Comparison of Sea Surface and Mixed Layer Temperatures

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 4               Comparison of Sea Surface and Mixed Layer Temperatures

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 7                      Semyon A. Grodsky, James A. Carton, and Hailong Liu

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11                                         April 16, 2008

12                    Submitted to the Journal of Geophysical Research, Oceans

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17   Department of Atmospheric and Oceanic Science

18   University of Maryland, College Park, MD 20742

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20   Corresponding author:

21   senya@atmos.umd.edu

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24   Abstract

25   Mixed layer temperature, TML , and SST are frequently used interchangeably or assumed to be

26   proportional in climate studies. This study examines the historical observational record 1960-

27   2007 for systematic differences between these variables. The results show that globally and time

28   averaged TML is lower than SST by approximately 0.1 oC. TML minus SST is even lower in

29   upwelling zones where abundant net surface warming is compensated for by cooling across the

30   base of the mixed layer. In the upwelling zone of the Equatorial East Pacific this negative TML -

31   SST difference varies out of phase with seasonal SST, reaching a negative extreme in boreal

32   spring when SST is warm, solar radiation is high, and winds are weak. In contrast, on interannual

33   timescales TML -SST varies in phase with SST with small differences during El Niños as a result

34   of low solar heating and enhanced rainfall. On shorter diurnal timescales during El Niños TML -

35   SST differences associated with temperature inversions occur in response to nocturnal cooling in

36   presence of nearsurface freshening. Near surface freshening produces persistent shallow (a few

37   meters depth) warm layers in the northwestern Pacific during boreal summer when solar heating

38   is strong. In contrast, shallow cool layers occur in the Gulf Stream area of the Northwest Atlantic

39   in boreal winter when fresh surface layers developed due to lateral interactions are cooled down

40   by abundant turbulent heat loss. The different impacts of shallow barrier layers on near surface

41   temperature gradients are explored with a one dimensional mixed layer model.

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43   1.         Introduction

44   Although instantaneous thermodynamic fluxes across the ocean-atmosphere interface are

45   affected by the temperature of the near surface ocean (< 1 m), many climate studies identify the

46   vertically average temperature of the ocean mixed layer to be the most relevant parameter or

47   even use a slab mixed layer as a proxy for the ocean [see e.g. Manabe and Stouffer, 1996]. In

48   general we may expect monthly depth average mixed layer temperature to be lower than SST by

49   a few tenths of a degree. This difference reflects the time average effect of the nearsurface

50   suppression of turbulence by daytime warming. In this study we compare historical analyses of

51   SST with contemporaneous temperature and salinity profile observations to identify the

52   conditions giving rise to systematic differences between mixed layer temperature and SST and to

53   identify the regions where this difference is essential.

54

55   SST is a difficult parameter to define exactly because the upper 10 m of the ocean has such

56   complex and variable vertical temperature stratification1. This variation in stratification occurs

57   more frequently under conditions in which the ocean surface fluxes cause gains or loses of heat

58   or freshwater or in situations of strong horizontal exchange. Surface fluxes are responsible for a

59   distinct diurnal cycle in the temperature in the uppermost few meters over wide areas of the

60   ocean when winds are weak and solar heating is strong. [Stuart-Menteth et al., 2003; Gentemann

61   et al., 2003; Clayson and Weitlich, 2007; Kawai and Wada, 2007]. This diurnal cycle is

62   particularly prominent in upwelling areas such as the eastern equatorial Pacific where vertical

63   advection of cool water leads to shallow stratification and thus shallow mixed layers (Deser and

64   Smith, 1998; Cronin and Kessler, 2002). In the warm pool region of the western equatorial

65   Pacific diurnal warming arises because the excess rainfall forms a nearsurface barrier layer of
     1
         See the GODAE Global High Resolution SST Pilot Project at http://www.ghrsst-pp.org/SST-Definitions.html


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66   low salinity water even though the seasonal thermocline is rather deep [Soloviev and Lukas,

67   1997].

68

69   In order to reduce the impact of diurnal effects the UK Met Office HadISST1 utilizes only the

70   night satellite SSTs (available beginning in 1981) and adjusts them to match in-situ

71   measurements collected by voluntary observing ships, drifters, and buoys (Rayner et al. 2003).

