Mixed layer heat balance in the western North Pacific

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C7, 3242, doi: /2002jc001536, 2003 Mixed layer heat balance in the western North Pacific Tangdong Qu International Pacific Research Center, School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, Hawaii, USA Received 15 July 2002; revised 15 January 2003; accepted 25 March 2003; published 29 July [1] The mixed layer heat balance in the western North Pacific is examined using historical temperature data and U.S. National Centers for Environmental Prediction reanalysis surface wind and heat flux. Although the data come from different sources, the mixed layer heat budget is balanced reasonably well in the region studied. For the annual average, the surface thermal forcing is significant mainly in three regions. One is in the Mindanao Dome, where the incoming surface heat flux is balanced in a large part by the vertical entrainment induced by Ekman pumping. The second occurs in the Kuroshio region, where much of the heat advected by the Kuroshio is released to the atmosphere. The third is located in the central subtropical gyre where the annually integrated surface thermal forcing is balanced primarily by vertical entrainment associated with the deepening of the mixed layer. Although seasonal variation in sea surface temperature (SST) is mainly due to surface thermal forcing, ocean dynamics also has an effect, particularly, during the developing stage of the summer monsoon. From late May to early July, there is a northeastward increase in SST (>29 C) in the region E, N, which coincides with the northeastward onset of the summer monsoon. As the summer monsoon develops, vertical entrainment becomes effective at cooling, leading to a decrease in SST against the incoming surface heat flux over a large part of the western North Pacific. INDEX TERMS: 4572 Oceanography: Physical: Upper ocean processes; 4576 Oceanography: Physical: Western boundary currents; 4504 Oceanography: Physical: Air/sea interactions (0312); KEYWORDS: summer monsoon, sea surface temperature, surface heat flux, heat advection, vertical entrainment Citation: Qu, T., Mixed layer heat balance in the western North Pacific, J. Geophys. Res., 108(C7), 3242, doi: /2002jc001536, Introduction [2] The western North Pacific is a key region of the Earth s climate system. The region s contrast between sea surface temperature (SST) and land temperature sets up the Asian- Australian monsoon. The atmospheric heating associated with the heavy monsoon rains may have a global impact on atmospheric circulation [e.g., Yasunari and Seki, 1992; Lukas, 1996; Wang et al., 2000]. The western North Pacific is also a major source of the interannual variations associated with the El Nino/Southern Oscillation (ENSO), in which a small change in SST may affect the ocean and atmosphere throughout the entire planet [e.g., Rasmusson and Carpenter, 1982; Palmer and Mansfield, 1984; Wang et al., 1999]. [3] The prediction of SST requires detailed information about the surface heat balance and thus a better understanding of the processes that control SST variability. However, because of the lack of spatially extensive and temporally continuous data, the surface heat balance in the western North Pacific has never been carefully examined. Niiler and Stevenson [1982] considered the time-averaged heat balance of the western Pacific warm pool by analyzing the volume Copyright 2003 by the American Geophysical Union /03/2002JC bounded by the mean 28 C isotherm. As a direct consequence of the control volume they chose, Niiler and Stevenson concluded that turbulent ocean mixing is the dominant mechanism counterbalancing the surface heat flux into the ocean. Later analyses of results from general circulation models [e.g., Qu et al., 1997] provided further evidence for Niiler and Stevenson s conclusion and showed that over the entire tropical western Pacific (i.e., 20 S 20 N, western boundary: 180 ) heat gained from the atmosphere is predominantly transported downward to the deep levels by vertical motion and mixing. On closer inspection of smaller areas of the tropical western Pacific, however, Qu et al. [1997] found that different mechanisms are balancing the surface heat budget in different regions, and in some of these regions, horizontal advection is not negligible. [4] Regarding time variations, Gill and Rasmusson [1983] and Meyers and Donguy [1984], among others, have documented cases in which ocean dynamics seems to play a role in generation of the observed SST anomalies in the tropical western Pacific, in particular, those associated with ENSO. But until recently, only preliminary results on timedependent heat budget have been published [McPhaden, 2002; J. M. Toole et al., On the time-dependent internal energy budgets of the tropical warm-water pools, submitted to Journal of Climate, 2002, hereinafter referred to as Toole 35-1

