicate maximum water in the central and eastern basins (ROEMMICHet al., 1991, Plate 5), which is the oldest and presumably most stagnant water type on the 24°N section. For a basic reference level of 3000 dbar, the mid-ocean southward geostrophic transport across 24°N is 30.4 x 106 m 3 s-1. A barotropic velocity of

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--0.019 cm s -1 is then added uniformly to the central and eastern portions of the section in order to make the southward mid-ocean geostrophic transport exactly balance the combined northward Ekman and Kuroshio transport of 40.3 x 106 m 3 s -~. No barotropic southward velocity is added to the western portion of the section across the Philippine Basin so that no additional southward transport is added below 2500 m depth where the basin is closed to the north. The mid-ocean geostrophic flow determined in this way has southward transport of 40.3 x 106 m 3 s-1 at an average temperature of 15:11 °C. A total of 21.6 x 106 m 3 s-1 flows southward geostrophically across the eastern portion of the section between North America and the Hawaii Ridge; 27.4 x 106 m 3 s -1 flows southward across the central section between the Hawaii and Izu-Ogasawara Ridges, and 8.8 x 106m 3 s -I flows northward across the western portion of the section across the Philippine Basin between the Izu-Ogasawara Ridge and Okinawa. Over the complete mid-ocean section, there is a small net southward transport of 2.4 x 106 m 3 s -1 of deep water colder than 2.5°C. The vertical distribution of the mid-ocean meridional circulation is discussed more thoroughly later in this paper.

NET HEAT TRANSPORT ACROSS 24°N IN THE PACIFIC The Ekman layer, western boundary current and mid-ocean gcostrophic components can be easily combined to determine the net heat transport across the transpacific 24°N section. The northward, wind-driven transport of 12.0 × 106 m 3 s -1 at an average temperature of 22.56°C and the northward Kuroshio transport of 28.3 x 106 m 3 s -1 at an average temperature of 18.46°C are balanced by the southward mid-ocean return transport of 40.3 x 106 m 3 s-1 at an average temperature of 15.11°C. Since the mass transports are in balance, the northward heat transport is then directly estimated to be 0.76 x 1015 W (Table 3). RoEt,iUlCtt and MCCALLISTER(1989) used this same 24°N hydrographic section as part of the data input for an inverse calculation of the large-scale circulation in the North Pacific Ocean. As part of the inverse solution, they estimated a similar northward heat transport across the 24°N section of 0.75 x 1015 W, but deferred discussion of the heat transport mechanisms to this present work. It is of interest that changing the details of the calculation procedure as to reference levels or to the constraints of other sections in the inversion has little effect on the estimate of meridional heat transport across 24°N. There are several ways to diagnose the mechanisms by which the ocean transports this heat northward and much of the confusion in the literature over the importance of various mechanisms really resides in the definitions used. In the most simple procedure, we might define the E k m a n component of the heat transport to equal the Ekman transport times the difference in temperature between the wind-driven surface layer and the southward return flow: 12.0 x 106m 3s -I x ( 2 2 . 5 6 - 15.11°C) x 4.1 x 1 0 6 j ° C - l m -3 = 0.37 x 1015 W. The western boundary current component would then be defined to equal the Kuroshio transport times the difference in temperature between the Kuroshio and southward return flow: 28.3 x 106m3s -1 x (18.46 - 15.11°C) x 4.1 x 106j °C-1 m -3 = 0.39 x 1015 W.

Ocean heat transport in the Pacific

309

In this procedure, the Ekman and western boundary current components each carry approximately half of the poleward heat transport of 0.76 x 1015 W across 24°N in the Pacific (Table 3, method 1). In the procedure used by HALL and BRYDEN (1982), the Kuroshio and mid-ocean geostrophic flows are each broken down into a baroclinic heat flux with no net mass transport and a barotropic transport at the volume-averaged temperature. The Ekman layer and western boundary current components of the heat transport are identified by the products of their transports times their temperature excesses over the average temperature of the section of 3.56°C (Table 3, method 2). Such methodology yields a much larger Ekman heat transport of +0.93 x 1015 W, a much larger Kuroshio heat transport of 0.88 x 1015 W, but there is a net southward baroclinic heat transport of 1.06 x 1015 W which compensates for the larger Kuroshio and Ekman heat transports. Clearly, the net heat transport remains the same but the definitions can change the relative importance of the various components. The northward Kuroshio flow and wind-driven Ekman transport surely return southward at temperatures warmer than the area-averaged temperature of 3.560C and the baroclinic heat transport accounts for the warmer temperature of the southward return flow. To consider the western boundary current component or the Ekman component in isolation from the complete meridional circulation then can give an exaggerated impression of their importance in transporting heat. For example, KRAUSand LEvrrus (1986) argue that the seasonal variability in Ekman heat transport in the North Pacific is enormous because they define the Ekman heat transport as the product of Ekman transport times the difference in the surface layer temperature of order 23°C and their section-averaged temperature of 3.61°C and consider the Ekman component in isolation from the baroclinic return flow. Until one considers the complete three-dimensional circulation, the magnitudes of various components in the heat transport remain overwhelmingly dependent on their definitions. MERIDIONAL CIRCULATION We consider that the best way to understand the mechanisms of ocean heat transport across 24°N in the Pacific is to distribute the meridional transport across the section into temperature classes for each of the components (Table 4, Fig. 4). Such a tabulation immediately shows that the wannest water flows northward for all three components: in the Kuroshio, in the wind-driven surface Ekman layers, and even in the mid-ocean where the warmest water on the western side of the section still flows northward. From the total transport for all three components (Fig. 5), it is clear also that the sum of northward Kuroshio and Ekman layer flows are balanced by the southward, mid-ocean geostrophic flow wanner than 5.50C, an isotherm whose maximum depth on the transpacific section is only 900 m. This upper water circulation carries essentially all of the heat transport across 24°N. The division of the heat transport into a horizontal circulation cell and a vertical meridional circulation cell is described in the next section. First, however, the intermediate and deep water circulation across 24°N below 800 m depth is described. While this deeper circulation accounts for only a small portion (0.02 x 1015 W) of the northward heat transport across 24°N, it is essential to show that the deep meridional flow is reasonable in order to form a basis for the upper water circulation and heat transport estimates.