72   The NOAA National Climatic Data Center SST extended analysis uses both day and night

73   satellite SSTs only to evaluate the spatial structure of analysis SST while rely on the same in situ

74   observations to adjust their SST analysis to reflect water temperature at an effective depth at ~1-

75   5 m (Smith and Reynolds 2003). A more precise definition of this analysis depth is impractical

76   for either product because of the variety of depths at which the in situ observations are available.

77   Most recently the Global Ocean Data Assimilation Experiment High Resolution SST Project has

78   introduced the concept of ‘Foundation SST’, defined as the temperature at a depth (10m) that is

79   below the depth of the diurnal cycle. But this 10m depth temperature, which generally lies

80   within the mixed layer, has not been measured frequently enough to calibrate the analyses.

81

82   In this study we focus on the difference between the mixed layer temperature and SST provided

83   by historical analyses of Rayner et al. [2003] and Smith and Reynolds [2003]. The mixed layer is

84   defined as the near-surface layer of uniform properties such as temperature and salinity. The

85   presence of weak stratification and the nearness to atmospheric momentum sources give rise to

86   values of the Richardson number consistent with flow instabilities and thus a high potential for

87   turbulent motion. Under conditions where density is primarily determined by temperature de

88   Boyer Montégut et al. [2004] (with a generalization introduced by Kara et al., 2000a) define the




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 89   base of the seasonal mixed layer to be the depth at which temperature changes by 0.2C from its

 90   value at 10m Foundation depth. From this we can define a seasonal mixed layer temperature,

 91   TML , as the vertical average temperature of the mixed layer, which when multiplied by the depth

 92   of the mixed layer and the specific heat of seawater gives the heat capacity of the layer of ocean

 93   in direct contact with the atmosphere on seasonal timescales.

 94

 95   The near surface processes that affect the monthly TML minus SST difference, dT= TML -SST, are

 96   dominated by the integrated effect of diurnal warming. But, a variety of processes including

 97   rain, river discharge, or lateral interactions may produce fresh barrier layers that trap the heat

 98   near the surface by shoaling the penetration depth of wind stirring and nocturnal convection

 99   [Lukas and Lindstrom, 1991; Soloviev and Lukas, 1997]. Moreover stable salinity profiles may

100   permit nocturnal temperature inversions due to radiative cooling [Anderson et al., 1996; Cronin

101   and Kessler, 2002] with magnitudes comparable to those of diurnal warming. Barrier layers are

102   observed over wide ocean areas, in particular, they are produced by abundant rainfall and river

103   discharge in the tropics, by an excess precipitation over the North Pacific, and by lateral

104   exchanges across the western boundary currents [de Boyer Montégut et al., 2007]. In all these

105   areas we also expect significant stratification of near surface layers that affect the difference

106   between TML and SST.

107

108   2.     Data and Methods

109   The mixed layer properties for this study are estimated from individual temperature profiles

110   provided by World Ocean Database 2005, WOD05, [Boyer et al., 2006] for the period 1960

111   through 2004. We use data from the mechanical bathythermographs (MBT), expendable



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112   bathythermographs (XBT), conductivity-temperature-depth casts (CTD), as well ocean station

113   data (OSD), moored buoys (MRB), and drifting buoys (DRB). The final four years of the

114   database contain an increasing number of profiles from the new Argo system (PFL). The Argo

115   profiles for the period through 2007 are obtained from the Argo Project web site. For better

116   characterization of the tropical Pacific region the data provided by the TAO/TRITON moorings

117   [McPhaden et al., 1998] are also used.

118

119   The mixed layer depth (MLD) may be defined in a number of different ways. In this study we

120   use the concept of the isothermal mixed layer depth (MLD) that is evaluated from individual

121   vertical profiles based on the temperature difference from the temperature at a reference depth of

122   10 m [de Boyer Montégut et al., 2004]. This reference depth was shown to be sufficiently deep to

123   avoid aliasing by the diurnal signal, but shallow enough to give a reasonable approximation of

124   monthly TML . Here the isothermal MLD is defined as the depth at which temperature changes by

125   | T | = 0.2oC relative to its value at 10m depth. Following Kara et al. [2000a], the isothermal

126   MLD is defined by the absolute difference of temperature, | T |, rather than only the negative

127   difference of temperature to account for mixed layers with temperature inversions in salt-

128   stratified situations (most common at high latitudes). The mixed layer temperature is evaluated

129   as the temperature vertically averaged above the base of the mixed layer assuming uniform

130   temperature above the reference depth, T ( z  10m)  T ( z  10m) .

