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Role of Horizontal Density Advection in Seasonal Deepening of the Mixed Layer in the Subtropical Southeast Pacific


doi: 10.1007/s00376-015-5111-x

  • The mechanisms behind the seasonal deepening of the mixed layer (ML) in the subtropical Southeast Pacific were investigated using the monthly Argo data from 2004 to 2012. The region with a deep ML (more than 175 m) was found in the region of (22°-30°S, 105°-90°W), reaching its maximum depth (∼200 m) near (27°-28°S, 100°W) in September. The relative importance of horizontal density advection in determining the maximum ML location is discussed qualitatively. Downward Ekman pumping is key to determining the eastern boundary of the deep ML region. In addition, zonal density advection by the subtropical countercurrent (STCC) in the subtropical Southwest Pacific determines its western boundary, by carrying lighter water to strengthen the stratification and form a "shallow tongue" of ML depth to block the westward extension of the deep ML in the STCC region. The temperature advection by the STCC is the main source for large heat loss from the subtropical Southwest Pacific. Finally, the combined effect of net surface heat flux and meridional density advection by the subtropical gyre determines the northern and southern boundaries of the deep ML region: the ocean heat loss at the surface gradually increases from 22°S to 35°S, while the meridional density advection by the subtropical gyre strengthens the stratification south of the maximum ML depth and weakens the stratification to the north. The freshwater flux contribution to deepening the ML during austral winter is limited. The results are useful for understanding the role of ocean dynamics in the ML formation in the subtropical Southeast Pacific.
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  • AVISO, 2008: SSALTO/DUACS User Handbook: (M)SLA and (M)ADT Near-Real Time and Delayed Time Products. Collecte Localisation Satellites,Agne, France, 39 pp.
    Colbo K., R. Weller, 2007: The variability and heat budget of the upper ocean under the Chile-Peru stratus. J. Mar. Res., 65, 607- 637.10.1357/0022240077836495109a29de2de7e9cab6dacd9484887092c9http%3A%2F%2Fwww.ingentaconnect.com%2Fcontent%2Fjmr%2Fjmr%2F2007%2F00000065%2F00000005%2Fart00002http://www.ingentaconnect.com/content/jmr/jmr/2007/00000065/00000005/art00002The persistent stratus clouds found west of Chile and Peru are important for the coupling of the ocean and atmosphere in the eastern Pacific and thus in the climate of the region. The relatively cool sea-surface temperatures found west of Peru and northern Chile are believed to play a role in maintaining the stratus clouds over the region. In October 2000 a buoy was deployed at 20S, 85W, a site near the center of the stratus region, in order to examine the variability of sea-surface temperature and the temporal evolution of the vertical structure of the upper ocean. The buoy was wellinstrumented and obtained accurate time series of the surface forcing as well as time series in the upper ocean of temperature, salinity, and velocity. The variability and the extent to which local forcing explains the temporal evolution of upper ocean structure and heat content was examined. The sources of heating (primarily surface fluxes with weaker contributions from Ekman convergence and transport) are found to be balanced by cooling from the gyre-scale circulation, an eddy flux divergence and vertical diffusion. The deduced eddy flux divergence term is bounded away from zero and represents an order one source of cooling (and freshening). We postulate that the eddy flux divergence represents the effect of the cold coherent eddies formed near the coast, which propagate westward and slowly decay. Direct advection of coastal upwelled water by Ekman transport is negligible. Thus the upwelled water does influence the offshore structure, but through the fluctuating mesoscale flow not the mean transport.
    de Boyer MontèGut, C., G. Madec, A. S. Fischer, A. Lazar, D. Iudicone, 2004: Mixed layer depth over the global ocean: An examination of profile data and a profile-based climatology. J. Geophys. Res,109, C12003, 481- 497.10.1029/2004JC002378549958876721e0d1668cf9a76d379a68http%3A%2F%2Fonlinelibrary.wiley.com%2Fdoi%2F10.1029%2F2004JC002378%2Ffullhttp://onlinelibrary.wiley.com/doi/10.1029/2004JC002378/fullA new 2 resolution global climatology of the mixed layer depth (MLD) based on individual profiles is constructed. Previous global climatologies have been based on temperature or density-gridded climatologies. The criterion selected is a threshold value of temperature or density from a near-surface value at 10 m depth (T = 0.2C or = 0.03 kg m613). A validation of the temperature criterion on moored time series data shows that the method is successful at following the base of the mixed layer. In particular, the first spring restratification is better captured than with a more commonly used larger criteria. In addition, we show that for a given 0.2C criterion, the MLD estimated from averaged profiles results in a shallow bias of 25% compared to the MLD estimated from individual profiles. A new global seasonal estimation of barrier layer thickness is also provided. An interesting result is the prevalence in mid- and high-latitude winter hemispheres of vertically density-compensated layers, creating an isopycnal but not mixed layer. Consequently, we propose an optimal estimate of MLD based on both temperature and density data. An independent validation of the maximum annual MLD with oxygen data shows that this oxygen estimate may be biased in regions of Ekman pumping or strong biological activity. Significant differences are shown compared to previous climatologies. The timing of the seasonal cycle of the mixed layer is shifted earlier in the year, and the maximum MLD captures finer structures and is shallower. These results are discussed in light of the different approaches and the choice of criterion.