310

H.L. BRYDENet al.

For intermediate waters colder than 5°C, a breakdown of meridional transport into finer temperature intervals of 0.2°C is helpful (Fig. 6). There is a general northward flow for all waters between 4,5 and 1.9°C, which correspond on average to depths between 800 and 2000 m. While this northward flow is strongest over the Philippine Basin in the west, it is also evident in the central and eastern portions of the section as a small reversal in the generally southward mid-ocean transport. We associate this flow with the remnants of the northward-flowing Antarctic Intermediate Water, although its characteristic salinity minimum has been eroded by mixing. By summing the transports in all temperature classes below 5.5°C which are northward, we estimate that there is a net northward transport of Antarctic Intermediate Water between 4.5 and 1.9°C of 4.3 × 106m3s -1, with 3.3 x 106 m3 s -1 of this flow occurring in the western portion of the section over the Philippine Basin. As a side note, the North Pacific Intermediate Water transport above 800 m is more difficult to characterize. While its salinity minimum is clearly evident in the thermocline throughout the entire 24°N section (ROEMMICHe t a l . , 1991, Plate 2), its general southward flow is part of the subtropical circulation of the thermocline in the central and eastern portions of the 24°N section, and it flows northward both in the Kuroshio and western

Table 4.

Northward transport as a function of temperature for Kuroshio, Ekman and mid-ocean geostrophic return flows

Temperature

Kuroshio

28 27 26 25 24 23 22 21 20 19 18 17 16 15 14 13 12 11 10 9 8 7 6 5 4 3 2 1

0.41 2.88 1.79 1.32 1.30 1.18 1.15 1.14 1.29 1.32 1.93 2.23 1.46 1.20 1.09 1.11 1.12 1.03 0.88 0.54 0.46 0.44 0.45 0.23 0.31

Ekman

0.56 2.39 1.89 1.82 2.37 1.03 0.29 0.49 0.48 0.29 0.17 0.12 0.09 0.01

Mid-ocean

Sum

0.03 2.26 0.05 - 1.12 -0.95 -2.93 -2.28 - 1.20 -5.82 -3.31 -4.79 -2.31 -2.45 -2.25 -2.18 -2.29 -2.13 -2.05 -1.47 - 1.25 -0.55 -1.33 -0.37 1.88 0.99 0,78 -3.22

0.41 2.91 4.61 3.76 2.07 2.05 0.59 -0.11 +0.38 -4.01 -0.90 -2.27 -0.68 -1.13 - 1.07 - 1.06 -1.17 -1.10 -1.17 -0.93 -0.79 -0.11 -0.88 -0.14 +2.19 0.99 0,78 -3.22

311

Ocean heat transport in the Pacific

portion of the section over the Philippine Basin. Thus, it goes with the flow everywhere and any estimate of its net southward transport across this section would depend overwhelmingly on the exact definition of the intermediate water properties. For the deep and bottom waters colder than 1.3°C (deeper than 3200 m) a breakdown of the meridional transport into even finer temperature intervals of 0.02°C for each of the western, central and eastern portions of the 24°N section shows the structure of the deep circulation (Table 5). Most notable is the northward transport of the coldest waters across each section. ROEMMICHet al. (1991) have identified the coldest water in the central basin as northwardflowing Antarctic Bottom Water. Here, we find a northward transport of 4.8 x 1 0 6 m 3 s - 1 of water colder than 1.05°C in the central basin. In the eastern basin, there is an additional northward transport of 0.1 x 1 0 6 m 3 s - 1 of water colder than 1.04°C. Thus, we estimate a total northward transport of Antarctic Bottom Water across 24°N in the Pacific of 4.9 x 10 6 m 3 s - 1 . In the Philippine Basin, the coldest water is also flowing northward. In contrast to the central and eastern basins, however, the coldest water is not the deepest water. Instead, the coldest water in the Philippine Basin is banked against the IzuOgasawara Ridge at depths of 4000-5000 m. We estimate that this water with potential

-4

"4

-2

0

2

4

,

,

6

I" I i I I I

Ol

,

,

I

Fig. 4. Northward transport ( x 106 m 3 s- l ) as a function of temperature for the Ekman ( b ) , Kuroshio (- - - ) and mid-ocean geostrophic (--) components. Transports are determined for each component over temperature intervals of 1.(PC.