131

132   An alternative definition of the mixed layer depth (based on the dynamical stability criterion)

133   defines it as the depth of a density uniform layer. Vertically average temperature of temperature

134   uniform layer is the same as vertically average temperature of density uniform layer if the latter



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135   layer is not deeper than the former (barrier layer). If a density uniform layer is deeper than a

136   temperature uniform layer (density compensation) their average temperatures may be different.

137   Here we follow de Boyer Montégut et al. [2004] and define the mixed layer as a layer vertically

138   uniform in both, temperature and salinity. Hence, the mean mixed layer temperature is the same

139   as the mean temperature of isothermal layer. The latter is referred in this study as the mixed layer

140   temperature, T ML .

141

142   The mixed layer archive and the seasonal and interannual variability of mixed layer properties

143   are described by Carton et al. [2008]. They show that the temperature difference criterion works

144   reasonably well even at high latitudes in the North Atlantic and provide further details on data

145   quality control procedures. After estimating TML at each profile location we then apply

146   subjective quality control to remove ‘bulls eyes’ and bin the data into 2ox2ox1mo bins with no

147   attempt to fill in empty bins.

148

149   SST is provided by Met Office Hadley Centre sea ice and sea surface temperature (HadISST1) of

150   Rayner et al. [2003] and by extended analysis (version 2) of Smith and Reynolds [2003]. Both

151   products are globally complete monthly averaged grids spanning time period beginning the late

152   19-th century onward. HadISST1 combines a suite of historical and modern in situ near surface

153   water temperature observations from ships and buoys with the recent satellite SST retrievals,

154   while the Smith and Reynolds [2003] data is mostly based on in-situ measurements. Neither of

155   these products uses the vertical temperature profiles from WOD05. The SST provided by these

156   archives is based to a large extend on measurements collected by voluntary observing ships and

157   is referred here as bulk SST or simply SST. Data adjustment to measurements taken from a few



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158   meters depth (where the diurnal signal is relatively weak) effectively attenuates but doesn’t

159   eliminate completely impacts of transient near surface processes on bulk SST.

160

161   The local response of the mixed layer to the forcing from the atmosphere is simulated using the

162   one dimensional hybrid mixed layer model of Chen et al. [1994]. This model is based on the

163   Kraus-Turner-type bulk mixed layer physics in which the depth of the mixed layer is determined

164   by a turbulent energy balance equation, while the temperature and salinity of the mixed layer is

165   determined by budget equations forced by surface fluxes and entrainment. These balances are

166   augmented in the Chen et al. formulation by the addition of convection and Richardson Number-

167   dependent mixing. The model is forced by 6-hour surface fluxes provided by the National Center

168   for Environmental Predictions/Department of Energy (NCEP/DOE) Reanalysis-2 of Kanamitsu

169   et al. [2002].

170

171   3.      Results

172   We begin by examining the average dT based on the 1960-2004 WOD05 data set (Figs. 1a,b).

173   Because of the distribution of observations only the Northern Hemisphere is well sampled. On

174   average, T ML is colder than SST by about -0.1oC, with large <-0.4C negative anomalies in the

175   region north of the Kuroshio-Oyashio extension and along the equator in the eastern Pacific, and

176   large >0.4C positive anomalies (temperature inversions) in the Gulf Stream region (the results

177   are similar for the two SST analyses). The equatorial Atlantic shows negative anomalies as well,

178   but not as large as the equatorial Pacific.