2 35-2 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 1. Spatial distribution of the temperature profiles (232,751) used for this study. et al., submitted manuscript, 2002]. Consequently, the effect of ocean dynamics on SST variability in the tropical western Pacific is not well understood. [5] In the off-equatorial region, the most significant SST variability is associated with the onset of the summer monsoon. Previous studies have shown that the summer monsoon starts at the lower latitudes of the western North Pacific in early to mid-june and then advances northeastward in the region E, N [e.g., Wu and Wang, 2001; Wu, 2002]. This northeastward march of the onset of the summer monsoon is attributed to the seasonal variation in SST [Wang, 1994; Murakami and Matsumoto, 1994; Ueda and Yasunari, 1996]. A better understanding of the seasonal heat budget in this particular region is therefore highly desirable. [6] This study is intended to provide a detailed description of the mixed layer heat balance in the western North Pacific, using all existing temperature profiles, most of which come from expendable bathythermograph (XBT) measurements, combined with reanalysis of surface wind and heat flux data prepared by the U.S. National Centers for Environmental Prediction (NCEP). The results of this study are presented as follows: section 2 describes the data and methods of analysis; section 3 shows the general characteristics of the western North Pacific; section 4 provides a theoretical background for the present analysis; sections 5 and 6 have an assessment of the annual mean and seasonal variation of the mixed layer heat balance, respectively; section 7 summarizes the results. An error analysis is included in the Appendix A. 2. Data and Methods of Analysis 2.1. Temperature Data [7] In order to develop a three-dimensional picture of the upper layer thermal structure and its seasonal variation, the temperature profiles at observed levels recorded on the CD- ROMs of World Ocean Database 1998 of NOAA/NESDIS/ NODC from the region 0 30 N, E were used. Those profiles flagged as bad or as not passing the monthly, seasonal, and annual standard deviation checks [Levitus, 1994] were dropped. Then those profiles with obviously erroneous records (e.g., temperature <10 C in the upper 100 m) and those profiles with exceptionally sparse sampling in the upper ocean (e.g., vertical distance between two samples >50 m in the upper 100 m) were removed. Some early observations were of poor quality; extreme outliers were not uncommon in some areas, requiring extensive hand editing to remove. The final data set for the present study consists of 232,751 temperature profiles (Figure 1). These temperature profiles span the period from the 1920s to the middle of the 1990s, of which more than 81% (189,687 profiles) were collected after The spatial distribution of the data is essentially the same for each season, except that the density of sampling is biased slightly toward northern summer (not shown). [8] The temperature profiles were first linearly interpolated onto a 5-db uniform pressure series, and then gridded in a 1 1 horizontal resolution for every 5 days (called pentad hereinafter) regardless of the year of observation. The upper ocean was sampled best in the northwestern region, with the sample size in each grid bin being usually greater than five and reaching sometimes several 10s. For the region with poor data coverage, the horizontal radius was widened so that each grid bin had at least five samples. Standard deviations were estimated during the averaging process (discussed in section 7) and used to edit the resulting mean fields. Observations that deviated from the preliminary ensemble mean by more than three standard deviations were discarded. Finally, the gridded pentad mean temperature data were smoothed using a two-dimensional

3 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC 35-3 Figure 2. Annual mean (a) sea surface temperature ( C) from the present data and (b) surface heat flux (W m 2 ) from NCEP reanalysis data. Gaussian filter, with e-folding scales of 15 days and 2 longitude and latitude. [9] A corresponding estimate of potential density (s q ) was derived from a climatological T/S relationship prepared by Qu and Lukas [2003]. This monthly averaged data set has a horizontal grid of , and consequently, resolves the narrow western boundary currents better than most, if not all, of the existing climatologies in the western North Pacific [e.g., Levitus, 1994] NCEP Reanalysis Data [10] The climatological pentad mean surface wind stress, shortwave radiative heat flux, longwave radiative heat flux, sensible heat flux, and latent heat flux used in the present study were constructed from the NCEP daily mean reanalysis data for the period , the period during which most of the temperature data used in the present study were collected. The original data set with a horizontal resolution of was interpolated onto a 1 1 grid to be consistent with the temperature data described above. 3. General Characteristics of the Region [11] Before proceeding to the heat budget analysis, this section shows the annual mean SST, surface heat flux, mixed layer depth (MLD), and surface buoyancy flux. These annual mean fields have the following characteristics Sea Surface Temperature [12] The annual mean SST in the western North Pacific has a pattern dominated by zonally oriented contours (Figure 2a). This pattern is slightly distorted near the maritime continent and East Asia, which is presumably due to the influence of the monsoon circulation. The SST is higher than 28 C in most of the tropical western Pacific (15 of the equator), with its maximum (higher than 29 C) near the equator northeast of the Papua New Guinea. The SST decreases toward the north and falls below 24 C at about 30 N. The annual SST range (maximum - minimum) has a pattern (not shown) similar to the annual mean field, with about 1 C near the equator to about 8 C at30 N Surface Heat Flux [13] The spatial distribution of the time-averaged NCEP reanalysis surface heat flux (Figure 2b) shows a pattern similar to that presented in the climatological atlas [e.g., Oberhuber, 1988] with heating of the ocean occurring in the tropics and cooling at higher latitudes. The