    Deser C., M. A. Alexand er, and M. S. Timlin, 1996: Upper ocean thermal variations in the North pacific during 1970-1991. J. Climate, 9, 1840- 1855.10.1175/1520-0442(1996)009<1840:UOTVIT>2.0.CO;286610051-77c7-4c72-b2a2-e0cf2e1e5193c149d3a1d806289914cb77607c8d2b4bhttp%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F1996JCli....9.1840Drefpaperuri:(5f9d5f0523469db4bf6890e0c476f166)http://adsabs.harvard.edu/abs/1996JCli....9.1840DAbstract A newly available, extensive compilation of upper-ocean temperature profiles was used to study the vertical structure of thermal anomalies between the surface and 400-m depth in the North Pacific during 1970-1991. A prominent decade-long perturbation in climate occurred during this time period: surface waters cooled by 651C in the central and western North Pacific and warmed by about the same amount along the west coast of North America from late 1976 to 1988. Comparison with data from COADS suggests that the relatively sparse sampling of the subsurface data is adequate for describing the climate anomaly. The vertical structure of seasonal thermal anomalies in the central North Pacific shows a series of cold pulses beginning in the fall of 1976 and continuing until late 1988 that appear to originate at the surface and descend with time into the main thermocline to at least 400-m depth. Individual cold events descend rapidly (65100 m yr 611 ), superimposed upon a slower cooling (6515 m yr 611 ). The interdecadal climate change, while evident at the surface, is most prominent below 65150 m where interannual variations are small. Unlike the central North Pacific, the temperature changes along the west coast of North America appear to be confined to approximately the upper 200-250 m. The structure of the interdecadal thermal variations in the eastern and central North Pacific appears to be consistent with the dynamics of the ventilated thermocline. In the western North Pacific, strong cooling is observed along the axis of the Kuroshio Current Extension below 65200 m depth during the 1980s. Changes in mixed layer depth accompany the SST variations, but their spatial distribution is not identical to the pattern of SST change. In particular, the decade-long cool period in the central North Pacific was accompanied by a 6520 m deepening of the mixed layer in winter, but no significant changes in mixed layer depth were found along the west coast of North America. It is suggested that other factors such as stratification beneath the mixed layer and synoptic wind forcing may play a role in determining the distribution of mixed layer depth anomalies.
    Huang R. X., 2010. Oceanic Circulation: Wind-driven and Thermohaline Processes. Cambridge University Press, Cambridge, 360- 369.
    Kara A. B., P. A. Rochford, and H. E. Hurlburt, 2003: Mixed layer depth variability over the global ocean. J. Geophys. Res.: Oceans (1978-2012), 108( C3), 209.10.1029/2000JC0007363b06fed7-a1db-4491-9560-bd359f234b67abd3c64f533899a42ab5fe0dccca1c14http%3A%2F%2Fonlinelibrary.wiley.com%2Fdoi%2F10.1029%2F2000JC000736%2Ffullrefpaperuri:(9554aced13de896e8244726590eaf470)http://onlinelibrary.wiley.com/doi/10.1029/2000JC000736/fullAbstract Top of page Abstract 1.Introduction 2.Data and Limitations 3.Mixed Layer Processes and MLD Criterion 4.Overview of Global MLD Variability 5.ILD and MLD Correspondence 6.Validation of ILD Versus MLD Correspondence 7.Conclusions Acknowledgments References Supporting Information [1] The spatial and monthly variability of the climatological mixed layer depth (MLD) for the global ocean is examined using the recently developed Naval Research Laboratory (NRL) Ocean Mixed Layer Depth (NMLD) climatologies. The MLD fields are constructed using the subsurface temperature and salinity data from the World Ocean Atlas 1994 [ Levitus et al. , 1994 ; Levitus and Boyer , 1994 ]. To minimize the limitations of these global data in the MLD determination, a simple mixing scheme is introduced to form a stable water column. Using these new data sets, global MLD characteristics are produced on the basis of an optimal definition that employs a density-based criterion having a fixed temperature difference of T = 0.8 and variable salinity. Strong seasonality of MLD is found in the subtropical Pacific Ocean and at high latitudes, as well as a very deep mixed layer in the North Atlantic Ocean in winter and a very shallow mixed layer in the Antarctic in all months. Using the climatological monthly MLD and isothermal layer depth (ILD) fields from the NMLD climatologies, an annual mean T field is presented, providing criteria for determining an ILD that is approximately equivalent to the optimal MLD. This enables MLD to be determined in cases where salinity data are not available. The validity of the correspondence between ILD and MLD is demonstrated using daily averaged subsurface temperature and salinity from two moorings: a Tropical Atmosphere Ocean array mooring in the western equatorial Pacific warm pool, where salinity stratification is important, and a Woods Hole Oceanographic Institute (WHOI) mooring in the Arabian Sea, where strongly reversing seasonal monsoon winds prevail. In the western equatorial Pacific warm pool the use of ILD criterion with an annual mean T value of 0.3鎺矯 yields comparable results with the optimal MLD, while large T values yield an overestimated MLD. An analysis of ILD and MLD in the WHOI mooring show that use of an incorrect T criterion for the ILD may underestimate or overestimate the optimal MLD. Finally, use of the spatial annual mean T values constructed from the NMLD climatologies can be used to estimate the optimal MLD from only subsurface temperature data via an equivalent ILD for any location over the global ocean.