179




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180   To illustrate this relationship in the Southern Hemisphere we examine average dT using the

181   spatially more homogeneous Argo profile data set which, however, is mainly restricted to the

182   years 2004-onward (Fig. 1c). The Argo results in the Northern Hemisphere show only a few

183   differences from the distribution of dT based on the WOD05 data set. In the Labrador Sea

184   positive values of dT (indicating nearsurface temperature inversions) are now more evident. In

185   contrast, the subtropical North Atlantic and North Pacific both show negative values in the

186   regions of weak winds where diurnal warming of the nearsurface is a frequent occurrence. In the

187   Southern Hemisphere dT based on Argo shows large negative anomalies at several longitudes,

188   in the South Pacific west of Chile as well as southwest of Australia and South of Cape of Good

189   Hope. We next focus on the Northern Hemisphere patterns because they are evaluated from

190   longer time records then the southern counterparts. To explore the causes of the largest

191   anomalies of dT we next examine in detail the time changes in the three regions in the Northern

192   Hemisphere identified in Fig. 1.

193

194   These three regions are distinguished by persistent shallow near surface stratification due to

195   either upwelling or impact of the barrier layers (nearsurface freshening) that trap warming

196   (cooling) in the near surface. On the other hand, the air-sea interactions are particularly strong

197   over these regions. It is illustrated by climatological maps of the net surface heat gain by the

198   ocean. During the northern winter (Fig. 2a) the turbulent heat loss in excess of 200 Wm-2 occurs

199   over the warm western boundary currents due to strong air-sea temperature contrast and

200   enhanced evaporation over warm SSTs. In northern summer (Fig. 2b) the ocean gains heat in

201   excess of 150 Wm-2 in the northwestern Pacific and over the shelf waters north of the Gulf

202   Stream. In both these areas the local increase of the ocean heat gain is due to a decreased




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203   evaporation over cool SSTs. The ocean also gains heat at a rate exceeding 100 Wm-2 in the

204   eastern equatorial Pacific cold tongue (Fig. 2b) due to abundant solar radiation and relatively

205   weak latent heat loss. In the cold tongue the heat gain is compensated for by entrainment cooling.

206   In the near surface it produces remarkable magnitudes of diurnal warming. We shall next analyze

207   the origins of persistent shallow stratifications in these three regions.

208

209   3.1    Eastern Equatorial Pacific

210   The equatorial Pacific thermocline shoals eastward in response to annual mean easterly winds

211   that along with entrainment cooling form a tongue of cool water in the east. Here, in the cold

212   tongue, the ocean gains heat from the atmosphere in excess of 100 Wm-2 (Fig.2b) that is

213   compensated for by entrainment cooling. In response to this surface heat flux the nearsurface

214   ocean develops substantial diurnal warming of SST, in excess of 0.2C in time average [Deser

215   and Smith, 1998]. Here dT averages approximately -0.4oC (Fig. 3a) with more negative values

216   ( T ML <SST) in March when SST reaches its monthly maximum and diurnal warming is large

217   (Fig. 3b) [Cronin and Kessler, 2002]. In contrast, on interannual timescales dT is weak

218   ( TML  SST) when El Nino warms SST, the mixed layer deepens, solar radiation decreases and

219   freshwater input increases, and dT has negative extreme during the La-Niñas when the mixed

220   layer shoals and atmospheric convection shifts westward [Cronin and Kessler, 2002; Clayson

221   and Weitlich, 2005]. This relationship is most clear after the early 1980s, as the data coverage

222   increases.

223

224   In order to understand the causes of the seasonal and interannual relationships we examine

225   conditions at the Tropical Ocean Atmosphere/TRITON mooring at 0N, 140W for the seven years



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226   1995-2001 encompassing the 1997-98 event (Fig. 4a). We focus on 0N, 140W location where

227   the records are continuous during the event. At this location 1m temperature, a proxy for SST,

228   increases by 5C during 1997 and then decreases by nearly 7C in mid-19982. Coinciding with the

229   drop in 1m temperature is a substantial development of negative dT meaning that the mixed

230   layer has developed some nearsurface temperature stratification. The negative values of dT are

231   even more striking in 1999 and 2000 when SST increases during January-March as part of the

232   climatological seasonal cycle at this location phase with interannual variation of dT .

233

234   To identify the mechanisms giving rise to differences in seasonal and ENSO changes in dT we

235   examine a one dimensional mixed layer model simulation beginning with homogeneous initial

236   conditions (Fig. 4b). The model responds seasonally to weakened winds in boreal spring (Fig.

237   4d) with increased near-surface stratification ( dT <0) as observed. The conditions arising during

238   the onset of El Nino such as occurred during the first half of 1997 are somewhat different.