312

H . L . BRYDENet

al.

~ •-8 30

-4 J

-2 i

(xlO6m3s"~) 0

2

4

I

t

J

8

m m ! 20-

m

ma m m [] 8-

m m [] o Fig. 5.

I

I

I

I

Net northward transport ( × 1 0 6 m 3 s - l ) across 24°N in the Pacific as a function of temperature. Transports are determined over temperature intervals of 1.0°C.

temperature less than 1.18°C flows northward with a transport of 0.7 x 10 6 m 3 s -a and returns at slightly warmer temperatures in the central and western Philippine Basin. The deep water mass between 1.9 and 1.05°C, that is between the Antarctic Intermediate Water and Antarctic Bottom Water, is called North Pacific D e e p Water (CocnRnr~E, 1956) or Common Water (MoNT6OME~Y, 1958). This water mass occupies more than half of the area of the 24°N section, between the depths of 2000 and 4700 m in the central and eastern basins and below 2000 m depth in the Philippine Basin. According to WoRrrtINCTOr~'S(1981) census, it is the most abundant of all water masses. As indicated by its high silicate values which achieve a maximum greater than 170pmol 1-1 in the eastern basin of the 24°N section (Ror~Mt~icn et al., 1991, Plate 5), it is the oldest water mass in the world ocean. Here, the North Pacific Deep Water flows southward across 24°N with a transport of 8.4 × 10 6 m 3 s - 1 . Because of its enormous area on the 24°N section, this Deep Water transport is most sensitive to reference level velocities and overall mass balance constraints. In their inverse calculation for the large-scale circulation of the North Pacific, RoEM~Icr~ and MCCALUSTER (1989) found a northward transport of 10 x 10 6 m 3 s - 1 of Antarctic B o t t o m Water across the 24°N section, which was almost exactly balanced by a southward transport of North Pacific Deep Water. They found essentially no northward transport of

Ocean heat transport in the Pacific

313

Antarctic Intermediate Water, which is consonant with the lack of any tracer signal of the Antarctic Intermediate Water at 24°N (ROEMmCX et al., 1991). Thus, different analysis procedures can lead to different intermediate, deep and b o t t o m water transports, so that uncertainty still remains in the actual deep circulation across a long zonal section such as the 24°N transpacific section. It is notable that the different deep water circulations do not appreciably affect the estimate of meridional heat transport as long as there is little net deep water transport, since the differences in temperature among the deep water masses are so small. Indeed, there remains a problem of order 2 x 106 m 3 s- 1 in the meridional circulation of deep and bottom waters found here. Since the North Pacific is closed at the northern boundary for the deep circulation, the southward-flowing D e e p Water across 24°N must have been formed in the northern North Pacific by mixing between the Antarctic Intermediate and Bottom Waters flowing northward across 24°N. While there is sufficient northward transport of Intermediate and B o t t o m Waters to balance the southward transport of Deep Water, the D e e p Water flow across 24°N has an effective temperature of 1.23°C, so close to the 1.00°C temperature of the northward-flowing Antarctic Bottom Water that only a small amount of Intermediate Water can be involved in its formation.

~¢~l~maa"t) •.6 • 0

-4 i

-2 I

0

2 I

4 i

6

4,0-

! | I ! I

0-8

l

i

i

I

Fig. 6. Northward transport (×106m3s -I) of intermediate, deep and bottom waters as a function of temperature below 5°C. Transports are determined over temperature intervals of 0.20C.

314

H.L. BRYDENet al. Table 5. Northward transport (106 m-3s -I) as a function o f potential temperature for the deep circulation colder than 1.3°C. The eastern section extends from California to the Hawaii Ridge; the central section from the Hawaii Ridge to the lzu-Ogasawara Ridge; and the western section from the Izu-Ogasawara Ridge to Okinawa across the Philippine Basin

Potential temperature (°C)

Western section Stas 306-369

Central section Stas 155-305

Eastern section Stas 1-152

1.29 1.27 1.25 1.23 1.21 1.19 1.17 1.15 1.13 1.11 1.09 1.07 1.05 1.03 1.01 0.99 0.97 0.95

-0.05 -0.09 -0.08 -0.25 -0.30 -0.32 +0.68

-0.05 0.09 -0.04 -0.02 -0.05 -0.11 -0.13 -0.22 -0.31 -0.43 -0.53 -0.20 0.56 0.74 1.13 1.04 0.40 0.92

-0.18 -0.21 -0.14 -0.33 -0.59 -0.83 -0.25 -0.10 -0.21 -0.49 -0.44 -0.32 -0.09 +0.12