zero line of surface heat flux extends roughly from about 17 N at the western boundary to about 7 N near the international dateline. In the northwest corner of the domain, the surface heat flux is < 100 W m 2, forming a sharp contrast with that ( 20 W m 2 ) in the interior ocean. The distribution of surface heat flux near the western boundary differs from that expected from solar radiation, indicating that ocean dynamics exert a strong influence on the heat budget, as will be discussed in the following sections. [14] The annual surface heat flux range (maximum - minimum) is generally large (not shown), running from about 50 W m 2 near the equator to more than 450 W m 2 in the northwestern subtropical gyre. This large seasonal variation is the primary driving force of the annual fluctuations in SST Mixed Layer Depth [15] The thickness of the mixed layer is a key factor determining SST. The criteria used to define the MLD have been diverse in literature. As in several earlier studies [e.g., Lindstrom et al., 1987; Lukas and Lindstrom, 1991; Sprintall and Tomczak, 1992], the MLD here is determined as where s q is equal to the sea surface s q plus the increment in s q equivalent to a desired net decrease in temperature (1 C). This increment in s q takes into account the contribution of both temperature and salinity. Given this criterion, when the salinity stratification within the surface layer is negligible, the MLD is the same as the thermocline depth when the latter is based on the desired temperature difference (1 C) from the sea surface. In the case when salinity is positively stratified, the MLD is shallower than the thermocline and the distance that separates the bottom of the mixed layer from the top of the thermocline is often referred to as the barrier layer [e.g., Lukas and Lindstrom, 1991; Sprintall and

4 35-4 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 3. Annual mean (a) MLD (m) and (b) surface buoyancy flux (10 4 erg m 2 s 1 ) forced by wind stirring and cooling convection. Tomczak, 1992; Vialard and Delecluse, 1998a, 1998b; Kara et al., 2000a]. [16] The sensitivity of MLD to its defining criteria needs some discussion. In a recent study, Kara et al. [2000b] suggested that an optimal estimate of MLD may be obtained using a density-based criterion of T = 0.8 C, though its usefulness is, in some cases, seasonally dependent. The T =0.8 C criterion was also tried in the present study and yielded an MLD field that has the same pattern as that obtained with the T =1 C criterion cited above, except for a slight difference in magnitude. Further analysis of mixed layer heat balance indicated that the present results are not affected by using the higher temperature increment of T = 1 C. [17] The MLD is generally large in the western North Pacific (Figure 3a), and the annual mean field contains two places where MLD is larger than 70 m: one extends westward along 15 N and the other extends southwestward along the western boundary of the northern subtropical gyre. The annual range (maximum - minimum) of the MLD is small in the tropics, and its minimum (<10 m) is associated with the Mindanao Dome at 6 8 N (not shown). It increases toward the north, and exceeds 100 m in the northwestern part of the domain. [18] Within the context of mixed layer dynamics, the deepening of the mixed layer represents a downward buoyancy flux, and this buoyancy flux gives the conversion rate between the turbulent kinetic energy and potential energy [Niiler, 1975]. Based on the turbulent kinetic energy budget, Davis et al. [1981] showed that the downward buoyancy flux can be parameterized by g Z 0 h m w 0 r 0 dz ¼ m 0 u 3 * þ m agh m ss m c ðjq 0 j Q 0 Þ; ð1þ 4rC p where w is vertical velocity, r is water density, g is the acceleration due to gravity, C p is specific heat, a is the thermal expansion coefficient, Q 0 is net surface heat flux, and h m is the MLD. The first two terms on the right-hand side of equation (1) denote the energy sources of wind stirring and shear production, respectively, where u * = jt/rj 1/2 and S ¼ R 0 h m ru 0 w z udz. The third term represents the energy sources stemming from surface cooling. [19] Assuming that the flow is unsheared within the mixed layer (i.e., m s = 0), and m 0 = 0.5 and m c = 0.83 [Qiu and Kelly, 1993], the annual mean buoyancy flux is calculated from the NCEP reanalysis surface wind and heat flux data (Figure 3b), and it shows a remarkable correspondence with the observed MLD. Both deep cores of the mixed layer shown in Figure 3a stand out as the maxima of buoyancy flux. The southern maximum is determined mainly by wind stirring associated with trade winds, while the northern one is determined almost equally by monsoonal wind stirring and by cooling winter convection. 4. Heat Budget Equations [20] The equation governing the mixed layer temperature can be expressed ¼ Q 0 q d rc p h m u e rt m u g rt m w entðt m T d Þ ; ð2þ h m where T m denotes the mixed layer temperature and is a good proxy of SST, q d is the downward radiative heat flux across the base of the mixed layer, T d is the water temperature below the mixed layer, u e is Ekman velocity, u g is geostrophic velocity, and w ent represents the entrainment rate of cold water from the below. These five terms in equation (2) will be referred to as temperature tendency, surface thermal forcing, Ekman advection, geostrophic advection, and vertical entrainment, respectively [c.f., Qiu and Kelly, 1993; Qu, 2001] Radiative Heat Flux [21] In computing the downward radiative heat flux of equation (2), Paulson and Simpson s [1977] empirical formula is used q=q 0 ¼ Re z=x 1 þ ð1 RÞe z=x 2; ð3þ