239   During those months the winds also weakened, but solar heating decreased (Fig. 4c) and

240   freshwater input increased (Fig. 4d) as a result of the eastward shift of convection. The decrease

241   in the ocean heat gain due to decreased solar heating is accompanied by increased latent heat loss

242   due to warmer SST (Fig. 4c). The result is weakening values of dT followed in the summer and

243   fall by occasional temperature inversions. In mid-1998 through early 1999 as El Nino

244   transitioned into cooler La Nina conditions, the nearsurface again becomes strongly stratified due

245   to enhanced solar heating and weaker latent heat loss and resulting diurnal warming of the

246   nearsurface.

247

      2
          TAO/TRITON moorings measure SST at z=1m. Time mean difference of T1m from HadISST1 at 0N, 140W is -
      0.3C while time correlation is 0.96.


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248   Intermittent temperature inversions (0.2-0.5C cooler SSTs) are evident in observations (Fig. 4a)

249   and simulations (Fig. 4b). They are associated with nocturnal cooling of shallow freshwater

250   lenses produced by enhanced rainfall (Fig. 4d). Stable salinity stratification (barrier layer)

251   produced by local rainfall captures the nocturnal convection in the near surface layer until the

252   cooling or wind stirring is strong enough. If the freshwater surface flux is set to zero, the one

253   dimensional model doesn’t simulate temperature inversions [see also Anderson et al., 1996].

254

255   As we have seen the stable salinity stratification produced by local rainfall may impact

256   significantly the near surface temperature stratification. An alternative mechanism of barrier

257   layers formation is associated with the lateral interactions. In particular, in the equatorial Pacific

258   near the dateline, salt and warm water can be subducted under the western Pacific warm fresh

259   water to form barrier layers [Lukas and Lindstrom, 1991]. This advection mechanism (that is not

260   in a one dimensional model physics) may be effective near the frontal interfaces and contribute

261   to temperature inversions during the seasons when the ocean loses heat.

262

263   3.2    Gulf Stream

264   We next consider dT in the western North Atlantic where time mean positive values

265   (temperature inversions) lie along the Gulf Stream path (Fig. 1). This regional anomaly may

266   result from differences in spatial interpolation of T ML and bulk SST. However, we’ll next see in

267   Fig. 3d that dT in the Gulf Stream region displays noticeable seasonal variations suggesting that

268   physical processes play a role as well.

269




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270   Spatial patterns of dT during the two contrasting seasons are illustrated in Fig.5 using CTD and

271   Argo profiles. In distinction from Fig. 1 where T ML is compared with bulk SSTs from climate

272   archives, the SST in Fig. 5 is taken from the same profiles as TML . This eliminates contribution

273   due to the difference in spatial interpolation of SST and T ML . In summer dT is less than zero

274   (Fig. 5a) due to the formation of a shallow a few meter deep diurnal warming trapped in a 10-

275   30m shallow fresh layer (Fig. 6a). This shallow barrier layer limits the depth of nocturnal

276   convection and mechanical stirring above the base of halocline and thus separates the shallow

277   near surface warm layer from the seasonal mixed layer. In contrast a near-surface temperature

278   inversion is largest in the cold seasons (Fig. 5b) when heat loss is very large (Fig.2a, also see

279   Dong and Kelly, 2004). An examination of the spatial structure of dT during the winter months

280   (Fig. 5b), indeed, shows large inversions, frequently exceeding 1C along the path of the Gulf

281   Stream, while SST is close to T ML in this area in summer. Variations of dT are similar if an

282   alternative, gradient-based MLD is used (Fig.2c).

283

284   Seasonal variations of dT in the Gulf Stream region occur in accord with the seasonal variations

285   of the net surface flux that displays a peak in heat loss over the warm Gulf Stream in winter (Fig.

286   2a) and strong warming over the cool shelf water in summer (Fig. 2b). In distinction from the

287   equatorial Pacific where interannual dT significantly correlates with local SST, these values are

288   weakly correlated in the Gulf Stream area (Fig. 3c). This weak correlation reflects probably an

289   impact of interannual migration of the Gulf Stream front that affects both the box averaged SST

290   and dT .

291




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292   Winter mixed layer is occasionally warmer than SST in the Gulf Stream sector (Fig. 5b).