We estimate that 1.2 × 106 m 3 S - 1 of 2.2°C Intermediate Water could mix adiabatically with 4.9 x 106 m 3 s -1 of 1.00°C Bottom Water to form 6.1 x 106 m 3 s -1 of 1.23°C Deep Water. Thus, we would be more contented if only 6.1 x 106m 3s - t (instead of 8.4 × 106 m 3 s -1) of North Pacific Deep Water were flowing southward across 24°N. For this reason, we conclude that there is an uncertainty of order 2 x 106 m 3 s -1 in the deep circulation found here. In summary, the meridional circulation across 24°N below 800 m depth can be characterized by a northward flow of 4.3 × 106mas -1 of Antarctic Intermediate Water with potential temperatures between 4.5 and 1.9°C; a southward flow of 8.4 x 106 m 3 s -1 of North Pacific D e e p Water with temperatures between 1.9 and 1.05°C; and a northward flow of 4.9 × 10 6 m 3 s - 1 of Antarctic Bottom Water with temperatures less than 1.05°C.

CONTRIBUTION OF HORIZONTAL AND VERTICAL CIRCULATION CELLS TO THE NORTHWARD HEAT TRANSPORT To separate the contributions of horizontal and vertical circulation cells in the heat transport, profiles of geostrophic meridional mass transport per unit depth (equal to the zonal integral of the geostrophic velocity across the section) are determined for the midocean and Kuroshio portions of the 24°N section. The southward mid-ocean geostrophic flow is larger than the Kuroshio flow so that the zonally averaged geostrophic transport across 24°N is southward from the sea surface down to 800 dbar (Fig. 7). From 800 to

Ocean heat transport in the Pacific

315

NORTHWARD TRANSPORT PER UNIT DEPTH (xlO4m2s ''1 -14

-f2

-10

-8

-6

-4

-2

0

2

4

6 ....

8

10

--°

1000-

I X~00-

4000-

6000-

~OOI

I

i

I

I

I

I

I

I

I

Fig. 7. Vertical profile of geostrophic northward mass transport per unit depth (x 104 rn2 s--l ) for the mid-ocean from California to Okinawa (--), for the Kuroshio over Okinawa Trough ( - - - ) and for the total 24"N section (--).

2140 dbar, the Antarctic Intermediate Water flows northward; from 2200 to 4800 dbar, the North Pacific Deep Water flows southward; and below 4800 dbar the Antarctic Bottom Water flows northward. To this profile must be added the northward wind-driven Ekman transport per unit depth of 24 x 104 m 2 s -1 in the upper 50 m in order to obtain the total, zonally averaged meridional circulation. Half of the northward Ekman transport is balanced by the net southward geostrophic flow over the upper 200 m; and all of the Ekman transport is compensated by the geostrophic flow above 720 dbar. In fact, vertically integrating the meridional transport per unit depth, we find that the net meridional transport is zero at 720, 940 and 3800 dbar, as well as at the sea surface and bottom. The northward ocean heat transport due to the vertical circulation cell, Hv, is estimated by integrating the zonally averaged meridional transport per unit depth, (v) x L(z), where L(z) is the distance across the 24°N section at depth z, times the zonally averaged temperature, (T), and times the specific heat over the entire column to be 0.38 x 1015W:

Hv = -pCp r (v)(T)L(z) dz = J- H

+0.38 x 10a5 W.

Because the vertically integrated mass transport is zero at 720, 940 and 3800 dbar, the northward heat transport due to the vertical meridional circulation cell can be estimated down to these depths as well. From the sea surface down to 720 dbar, the circulation transports 0.36 × 10 '5 W of heat northward, which is virtually all of the transport by the

H.L. BRYDENeta1.

316

top-to-bottom zonally averaged circulation. Thus, the meridional heat transport due to the vertical circulation cell, Hv is essentially accomplished by the flow in the upper 700 m of the water column. The deeper circulation below 700 m contributes only 0.02 × 1015 W to the northward heat transport. The heat transport per unit depth due to the horizontal circulation cell, HH(Z), is estimated by multiplying the difference between the meridional velocity at each station pair and the zonally averaged velocity at each depth, v - (v), times the difference between the potential temperature at each station pair and its zonally averaged value, T - (T), and integrating across the section: HH(Z) = pCp I (v -- (v))( T -

(T)) dx.

Because there is no net mass transport at any depth for this component, the vertical distribution of the meridional heat transport due to the horizontal circulation cell can be unambiguously defined (Fig. 8). Vertical integration of this profile of heat transport per unit depth from the sea surface to the bottom yields a northward heat transport of 0.38 x 1015 W due to the horizontal circulation cell. Again, effectively all of the northward heat transport in this component is contributed by the horizontal circulation in the upper 700 m of the water column.

NORn.NVAI~ PEAT RJJX PER UNIT DEPTH (xX:~ZW M ~) pCp ( v - -1

0

1 I

0

) ( T - <'T> ) x L ( z ) 2 I

3 I

I

L

1000-

~

2000-

#~00

-

I

Fig. 8. Verticalprofile of the northward heat flux per unit depth (x 1012W m- t ) for the purely horizontal circulation cell which has no mass transport at any depth.