    Large W. G., S. G. Yeager, 2008: The global climatology of an interannually varying air-sea flux data set. Climate Dyn.,24, 341-364, doi: 10.1007/s00382-008-0441-3.10.1007/s00382-008-0441-36f6ad5b7-9ad1-4149-ae6f-befe4def052d730a7711ab2b95163dff278b8150175fhttp%3A%2F%2Fwww.springerlink.com%2Fcontent%2Fa1868543150g2678%2Frefpaperuri:(31d84218ab793598b06e6948a1b5fa04)http://onlinelibrary.wiley.com/resolve/reference/ADS?id=2009ClDy...33..341LThe air鈥搒ea fluxes of momentum, heat, freshwater and their components have been computed globally from 194802at frequencies ranging from 6-hourly to monthly. All fluxes are computed over the 2302years f
    Liu C. Y., Z. M. Wang, 2014: On the response of the global subduction rate to global warming in coupled climate models. Adv. Atmos. Sci.,31(1), 211-218, doi: 10.1007/s00376-013-2323-9.10.1007/s00376-013-2323-90adb0d71-aa10-4a5d-842b-fbdf69b862a313fef6623eccd076709e69322e54fb90http%3A%2F%2Fwww.cqvip.com%2FQK%2F84334X%2F201401%2F48212579.htmlrefpaperuri:(f3471fcc3ba0df7b8138152288846928)http://d.wanfangdata.com.cn/Periodical_dqkxjz-e201401020.aspxThe response of the global subduction rate to global warming was assessed based on a set of Intergovernmental Panel on Climate Change (IPCC) Fourth Assessment Report (AR4) models. It was found that the subduction rate of the global ocean could be significantly reduced under a warming climate, as compared to a simulation of the present-day climate. The reduction in the subduction volume was quantitatively estimated at about 40 Sv and was found to be primarily induced by the decreasing of the lateral induction term due to a shallower winter mixed layer depth. The shrinking of the winter mixed layer would result from intensified stratification caused by increased heat input into the ocean under a warming climate. A reduction in subduction associated with the vertical pumping term was estimated at about 5 Sv. Further, in the Southern Ocean, a significant reduction in subduction was estimated at around 24 Sv, indicating a substantial contribution to the weakening of global subduction.