293   Associated temperature inversions are aligned along the Gulf Stream northern wall suggesting

294   that lateral cross-frontal interactions between water masses may play a role. Subduction of warm

295   and salt water carried by the Gulf Stream below the cold and fresh shelf water produces a

296   shallow cold layer (Fig. 6b) that is further cooled down as the ocean losses heat. The temperature

297   inversion doesn’t overturn until the stable salinity stratification is overcome by instability

298   introduced by near surface cooling or wind stirring. In distinction from numerous occurrences of

299   the near surface warm layers in summer (Fig. 5a), the near surface cold layers are observed only

300   sporadically in winter (Fig. 5b). In fact, they are destroyed by transient storms that stir these

301   shallow density compensated layers. Despite their intermittence, the near surface temperature

302   inversions might impact wintertime infrared SSTs because passing storms that eventually destroy

303   the inversions are normally associated with the cold air outbreaks and significant convection

304   cloudiness.

305

306   Discussions above emphasize impacts of salinity on the nearsurface temperature stratification.

307   Next the temperature response to the presence of the near surface salinity gradients (occurring in

308   the Gulf Stream area) is explored with one-dimensional mixed layer model (Fig. 7). To contrast

309   the impact of salinity, the twin runs are compared. Each pair of model runs is forced by the same

310   fluxes but differs in initial conditions. The first (control) run starts from the vertically

311   homogeneous temperature and salinity while the initial salinity profile for the second run has

312   salinity decreasing toward the surface in the upper 20 m at a rate of 0.1 psu m-1 (in accord with

313   observations in Fig. 6).

314




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315   Fig. 7b illustrates simulations during the warm season. It displays the difference in temperature

316   between the two runs that evidents an impact of the near surface freshening. In the presence of

317   stabilizing salinity gradient the diurnal warming is stronger during the first day of simulations

318   (Fig. 7b), but is surprisingly similar during the second day when it is limited by the shear

319   instability of diurnal currents. Relative warming in the upper 20 m is even stronger as wind

320   strengthens. This is explained by slower deepening of the mixed layer and weaker entrainment

321   cooling in salinity stratified case. Although the one-dimensional mixed layer model simulates

322   warmer near surface temperature in salinity stratified case, the simulated temperature

323   stratification in the upper 10 m column doesn’t exceed a few tenth of degree in contrast with

324   observations (Fig. 6a). This is explained in part by relatively short (only a few days long) run as

325   well as by limitations of the model. If a strong (  ~ 1 day) relaxation of salinity to its initial

326   conditions is introduced (to account indirectly for mechanisms producing shallow halocline) the

327   temperature gradient in the upper 10m amplifies up to 1 oC but never reaches values shown in

328   Fig. 6a.

329

330   In winter the mixed layer model simulates 1oC colder mixed layer in salt stratified case than in

331   control run (Fig. 7d). The difference is due to the stably stratified halocline that limits

332   penetration depth of wind stirring. In turn, the shallower mixed layer is the faster it cools down

333   due to net surface heat loss. Although the anomalous cooling of 1oC compares well with

334   observations (Fig. 6b), the simulated mixed layer is relatively deep. Therefore, the stratification

335   is weak in the upper 10 m in distinction from observations. This suggests again that lateral

336   interactions (missing by one-dimensional model) are important for establishing winter




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337   temperature inversions in the region, while the net surface heat loss further amplifies existing

338   anomalies.

339

340   3.3     Northwestern Pacific

341   Salinity in the Northwestern Pacific decreases towards the surface. This stable halocline is

342   produced by annual mean excess of precipitation over evaporation north of 30N and is

343   maintained by upward vertical pumping driven by cyclonic wind curl [Kara et al., 2000b].

344   Although the regional precipitation peaks in winter, the near surface freshening persists year

345   around. In summer when the ocean heating is particularly strong (Fig. 2b), the shallow stably

346   stratified halocline localizes the ocean heat uptake in the nearsurface layer (Fig.1) by limiting the

347   penetration depth of wind stirring and nocturnal convection. In distinction from the Gulf Stream

348   region where shallow warm layers develop mostly in the cold sector of the front, the shallow

349   warm layers are observed randomly in the Northwestern Pacific (Fig.5c). They are not destroyed

350   by nocturnal convection (see sample profile taken at 20:30 local time, Fig. 8a). Meridional

351   variations of dT follow the meridional variations of net surface heating and are similar if

352   different a gradient-based definition of the mixed layer depth is used (Fig. 2d). Occasional SST

353   inversions seeing in Fig. 5c are associated with nocturnal cooling of freshwater lenses (Fig. 8b).