5 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC 35-5 Figure 4. Annual mean (a) surface thermal forcing and (b) horizontal advection and vertical entrainment in 10 8 Cs 1. Contour interval is Cs 1, and areas with positive values in Figure 4a and negative values in Figure 4b are shaded. where q is the radiative flux at depth z, q 0 is the radiative flux at the sea surface, R = 0.62, x 1 and x 2 are attenuation lengths equal to 1.5 and 20 m, respectively, and z is the vertical space coordinate, positive upward with the origin at the sea surface. Given that the mixed layer in the western North Pacific is deep (Figure 3a), q d is generally small compared with Q 0, except in the northwestern subtropical Pacific, where MLD is 20 m in summer and the downward radiative heat flux across the base of the mixed layer becomes noticeable Mixed Layer Temperature [22] The temperature of the mixed R layer is simply calculated in this study as T m ¼ 1=h m h m 0 TðÞdz, z where T(z) denotes temperature at depth z. The temperature jump across the base of the mixed layer, T m T d, is determined by choosing T d at the depth of 10 m below h m. Different values of T m T d have been chosen in literature, depending on the vertical resolution of the data used. Among others, Qiu and Kelly [1993] chose T m T d =1 C, while Qu [2001] chose T d at a depth of 5 m below h m. The vertical resolution of the historical data used here is generally poor, and the strength of the thermocline might be considerably underestimated. For this concern, a T d at a depth of 10 m is preferable to that at a depth of 5 m below h m. Given this selection, the typical value of T m T d is 1 C, ranging from about 0.7 C near the equator to about 1.5 C in the central subtropical gyre. Using a T d at the depth of 5 m below h m the typical value of T m T d will decrease by about 0.2 C; this does not significantly alter the mixed layer heat balance presented in the following sections Horizontal Circulation [23] Horizontal circulation consists of two components: Ekman current and geostrophic flow. Ekman velocity is defined by u e = T k/(rfh m ), where T is the surface wind stress vector constructed from NCEP reanalysis data and f is the Coriolis parameter. The geostrophic flow field used is based on the climatology prepared by Qu and Lukas [2003], which assumes a 1200-db reference level. This climatology resolves the narrow western boundary currents in the western North Pacific better than most, if not all, of the existing climatologies [e.g., Levitus, 1994] Entrainment Rate [24] The entrainment rate, w ent, is determined according to the rate of the mixed layer m /@t, the vertical velocity of water parcel at the base of the mixed layer, w mb, and the horizontal advection of water parcels below the mixed layer, u rh m [e.g., Williams, 1989; Qu, 2001], i.e., w ent m þ w mb þ u rh m if þ w mb þ u rh m > w ent ¼ 0 otherwise: Here w mb = w emb + w gmb and u rh m = u e rh m + u g rh m and the subscript e and g denote Ekman and geostrophic component, respectively. 5. Annual Mean Heat Balance [25] Heat budget terms of equation (2) were calculated from the pentad mean data, and these values were then averaged to produce the annual mean values. Both horizontal advection and vertical entrainment have a singularity at the equator, where f = 0. So our following discussion will focus on a smaller area, which is 5 away from the equator, namely between 120 and 160 E and between 5 and 30 N. In Figure 4a, we see that the surface thermal forcing in the western North Pacific is dominated by a pattern that consists of three regimes. One is located southeast of Mindanao, and the annually integrated contribution of the surface thermal forcing (Q 0 q d )/rc p h m to T m in this area is positive (> Cs 1 ), that is, the ocean gains heat from the atmosphere on the annual average. The second regime is along the ð4þ

6 35-6 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC western boundary of the subtropical Pacific, where atmosphere cools the ocean at a rate exceeding Cs 1. It is interesting that the annually integrated surface thermal forcing warms (> Cs 1 ) the mixed layer in a large part of the interior subtropical gyre, namely north of 25 N between 130 and 160 E, despite the negative annual mean surface heat flux in the region (Figure 2b). Similar phenomenon has been reported in the Kuroshio Extension region [Qiu and Kelly, 1993], and this apparent paradox arises due to the changes in MLD which allow heat gained from the atmosphere to warm the shallow mixed layer in summer more rapidly than that released to the atmosphere to cool the deep mixed layer in winter. [26] For the long-term mean, the surface thermal forcing should be counterbalanced by horizontal advection and vertical entrainment, because annual equilibrium of T m requires the m /@t to be zero (equation (2)). This is essentially true with the mixed layer heat balance in the western North Pacific (Figures 4a and 4b). In particular, the surface heating southeast of Mindanao is primarily counterbalanced by the vertical entrainment (Figure 5a) associated with the Mindanao Dome [Masumoto and Yamagata, 1991; Tozuka et al., 2002]. Along the western boundary of the subtropical gyre, the Kuroshio brings warm water toward higher latitudes (Figure 6a), accomplishing a warming advection of the upper ocean (Figure 5b) that counterbalances most of the surface cooling by the atmosphere (Figure 4a). In the interior subtropical region, a significant part of the surface thermal heating is counterbalanced