Ocean heat transport in the Pacific

317

Formally then, the total northward heat transport of 0.76 x 1015 W across 24°N in the Pacific is due half to the zonally averaged vertical meridional circulation cell and half to the horizontal circulation cell. The vertical meridional cell is a shallow one, however, in which the northward, wind-driven Ekman transport of warm surface waters returns southward only slightly deeper and colder, all within the upper 700 m of the water column. In terms of its meridional heat transport, the horizontal circulation cell is also shallow with effectively all of its northward heat transport contribution in the upper 700 m of the water column. Thus, meridional heat transport across 24°N in the Pacific Ocean is effected primarily by the circulation above 700 m depth. Warm water flows northward on the westena side of the Pacific and in the wind-driven near-surface layers and returns southward in the central and eastern Pacific at colder temperatures. Because the water is colder in the eastern Pacific at similar depths than in the western Pacific, much of the heat transport can be carried by a purely horizontal circulation. North of 24°N, northward-flowing warm water must be converted into water only as cold as 5.5°C to form the colder, southward flow across 240N at depths shallower than 700 m. No deep convection in the northern Pacific is necessary to maintain the northward heat transport across 24°N and this is consistent with there being no deep water formed in the North Pacific (WAssEN, 1983).

ERROR ESTIMATES

To determine the uncertainty in this estimate of North Pacific heat transport, several alternatives are examined. First, in terms of the overall circulation, we note that the North Pacific Deep Water exhibits about 2 x 106 m 3 s -1 too much southward transport, in comparison with heat budget constraints on its formation. Since the Deep Water between 1.05 and 1.9°C occupies a little more than half of the area on the 24°N section, its excess transport suggests that the upper water circulation components may have errors of order 4 x 106 m s s -1. For example, if the combined northward Ekman and Kuroshio transports were 36 x 106 m 3 s -1 instead of 40 x 106 m 3 s -z, the southward Deep Water flow would be reduced by 2 x 106 m 3 s -1 and would be consonant with its heat budget considerations. Another possibility which would affect the heat transport is that there could be an additional southward geostrophic flow of order 4 x 106m 3 s -1 over relatively shallow bathymetry, such as near the Izu-Ogasawara Ridge where southward bottom velocities over the slope region between 1000 and 3000 m depths would not be unreasonable. In terms of heat transport error, the worst case would be 4 x 106 m 3 s -1 change in the Ekman component, which has the warmest temperature of 22.56°C above the section-averaged temperature of 3.56°C. Thus, for a maximum change of 4 x 106 m 3 s -1 in the overall circulation, the change in northward heat transport is no larger than 0.3 x 1025 W. For changes in the circulation which do not affect the deep circulation below 700 m, the consequences for heat transport are less severe. If the variations in the Kuroshio or Ekman transports are compensated by changes in the mid-ocean geostrophic flow above 700 m depth, which has an effective temperature of 15.110C (Table 3), the northward heat transport changes in proportion to the difference between the temperature of the Kuroshio (18.46°C) or the Ekman (22.56°C) flows and the effective temperature of the upper water return flow. Such differences are only 3.4°C for Kuroshio variations and 7.4°C for Ekman variations. Thus, for a 10 x 106 m 3 s -1 change in the Kuroshio or Ekman transports, the northward heat transport would vary only by 0.14 or 0.30 x 1025 W.

318

H . L . BRYDENet al.

Lastly, there is the possible contribution of a meridional eddy heat flux across 24°N due to temporal fluctuations, which are not adequately resolved by the single transpacific section. Long-duration current meter measurements are the best way to estimate the eddy heat flux. Unfortunately, there are no current meter records along the 24°N section, as far as we are aware. The closest analysis that we can find is by NnLER and HALL (1988) for a set of 3-year time series measurements at 28°N, 152°W, which may be typical of the mid-ocean North Pacific. They found southward eddy heat fluxes in the upper 1000 m of the water column which averaged - 0 . 2 5 ° C cm s-l. Multiplying this flux by the area of the 24 ° section above 1000m depth yields an estimate of the eddy heat transport across 24°N of - 0 . 1 3 × 1015 W. While it is dangerous to extrapolate a single measurement of eddy heat flux across the entire transpacific section, it is not unreasonable to think that the eddy heat flux component might be of order 0.1 × 1015W in the middle of the North Pacific subtropical gyre. Clearly, more direct measurements of the eddy heat flux and further analysis of existing time series measurements in the North Pacific are needed to sharpen the estimate of eddy heat transport beyond this order-of-magnitude calculation. Based on an uncertainty in the overall circulation, including the deep water circulation of 4 × 106 m 3 s -1, or an uncertainty in the upper water circulation of 10 × 106 m 3 s -1, we conclude that the error in the northward heat transport across 24°N in the Pacific of 0.76 x 1015 W is +0.3 x 1015 W. ANNUAL VARIABILITY The above discussion of errors naturally leads to a consideration of the annual variability in heat transport across 24°N in the Pacific. First, the transpacific 24°N section could be considered to provide an instantaneous estimate of the A p r i l - M a y heat transport if the April-May Ekman layer transport of 13.8 x 106 m 3 s -1, 1.8 × 106 m 3 s -1 larger than the annual average, were used. Since such an increase is within the possible variation in the overall circulation of 4 x 106 m 3 s -1 discussed above, the mid-ocean, geostrophic southward flow could increase by 1.8 x 106m3s-1 at the section-averaged temperature of 3.56°C. Thus, the northward heat transport for April-May would be estimated to be 0.90 x 1015 W with only small modifications to Table 3, m e t h o d 2. From differences in sea level across Tokara Strait, BLAHA and REED (1982) estimate an annual variation in Kuroshio transport of _+5 x 106 m 3 s -1, with maximum transport in July. The HELLERMAN and ROSENSTEIN (1983) wind stress tabulations suggest an annual variation in northward E k m a n transport across 24°N from 16.8 x 106 m 3 s -1 in November to 4.6 x 106 m 3 s -1 in January about an annual mean of 12.0 x 106 m 3 s -1. There is little information on the annual variation in mid-ocean geostrophic flow since the transpacific 24°N hydrographic section has only been made once. It is reasonable to assume that there is little seasonal variation in the intermediate and deep water circulation for temperatures less than 5.5°C below 700 m depth. The annual variability in Kuroshio or E k m a n transport then would be balanced by variability in the mid-ocean circulation above 700 m depth. The temperature difference between the Kuroshio or E k m a n flows and the mid-ocean circulation in the upper 1000 m is of order 6°C and because of phase differences the annual variation in Kuroshio plus E k m a n transport appears to be about 8 x 106 m 3 s -1. Thus, the annual variation in northward heat transport is of the order of 0.2 x 1015 W about the mean heat transport of 0.76 x 1015W with a maximum in July, as long as the deep circulation exhibits no annual variation.