354

355   Shallow stratified layers observed in the Northwestern Pacific in summer are destroyed by winter

356   storms (Fig. 5d and Fig. 3f). Despite similarly strong heat loss over the warm western boundary

357   currents (Fig. 2a), the winter SST inversions are not observed in the Kuroshio region in

358   distinction from the Gulf Stream region (Fig.1). This may be linked to the differences in spatial

359   patterns of salinity. In fact, the spatial gradients of salinity (that are vital for producing the




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360   temperature anomalies) are significantly weaker in the Northwestern Pacific in comparison with

361   the Northwestern Atlantic [see e.g. Antonov et al., 2006].

362

363   4.     Summary

364   This study compares the magnitudes of two ocean temperature variables frequently used in

365   climate studies, mixed layer temperature and bulk SST as represented by the widely used

366   analyses of Rayner et al. [2003] and Smith and Reynolds [2003]. Mixed layer temperature is

367   defined as the vertical average temperature above the mixed layer base, and the depth of the base

368   here is defined following Kara et al. [2000a] and de Boyer Montégut et al. [2004] as a function

369   of the temperature difference relative to 10m temperature. Our analysis shows that areas with

370   shallow temperature stratification such as upwelling zones frequently have significant

371   differences between mixed layer temperature and SST. Shallow temperature stratification also

372   occurs in regions of near surface freshening (barrier layers) which limits the depth of convection

373   and wind stirring. In both cases shallow stratification occurs in zones of strong air-sea heat

374   exchange. In the northern hemisphere the local peaks of heat gain by the ocean are observed in

375   local summer over the areas of equatorial cold tongues and over the areas of cold SSTs to the

376   north of the Kuroshio extension front and the Gulf Stream north wall. While in winter the ocean

377   loses much heat over the warm SSTs of the western boundary currents.

378

379   We examine the temporal relationship between SST and T ML in the Equatorial East Pacific

380   where abundant net surface warming is compensated for by cooling across the base of the mixed

381   layer. Here T ML is persistently cooler than SST by approximately -0.4oC. On seasonal time

382   scales, it has a negative extreme during the boreal spring warm season when winds are weak. In



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383   contrast, on interannual timescales the magnitude of dT = TML -SST increases during La Ninas

384   and weakens during El Niños as a result of increases/decreases in solar radiation and

385   decreases/increases in precipitation. Increased precipitation during El Niños produces freshwater

386   stratified barrier layers leading to nocturnal cooling.

387

388   In the subtropics negative values of dT are found in the Gulf Stream area of the western North

389   Atlantic. In summer the shallow warming in excess of 1 oC develops above the cool shelf waters

390   to the west and north of the Gulf Stream where the ocean gains heat at a rate of 150 Wm-2. The

391   presence of nearsurface freshening prevents the nighttime destruction of this shallow warm layer.

392   In contrast, during winter the near surface layer within the Gulf Stream itself has an inverted

393   temperature structure (the time averaged dT=0.6C) as the result of strong surface cooling in the

394   presence of a nearsurface barrier layer.

395

396   Another region where the salinity stratified barrier layers are present is the Kuroshio extension

397   region of the Northwest Pacific. Here the barrier layer is produced due to excess of precipitation

398   accompanied by upward Ekman pumping preventing the vertical exchange of this freshwater. As

399   in the case of the Gulf Stream region, the ocean gains heat in the summer at a rate of 150 Wm-2

400   producing a warm surface layer during the day which has the time averaged dT =-0.5 oC. In

401   winter, T ML and SST match in this region.

402

403   One of the persistent issues in coupled atmosphere-ocean general circulation models is tendency

404   to develop cold biases in the eastern equatorial Pacific [Davey et al. 2002]. However the surface

405   temperature of such models is actually more analogous to mixed layer temperature since the



                                                                                                          17
406   uppermost ocean gridpoint is well below the ocean surface and diurnal processes are generally

407   neglected. Thus, any systematic differences in SST and T ML are likely to be reflected in the

408   evaluation of model SST bias. Indeed, Danabasoglu et al. [2006] have shown that adding the

409   diurnal cycle to the daily mean incoming solar radiation does warm the model eastern equatorial

410   Pacific SST and shoals the ocean boundary layer in better agreement with SST observations.

411   Even greater improvements to model SST estimates seem possible if the nearsurface

412   stratification of temperature and salinity can be more accurately represented.