by vertical entrainment (Figures 4b and 5a), presumably as a result of the deepening of MLD from late fall to early spring (discussed in section 6). [27] In addition to the three regimes described above, there is a broad region in the latitude band roughly between 15 and 25 N, where the atmosphere slightly cools the ocean at a rate of > Cs 1. This cooling effect on T m is mainly counterbalanced by Ekman advection (Figure 5c). With the prevailing easterly winds, the annual mean Ekman current in the western North Pacific is predominantly northward (Figure 6b); its velocity decreases from 10 cm s 1 near the equator to near zero at about 25 N. Meridional temperature gradient is weak in the tropical region but becomes significant at higher latitudes. As a result, a warming Ekman advection exists roughly between 15 and 25 N, with a maximum of C s 1 around 20 N (Figure 5c). In the interior western North Pacific, geostrophic flow tends to follow isotherms (Figure 6a), and the effect of geostrophic flow on T m is generally small. 6. Seasonal Heat Balance [28] The high correlation between temperature m /@t and surface heat flux Q 0 q d shows that surface heat flux is the primary controlling factor for the gross seasonal cycle of the mixed layer temperature (Figure 7a). In a large part of the western North Pacific, the correlation coefficient between temperature tendency and surface heat flux is >0.75 with its maximum exceeding 0.90 in the northeastern part of the domain. This correlation coefficient increases if the seasonal variation of MLD is included (Figure 7b), and it increases even further if horizontal Figure 5. Annual mean (a) vertical entrainment, (b) geostrophic advection, and (c) Ekman advection in 10 8 Cs 1. Contour interval is Cs 1. Values smaller than Cs 1 in magnitude are not shown. advection and vertical entrainment are also included (Figure 7c). Adding the contribution from horizontal advection and vertical entrainment better explains the temperature tendency, particularly, in the region between 15 and 20 N, where the correlation coefficient exceeds 0.9 and can be as high as 0.95 (Figure 7c).

7 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC 35-7 Figure 6. Annual mean (a) geostrophic flow and (b) Ekman current superimposed with mixed layer temperature ( C). [29] Ideally, the correlation between temperature tendency and the sum of surface thermal forcing, horizontal advection, and vertical entrainment should be one. But in reality, due to the lack of sufficient data, our estimates of heat budget components are not accurate, and as a consequence, the correlation is always less than one (Figure 7c). Neglecting eddy heat flux is probably another reason for the lower correlation, and this is particularly true in the region southeast of Mindanao, where eddy activities associated with the Mindanao Dom are extremely strong SST Distribution Associated With the Onset of the Summer Monsoon [30] The onset of the summer monsoon is the most significant phenomenon of the regional climate. It is manifested by the start of rainy season and frequent convective activity and is accompanied by dramatic changes in the large-scale atmospheric circulation [Wu and Wang, 2001]. Over the western North Pacific, previous studies [e.g., Wang, 1994; Murakami and Matsumoto, 1994] have shown that there is a distinct northeastward march of the onset of the summer monsoon from about 120 to 160 Eat10 20 N. Wu and Wang [2001] further indicated that this northeastward march follows a northeastward migration of warm SST, suggesting that the seasonal SST change in the western North Pacific plays an important role in the northeastward march of the summer monsoon onset. Here we see that warm SST (>29 C) first appears in the southwestern Philippine Sea in late May, penetrates northeastward in June, and reaches its northeasternmost position in early July (Figure 8). It remains at almost the same position until late July when the summer monsoon is fully developed. [31] The presence of warm SST can increase the surface air temperature and humidity through enhancing surface heat flux, providing favorable background for convective instability. It can also decrease local pressure that, in turn, can induce low-level convergence. Wu and Wang [2001] interpreted the northeastward march of the onset of the summer monsoon as a result of the northeastward progress of high convective instability, low-level convergence, easterly vertical shear, and monsoon trough, all of which are related to the northeastward migration of warm SST noted above Monsoon Feedback on the SST [32] The onset of the summer monsoon can alter the upper layer thermal structure and eventually set down the SST in the western North Pacific. In addition to the enhanced evaporation and cloud blocking of solar radiation, as previously noted by Wu and Wang [2001], ocean dynamics can also play a role in the decrease of SST. As an example, Figure 9 shows the mixed layer heat balance in pentad 44 (4 8 August), when the summer monsoon approaches its maximum strength. Here we see a dramatic decrease of SST (@T m /@t < Cs 1 ) in the Philippine Sea. Although a core of weak surface thermal forcing (< Cs 1 ) does exist in the region (Figure 9b) in response to the enhanced evaporation and cloud blocking of solar radiation [Wu and Wang, 2001], it cannot explain the negative temperature tendency during that period of time (Figure 9a). [33] To further address this problem, we have calculated all heat budget components of equation (2). In summer, horizontal temperature gradient is weak (Figure 10), and as a result, Ekman