Ocean heat transport in the Pacific

319

KRAUS and LExaaxJs (1986) found much larger annual and interannual variations in North Pacific meridional heat transport across 24°N, of order 0.6 × 1015 W, because they assumed implicitly that Ekman transport variations are compensated by changes in the deep circulation with an average temperature of 3.61°C. We know of no evidence for seasonal variations in the deep Pacific circulation, however, and we think it is more reasonable at the present time to assume that the deep circulation exhibits little annual variability. Thus, we maintain that the annual variation in northward heat transport across 240N is only 0.2 x 1015W for the combined annual variation in Kuroshio and Ekman transports. For more refined estimates of the annual variation in northward heat transport, it is clear that more information is required on the seasonal cycle of the Kuroshio, on the annual variations in the mid-ocean geostrophic flow in the upper 700 m, and on any evidence for seasonal variability in the intermediate and deep water circulation. DISCUSSION The mass balance achieved for the transpacific 24°N section, in order to determine the heat transport, has intrinsic interest for water mass transports. The northward flow of Antarctic Intermediate Water across 24°N would not be suspected on the basis of meridional sections such as the 160°W section of REID (1965, Fig. 3) or the GeOsecs sections (Cv,AIGet al., 1981, Plates 5 and 25) because the tongue of low salinity associated with Antarctic Intermediate Water dies out before it reaches the equator. REID'S (19651 Fig. 7) 27°N section exhibits a slight reversal in isotherm slopes below 1000 m depth, perhaps suggesting an intermediate maximum in northward velocity, but the data sources are too diverse to warrant much confidence. The 24°N transpacific section clearly shows the reversal in isotherm slope at about 3.50C (RoF.MMICHet aI., 1991, Plate 1), which marks the maximum in northward flow here. It is the overall mass budget carried out for the 240N section which allows the magnitude of this northward velocity maximum to be determined and hence the northward transport of Antarctic Intermediate Water with potential temperatures between 1.9 and 4.50C to be estimated to be 4.3 x 106 m 3 s -1. Similarly, the overall mass budget allows the southward transport of North Pacific Deep Water with potential temperatures less than 1.50C to be estimated to be 8.4 x 106 m 3 s -1. This Deep Water has the greatest volume of all the ocean's water masses (WoI~TmNGTOr% 1981). Summing the volumes of all North Pacific water with potential temperatures less than 1.50C according to Worthington's census yields a volume for the North Pacific Deep Water of 1.23 x 108 km 3. Dividing this volume by its southward transport across 240N of 8.4 x 106 m 3 s -1 suggests a renewal time scale for the North Pacific Deep Water of order 500 years. The northward transport of Antarctic Bottom Water of 4.9 x 106 m 3 s -1 across 240N shows a reasonable reduction from the northward transports recently estimated in the tropical Pacific. TAFT et al. (1991) found 12 x 106 m 3 s -1 of Antarctic Bottom Water flowing northward through the Samoan passage into the central Pacific basin and JonNsor~ (1990) estimated a total transport of Antarctic Bottom Water across the 10°N transpacific section of 8.4 x 106 m 3 s -1. The salinity flux across the 240N section is readily calculated to be northward at a rate of 5.5 × 106 kg s -1 (5.5 Sv%o) by multiplying salinity times velocity and integrating over the entire transpacific section. This flux implies a freshening of the water north of 240N,