413

414

415   Acknowledgements. We gratefully acknowledge the Ocean Climate Laboratory of the National

416   Oceanographic Data Center/NOAA, under the direction of Sydney Levitus for providing the

417   database upon which this work is based. Mixed layer temperature estimate based on the

418   Lorbacher et al. [2006] approach has been provided by Dietmar Dommenget, IFM-GEOMAR.

419   Support for this research has been provided by the National Science Foundation (OCE0351319)

420   and the NASA Physical Oceanography Programs.

421




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488




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489

      22
490   Figure 1. Time mean difference, dT , of mixed layer averaged temperature, TML , and bulk SST.
491           (a) T ML from WOD05 and bulk SST from HadISST1, (b) T ML from WOD05 and SST
492           from Smith and Reynolds extended v.2. In (a) and (b) grid points with less than one year
493           of data aren’t shown. (c) T ML from Argo floats and bulk SST from HadISST1 at grid
494           points with at least 6 months of data.  dT  is the global and time mean difference.
495




                                                                                                     23
496
497
498   Figure 2. (Left) Climatological net surface heat flux for (a) January, (c) August (solid contours
499           indicate heat gain by the ocean). Boxes show the same areas as shown in Fig.1a. (Right)
500           Mixed layer minus SST temperature difference, dT , zonally averaged over longitude
501           belts shown in the left panels. Solid lines show result of this study (against bottom x-
502           axis) while dashed lines show results based on depth estimates of Lorbacher et al. [2006]
503           (against top x-axis) that is based on the gradient-based definition of the mixed layer
504           depth.
505
506




                                                                                                    24
507




508
509
510   Figure 3. (a), (c) ,(e) Time series of annual running mean box-averaged dT , standard deviation
511           of dT (shading), and anomalous SST for the three boxes, equatorial east Pacific, Gulf
512           Stream, and northwestern Pacific. (b), (d), (f) Seasonal cycle of box-averaged dT and
513           SST based on HadISST1 data. Time series combine dT evaluated from WOD05 data
514           through 2004 and Argo data afterwards.
515




                                                                                                    25
516
517   Figure 4. (a) Time series of 1m temperature, T1m , and mixed layer temperature gradient, dT ,
518           from (a) TAO/TRITON mooring at 0N, 140W, (b) mixed layer model (MLM). (c)
519           monthly running mean short wave radiation (SWR) and latent heat loss (LHTFL), (d) 6-
520           hour precipitation (PRECIP) and monthly zonal wind stress (TAUX).
521



                                                                                                  26
522
523   Figure 5. dT during June-August (JJA) and October-March (ONDJFM) evaluated from
524           individual CTD and Argo profiles in (a,b) Gulf Stream area, (c,d) northwestern Pacific.
525           Circles mark locations of vertical profiles shown in Figs. 6 and 8.
526




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527




528
529   Figure 6. Sample vertical profiles of temperature and salinity with (a) negative and (b) positive
530            dT . Profiles are taken in the northwestern Atlantic in (a) boreal summer and (b) boreal
531           winter at locations shown in Figs. 5a and 5b, respectively. ‘LT’ indicates the local sun
532           time.
533


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534
535
536   Figure 7. Mixed layer model response to sample winds and net surface flux in the Gulf Stream
537           area in (a,b) summer (c,d) winter as a function of salinity stratification. (b,d) Difference
538           of temperature between two simulations with the same surface forcing and with the same
539           vertically uniform temperature initial conditions but with different salinity initial
540           conditions. In the first experiment initial salinity is vertically uniform while in the second
541           initial salinity has a uniform vertical gradient, S / z =0.1 psu m-1.
542

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543




544
545   Figure 8 Sample vertical profiles of temperature and salinity with (a) negative and (b) positive
546            dT . Profiles are taken in the northwestern Pacific in boreal summer at locations shown
547           in Fig. 5c. ‘LT’ indicates the local sun time.




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