advection is generally small in the tropical western Pacific. Ekman advection is significant (> Cs 1 ) only at higher latitudes (Figure 11a), where southeasterly winds drive surface water to flow right in the direction of temperature gradient (Figure 10b). Geostrophic advection is negligible (not shown) for the same reason as discussed in section 5. So vertical entrainment appears to be the only process that cools down the SST in the tropical western Pacific (Figure 11b). Adding the contribution of vertical entrainment partially explains the negative temperature tendency during that period of time (Figure 11c). [34] The importance of vertical entrainment in the seasonal variation of SST is also evident in its longitude-time distribution. In Figure 12, we show the time evolution of heat budget components along 17.5 N. Near the western boundary, temperature tendency reverses its sign from

8 35-8 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 8. early July. Location of the 29 C isotherm from late May to negative to positive in late February (Figure 12a). After this it increases toward its seasonal maximum (> Cs 1 ) around the end of April, and changes its sign from positive to negative in early July. The northeast-southwest orientated contours of temperature tendency from early July to late September clearly demonstrate the eastward propagation of the onset of summer monsoon. The surface thermal forcing bears essentially a similar pattern to the temperature tendency (Figure 12b). A careful inspection, however, reveals that the atmosphere heats the ocean from mid-march to early October, and starts to cool it almost 3 months after SST starts to decrease near the western boundary. This phase difference is significantly narrowed, if contribution from ocean dynamics (mainly vertical entrainment) is included (Figure 12c). Figure 7. Linear correlation coefficient (a) between temperature m /@t and net surface heat flux, Q 0 q d (b) between temperature tendency and surface thermal forcing, (Q 0 q d )/rc p h m, and (c) between temperature tendency and sum of surface thermal forcing, horizontal advection, and vertical entrainment. The light and dark shadings denote areas where the correlation coefficient is larger than 0.8 and 0.9, respectively Vertical Entrainment [35] Vertical entrainment near the western boundary is effective at cooling from early June to late January (Figure 13a), and this explains much of the temperature tendency during the summer monsoon. The dominant term among those of equation (4) at latitudes between 15 and 20 N is the change in MLD. From early February to late May, MLD shoals against the diminishing of winter monsoon (Figure 13b). No vertical entrainment occurs during that period of time, and as a result, heat gained from the atmosphere is trapped in a shallow surface mixed layer, leading to a dramatic increase in SST. The onset of summer monsoon marks the end of this warming condition. MLD starts to deepen in late May, as a result of enhanced wind stirring, and the corresponding vertical entrainment arrests further SST increase due to the incoming surface heat flux. Local Ekman pumping, though somewhat small in magnitude, also contributes to this cooling process during the period from July to December (Figure 13c) Area-Averaged Heat Balance [36] Averaging over the central area of the summer monsoon ( E, N) shows that surface thermal forcing is by far the most important process that

9 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC 35-9 Figure 9. Pentad mean (a) temperature tendency and (b) surface thermal forcing in 10 8 Cs 1 during the period from 4 to 8 August (pentad 44). Areas with negative values are shaded. determines the seasonal variation of mixed layer temperature, but the effect of ocean dynamics is not negligible (Figure 14a). The phase offset between temperature tendency and the sum of surface thermal forcing, horizontal advection, and vertical entrainment is quite consistent, except for a phase difference of about 1 month. The discrepancy in magnitude (called residual below) between the two terms is largest in spring (> Cs 1 ) and falls below Cs 1 during most of the following seasons. In addition to eddy heat fluxes, the uncertainties in the NCEP reanalysis surface heat flux and the temperature climatology can also contribute to this residual (see the Appendix A). [37] Horizontal advection can be further decomposed into Ekman and geostrophic advection (Figure 14b). Ekman advection is the primary heating process in winter, which counterbalances as much as 25% of the surface cooling by the atmosphere, and it becomes negligibly small as horizontal temperature gradient decreases in summer. Vertical entrainment is effective at cooling in summer, while geostrophic advection is negligible throughout the year. [38] As already shown in Figure 7, adding the contribution from ocean dynamics better explains the temperature tendency in the central area of the summer monsoon. The correlation coefficient in this small area between temperature tendency and surface thermal forcing is 0.83 (Figure 14), and it reaches 0.94 if the contribution of ocean dynamics is included. 7. Summary and Discussion [39] This study provides a detailed description of the processes contributing to the mixed layer heat balance in the western North Pacific, using all existing temperature Figure 10. Pentad mean (a) geostrophic flow and (b) Ekman current superimposed with mixed layer temperature ( C) during the period from 4 to 8 August ( pentad 44). Contour interval is 0.4 C and an additional contour of 29.1 C is shown.