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consistent with BAUMGARTNERand REICHEL'S (1975, Plate 27) net precipitation maps for the North Pacific, and a southward freshwater flux. Because the small flow out of the North Pacific through the Bering Strait of 0.8 × 106 m 3 s -1 has such a low salinity of 32.5%o (COACHMANand AAGARD, 1988) compared with the section-averaged salinity of 34.6%0 at 24°N, it is important to take account of the resulting salinity flux convergence of 1.7 × 100 kg s -1 in the freshwater budget for the region north of 24°N. The total freshwater divergence for the North Pacific between 24°N and the Bering Strait is then estimated by dividing the salinity flux by the section-averaged salinity to be 0.21 × 106 m 3 s-1, which is in reasonable agreement with the independent determination from hydrological data of 0.26 x 106 m 3 s -1 by BAUMGARTNERand REICHEL (1975, Table xxxv). The northward heat transport across 24°N in the Pacific of 0.76 x 1015 W is comparable with the northward heat transport across 24°N in the Atlantic of 1.22 × 1015 W. The mechanisms of the transports are quite different for the two oceans, however. As shown above, the Pacific heat transport is effected primarily by a circulation in the upper 700 m in which warmer water flows northward on the western side of the basin and in the surface layers and returns southward in the central and eastern Pacific at colder temperatures. There is only a shallow vertical meridional circulation cell and vertical convection is limited to depths of order 700 m. In the Atlantic, the vertical meridional cell with associated deep convection effects the northward heat transport. There, a net northward flow across 24°N of 18 x 106 m 3 s -1 of water warmer than 6°C in the upper 1000 m of the water column is converted in northern regions into North Atlantic Deep Water and returns southward across 24°N as intermediate and deep water between 1100 and 4500 m depths at temperatures of 2--6°C (HALL and BRYDEN, 1982). The horizontal circulation across 24°N, in which warmer water flows northward on the western side of the Pacific and colder water returns southward in the central and eastern Pacific at nearly the same depth, requires a loss of heat by the ocean to the atmosphere north of 24°N at the rate of 0.76 x 1015 W. Such a heat loss north of 24°N is in marked contrast with the recent tabulations of air-sea fluxes from bulk formula by Clark for the North Pacific which yield a net heat gain by the ocean north of 25°N of 0.17 x 1015 W. Such bulk formula calculations must be in error. For a heat loss of 0.76 x 1015 W over the 35 × 1012 m Esurface area of the Pacific north of 24°N, the average heat loss by the ocean is 22 W m -2, while Clark's air-sea fluxes yield an average net gain of 5 W m -2. Such difference reinforces the common conclusion that bulk formula calculations can be consistently wrong by 30 W m -2 (BRETHERTONet al., 1982). Direct estimates of ocean heat transport such as this one for 24°N in the Pacific may be the best method for calibrating bulk formula calculations of air-sea fluxes. It is surprising that BUNKER'S (1976) bulk formula tabulations of Atlantic air-sea heat exchange agreed so well with the direct estimate of ocean heat transport across 24°N while Clark's tabulations for the Pacific, ostensibly using the same procedure, are so much in error. Of the air-sea flux calculations for the North Pacific summarized by TALLEY(1984), the best agreement with the direct estimate was achieved by ESBENSEN and KUSHNIR (1981). They used Budyko's method of first calculating monthly averaged meteorological variables and then applying drag coefficients appropriate for monthly averages to estimate a heat loss by the ocean north of 24°N of 0.75 x 1015 W. HASTENRATH'S(1980) summary of ocean heat transport derived from air-sea heat exchanges yielded a larger northward heat transport across 24°N in the Pacific of 1.14 × 1015 W. Hastenrath, however, used BuoYr o ' s (1963) fluxes north of 30°N but obtained values of net heat loss by the ocean a factor of