10 35-10 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 11. Same as Figure 9 except for (a) Ekman advection, (b) vertical entrainment, and (c) surface thermal forcing plus horizontal advection and vertical entrainment. data combined with NCEP reanalysis surface wind and heat flux. Although the data come from different sources, the mixed layer heat budget is balanced reasonably well. The results are summarized as below. [40] The analyses reveal that the annual mean surface thermal forcing is significant (> C s 1 in magnitude) mainly in three regions. One is in the Mindanao Dome, where the incoming surface heat flux is balanced by the vertical entrainment caused by Ekman pumping. The second occurs in the Kuroshio region, where much of the heat advected by the Kuroshio is released to the atmosphere. The third is located in the central subtropical gyre where the annually integrated surface thermal forcing is balanced primarily by the vertical entrainment associated with the deepening of the mixed layer. [41] In a large part of the interior ocean, roughly between 10 and 20 N, where the surface winds are predominantly easterly, Ekman advection is mainly responsible for balancing the annual mean outgoing surface heat flux. Geostrophic flow (i.e., the NEC) in this region basically follows the east-west orientated isotherms and thus cannot contribute significantly to the annual mean heat budget. [42] Entrainment cooling is generally weak in the western North Pacific. Godfrey and Lindstrom [1989] suggest that turbulent mixing in the upper ocean of the western Pacific is only on the order of 10 W m 2 near the equator and even less at higher latitudes. The weak turbulent mixing in the western Pacific is supported by the present analysis. For the long-term mean, cooling by entrainment is more than Cs 1 in only a small area southeast of Mindanao (Figure 5). (The contribution of 10 W m 2 to the mixed layer temperature is about Cs 1 at a typical MLD of 60 m.) All the heat budget components are small in the tropical western Pacific, however, and thus vertical entrainment can still play a role in maintaining the annual mean SST, as suggested by previous studies [e.g., Niiler and Stevenson, 1982; Godfrey et al. 1991; Qu et al., 1997]. [43] Although surface thermal forcing is the primary factor determining the seasonal variation of SST, ocean dynamics also plays a role. Adding horizontal advection and vertical entrainment better explains the seasonal variation of SST. In the region where the summer monsoon prevails (say between 10 and 20 N), the correlation between the temperature tendency and the full set of terms on the right-hand side of equation (2) is significantly enhanced over the correlation between the temperature tendency and surface thermal forcing alone. [44] Observations in the western Pacific show that from late May to early July the SST (>29 C) increases in a northeastward direction, coinciding with the northeastward march of the summer monsoon [e.g., Wang, 1994]. This change in SST has been attributed to cloud-radiation and wind-evaporation feedback [e.g., Wu and Wang, 2001; Wu, 2002]; the present study shows that this northeastward movement of warm SST results in part from upper ocean dynamics. From late spring to early summer, when MLD shoals toward its seasonal minimum, heat gained from the atmosphere is trapped in a shallow mixed layer. This leads to a rapid increase in SST and provides favorable conditions for the onset of the summer monsoon. The onset of the summer monsoon, though, marks the end of this warming process, and the corresponding deepening of mixed layer enhances the cooling by vertical entrainment that eventually cools down the SST. [45] The importance of vertical entrainment in the seasonal variation of SST in the western North Pacific may

11 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 12. Time-longitude distribution of (a) temperature tendency, (b) surface thermal forcing, and (c) sum of surface thermal forcing, horizontal advection, and vertical entrainment in 10 8 Cs 1 along 17.5 N. Areas with negative values are shaded. Figure 13. Same as Figure 12 except for (a) vertical entrainment (10 8 Cs 1 ), (b) rate of mixed layer deepening (10 6 ms 1 ), and (c) Ekman pumping velocity (10 6 ms 1 ).

12 35-12 QU: MIXED LAYER HEAT BALANCE IN THE WESTERN NORTH PACIFIC Figure 14. Time series of (a) temperature tendency (heavy solid curve), surface thermal forcing (light solid curve), sum of horizontal advection and vertical entrainment (light short dashed curve), sum of surface thermal forcing, horizontal advection, and vertical entrainment (heavy dashed curve), and the residual defined as the difference between temperature tendency and the sum of surface thermal forcing, horizontal advection, and vertical entrainment (light solid curve with open cycle) and (b) Ekman advection (light solid curve), geostrophic advection (light long dashed curve), and vertical entrainment (light short dashed curve) averaged in the region E, N. Unit is 10 8 Cs 1. provide a useful hint in understanding the dynamics of ENSO and its connection with the Asian monsoon. Previous studies have shown that the off-equatorial ocean-atmosphere interaction in the western Pacific is a key process in determining the ENSO phase transition [Wang et al., 1999, 2000]. A separate study that uses results from an ocean general circulation model is now underway to further investigate the exact role of ocean dynamics in this process. Appendix A: Error Analysis [46] The uncertainty of the mean temperature field is measured by standard deviation (STD) information also obtained during averaging. The spatial distribution of temperature STD, though not shown, are chiefly east-west orientated. Its typical value at the sea surface is 0.5 C, ranging from about 0.3 C near the equator to about 0.8 C at higher latitudes. The standard errors, defined as the STDs divided by the square root of the number of measurements, are smaller by a factor of 2 3, based on 5 10 measurements in each grid bin. In addition to the bad quality of the data (discussed in section 2), these uncertainties can be attributed to interannual variability and eddy activities. [47] The uncertainty of the velocity field used for this study is not available. The difficulty lies in the fact that the dynamic height standard deviations cannot be evaluated from individual profiles directly because many are shallower than the reference level Qu and Lukas [2003] chose (i.e., 1200 db). The only advection uncertainty we can 0 evaluate here is due to u rt m where u represents the mean velocity and T m represents the uncertainty in the mixed layer temperature. Owing to the limitation of climatological data, the present study neglects all the processes associated with meso-scale eddies whose decorrelation scale is up to 1000 km in the tropical western Pacific [Qu et al., 2002], but rather focuses on the large-scale (>1000 km) phenomena. Based on 5 10 independent profiles in each grid bin and assuming that the adjacent bins are uncorrelated, a typical horizontal velocity of 0.1 m s 1 combined with a typical temperature standard deviation of 0.5 C can be roughly translated into an advection uncertainty of Cs 1 across a distance of 1000 km. [48] Another source of uncertainty in the mixed layer heat balance is due to surface heat flux. Recent studies have shown that the NCEP reanalysis surface heat flux significantly underestimates the ocean heat gain in the subduction region of the Northeast Atlantic [e.g., Josey, 2001]. The possibly large uncertainty of this data set in the tropical western Pacific was also noticed (e.g., Toole et al., submitted manuscript, 2002), but to the best of our knowledge, no accurate estimate has been published. To test the sensitivity of our results to different surface heat flux products, this study also included Oberhuber s [1988] climatology based on the Comprehensive Ocean- Atmosphere Data Set (COADS). The results from these two data sets show remarkable agreement. [49] Although there is no easy way to quantify the uncertainties of all heat budget components, especially the possibly large uncertainty in surface heat flux, the existing estimate shows an uncertainty of heat advection < Cs 1. With this uncertainty, most of the large-scale phenomena (>1000 km) presented in this study are representative. [50] Acknowledgments. This research was stimulated by conversations with B. Wang. The author is grateful to J.P. McCreary, H. Mitsudera, T. Jensen, T. Miyama, Y.-Y. Kim, and H.-W. Kang for useful communication on the present topic, J. Toole and an anonymous reviewer for constructive suggestions, and G. Speidel for careful editing of the earlier manuscript. Supports by National Science Foundation through grant OCE and by Frontier Research System for Global Change through its sponsorship of the International Pacific Research Center (IPRC) are acknowledged. School of Ocean and Earth Science and Technology (SOEST) contribution 6086, and IPRC contribution IPRC-192. References Davis, R. E., R. deszoeke, and P. P. Niiler, Variability in the upper ocean during MILE, part II, Modeling the mixed layer response, Deep Sea Res., Part A, 28, , 1981.

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