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two larger than BUDYKO'S(1974) values across 30°N and a factor of three larger than EMIG'S (1967) summary of BUDYKOet al.'s (1962) fluxes. There are obviously several different sets of Budyko's air-sea fluxes. The combination of the northward ocean heat transport across 24°N in the Pacific of 0.76 × 1015W and the northward heat transport across 24°N in the Atlantic of 1.22 × 1015W yields for the first time a reliable estimate of the total poleward ocean heat transport across a complete latitude circle, a northward heat transport of 2.0 x 1015 W across 24°N. Across the same latitude, the atmospheric energy transport is 1.7 x 1015 W (CARlSSIMOet al., 1985). Thus, the ocean transports more heat poleward across 24°N than does the atmosphere; the ocean therefore can be considered to be more important than the atmosphere at this latitude in maintaining the Earth's heat budget. For the first time, direct estimates of ocean and atmosphere heat transports across a complete latitude circle yield a total atmosphere-ocean poleward heat transport of 3.7 x 1015 W across 24°N. Such a heat transport is in reasonable agreement with traditional estimates of the radiation budget requirements at the top of the atmosphere. For example, PALMENand NEWTON(1969) and SELLERS(1965) found that the radiation budget required a combined atmosphere-ocean heat transport across 24°N of 4.0 x 1015 W. With the advent of satellite radiation measurements, however, the heat transport across 24°N required by radiation budgets has been increasing rapidly. VONDER HAAR and OORT'S (1973) first calculation raised the required heat transport across 24°N to 4.8 x 1015 W, and the most recent calculation (CARISSlMO et al., 1985) requires a combined oceanatmosphere heat transport of 5.3 x 1015 W across 24°N, a 30% increase over traditional values. The satellite radiation measurements, however, do have unknown biases which are generally removed independently of latitude but in fact may be latitudinally dependent. An interesting situation has arisen with the new satellite radiation measurements requiring larger total atmosphere-ocean heat transport. Meteorologists generally subtract their direct estimates of atmospheric energy transport from the radiation budget requirements and conclude that the ocean must be transporting an enormous amount of heat poleward (e.g. CARISSIr,tO et al., 1985), an amount far too large to be acceptable to oceanographers. On the other hand, oceanographers subtract their estimates of ocean heat transport from the radiation budget requirements and conclude that the atmosphere must be transporting an enormous amount of heat (e.g. TALLEY,1984), an amount far too large to be acceptable to meteorologists. But are the satellite radiation budget requirements correct? There are many ways to remove biases in satellite radiation measurements. Reducing the amount of incoming radiation in tropical regions or decreasing the amount of outgoing radiation in polar regions would decrease the size of the poleward transport required in the radiation budgets. One way to diagnose the biases in radiation measurements is to calibrate the radiation budget with accurate direct estimates of ocean and atmosphere heat transports across selected latitudes. The first such direct estimate presented here suggests that the combined atmosphere-ocean heat transport across 24°N is 3.7 x 1015 W, far less than the 5.3 x 1015W required by the radiation budget, when biases are removed independently of latitude. Clearly, more such direct estimates are needed to diagnose the biases in satellite radiation measurements. The recent rapid growth in poleward heat transport by the atmosphere and ocean required by radiation budget analyses strongly suggests that the satellite radiation measurements deserve close scrutiny to determine how best to remove their biases.

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Transoceanic hydrographic sections such as the 24°N transpacific measurements discussed in this work are extremely useful for determining the overall meridional ocean circulation, particularly for the deep waters where consistent and complete high-quality measurements are required to estimate the relatively small net meridional flows of the various water masses. From these new measurements, the net northward flow of Antarctic Intermediate Water of 4.3 x 106 m 3 s -1, the net southward flow of North Pacific Deep Water of 8.4 x 106m3s - t , and the net northward flow of Antarctic B o t t o m Water of 4.9 x 106 m 3 s- 1 across 24°N can be estimated quantitatively. H e r e , the emphasis has been directed toward estimating the northward ocean heat transport across 24°N, which requires an accurate determination of how the meridional circulation is distributed by temperature. The northward heat transport across 24°N in the Pacific is estimated to be 0.76 x 1015 W which is surprisingly large. Such heat transport has profound implications for bulk formula calculations of air-sea heat exchange, some of which are apparently in error by 30 W m -2. Combining the Pacific heat transport with the established value of Atlantic heat transport across 24°N yields the first reliable direct estimate of ocean heat transport across a complete latitude circle. The total ocean heat transport of 2.0 x 1015 W across 24°N is larger than the atmospheric energy transport of 1.7 × 1015 W across 24°N. The ocean then is more important than the atmosphere in transporting heat poleward across 24°N to maintain the global heat balance. The combined o c e a n - a t m o s p h e r e heat transport across 24°N of 3.7 x 1015 W is in reasonable agreement with traditional values derived from the radiation budget at the top of the atmosphere, but it is substantially smaller than recent values derived from satellite radiation budget analyses. Satellite radiation measurements, however, have biases which can be removed in various ways. Comparison of direct estimates of ocean-atmosphere heat transport with the satellite radiation budget requirements may help to diagnose the causes of bias in the satellite measurements. With the completion of the 24°N transoceanic sections, it is clear that similar sections across other latitudes are needed to provide accurate determinations of ocean heat transport across tropical and subpolar latitudes. Di~'ect estimates of ocean heat transport can be used to calibrate air-sea heat exchange calculations from bulk formula which may be accurate in some regions but not in others. Direct estimates of ocean heat transport can be compared with values of atmospheric energy transport to understand the roles of the atmosphere and ocean in maintaining the global heat balance. Ocean heat transport can be added to the atmosphere energy transport and the total can be compared with the transport required by the radiation budget in order to diagnose the sources of biases in satellite radiation measurements. Most importantly, the procedure by which meridional ocean heat transport is estimated provides insight into the magnitude of the meridional circulation of various water masses and into the relationship between the water masses and the processes by which the ocean circulation transports heat. transpacific24°Nhydrographicsection and its analysiswere supported by the National Science Foundation. James Swift and Melinda Hall and the PACODF group from Scripps Institution of Oceanography were instrumental in making the high-quality CTD measurements necessary to carry out this analysis of ocean heat transport. Terrence Joyce organized the acoustic Doppler velocity measurements. Ann Spencer made the calculationsreported here. John Church's visit to Woods Hole OceanographicInstitution was supported by a Travelling Fellowship from the Australian government. Comments on an early version of this

Acknowledgements--The

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manuscript by Stuart Godfrey, Melinda Hall, Shiro Imawaki, Terrence Joyce, Trevor McDougall and Peter Rhines led to substantial improvements in the final manuscript. Woods Hole Oceanographic Institution Contribution no. 6570.

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