    Liu H. L., W. Y. Lin, and M. H. Zhang, 2010: Heat budget of the upper ocean in the south-central Equatorial Pacific. J. Climate, 23( 7), 1779- 1792. doi: 10.1175/2009JCLI3135.110.1175/2009JCLI3135.1de052b63b0e06393bc5a57ae8c4f6a80http%3A%2F%2Fwww.cabdirect.org%2Fabstracts%2F20103168346.htmlhttp://www.cabdirect.org/abstracts/20103168346.htmlThe double intertropical convergence zone (ITCZ) over the tropical Pacific, with a spurious band of maximum annual sea surface temperature (SST) south of the equator between 500S and 1000S, is a chronic bias in coupled ocean09 tmosphere models. This study focuses on a region of the double ITCZ in the central Pacific from 500S to 1000S and 17000E to 15000W, where coupled models display the largest biases in precipitation, by deriving a best estimate of the mixed layer heat budget for the region. Seven global datasets of objectively analyzed surface energy fluxes and four ocean assimilation products are first compared and then evaluated against field measurements in adjacent regions. It was shown that the global datasets differ greatly in their net downward surface energy flux in this region, but they fall broadly into two categories: one with net downward heat flux of about 30 W m0903-2 and the other around 10 W m0903-2. Measurements from the adjacent Manus and Nauru sites of the Atmospheric Radiation Measurement Program (ARM), the Tropical Atmosphere Ocean (TAO) buoys, and the Tropical Ocean and Global Atmosphere Coupled Ocean09 tmosphere Response Experiment (TOGA COARE) are then used to show that the smaller value is more realistic. An energy balance of the mixed layer is finally presented for the region as primarily between warming from surface heat flux of 7 W m0903-2 and horizontal advective cooling in the zonal direction of about 5 W m0903-2, with secondary contributions from meridional and vertical advections, heat storage, and subgrid-scale mixing. The 7 W m0903-2 net surface heat flux consists of warming of 210 W m0903-2 from solar radiation and cooling of 53, 141, and 8 W m0903-2, respectively, from longwave radiation, latent heat flux, and sensible heat flux. These values provide an observational basis to further study the initial development of excessive precipitation in coupled climate models in the central Pacific.
    Liu L. L., R. X. Huang, 2012: The global subduction/obduction rates: Their interannual and decadal variability. J. Climate, 25( 4), 1096- 1115.10.1175/2011JCLI4228.1408135e09061689c43b4a941d3fa4286http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2012JCli...25.1096Lhttp://adsabs.harvard.edu/abs/2012JCli...25.1096LVentilation, including subduction and obduction, for the global oceans was examined using Simple Ocean Data Assimilation (SODA) outputs. The global subduction rate averaged over the period from 1959 to 2006 is estimated at 505.8 Sv (1 Sv 09-03 106 m3 s0903-1), while the corresponding global obduction rate is estimated at 482.1 Sv. The annual subduction/obduction rates vary greatly on the interannual and decadal time scales. The global subduction rate is estimated to have increased 7.6%% over the past 50 years, while the obduction rate is estimated to have increased 9.8%%. Such trends may be insignificant because errors associated with the data generated by ocean data assimilation could be as large as 10%%. However, a major physical mechanism that induced these trends is primarily linked to changes in the Southern Ocean. While the Southern Ocean plays a key role in global subduction and obduction rates and their variability, both the Southern Ocean and equatorial regions are critically important sites of water mass formation/erosion.
    Luo Y. Y., Q. Y. Liu, and L. M. Rothstein, 2011: Increase of South Pacific eastern subtropical mode water under global warming. Geophys. Res. Lett., 38, L01601.10.1029/2010GL0458789f9d7a49b10f4b0af30817eb2880b400http%3A%2F%2Fonlinelibrary.wiley.com%2Fdoi%2F10.1029%2F2010GL045878%2Ffullhttp://onlinelibrary.wiley.com/doi/10.1029/2010GL045878/full[1] The response of South Pacific Eastern Subtropical Mode Water (SPESTMW) to global warming is investigated by comparing solutions from a set of Intergovernmental Panel on Climate Change (IPCC) Fourth Assessment Report (AR4) coupled models between a present-day climate and a future, warmer climate. Under the warmer climate scenario, the SPESTMW extends southwestward and is significantly increased in volume. This is because all the local surface forcing mechanisms (i.e., wind stress, heat and freshwater fluxes) in the eastern subtropical South Pacific tends to de-stratify the upper ocean and thus deepen the mixed layer. Further, a suite of process-oriented experiments with an ocean general circulation model suggest that it is the intensified southeast trade winds under the warmer climate that promotes more heat flux from the ocean into the atmosphere that then results in a deepening of the mixed layer in the eastern subtropics of the South Pacific.
    Nishikawa S., A. Kubokawa, 2012: Mixed layer depth front and subduction of low potential vorticity water under seasonal forcings in an idealized OGCM. Journal of Oceanography, 68( 1), 53- 62.10.1007/s10872-011-0086-4c0064573-4ed8-446a-9cdf-586742dd2c7850778201268153-62Abstract<br/>The mixed layer depth (MLD) front and subduction under seasonal variability are investigated using an idealized ocean general circulation model (OGCM) with simple seasonal forcings. A sharp MLD front develops and subduction occurs at the front from late winter to early spring. The position of the MLD front agrees with the curve where <span class="a-plus-plus inline-equation id-i-eq1"><span class="a-plus-plus equation-source format-t-e-x">${\rm D}T_{\rm s}/{\rm D}t = \partial T_{\rm s} /\partial t + {\user2{u}}_{\rm g} \cdot \nabla T_{\rm s} = 0$</span></span> is satisfied (<em class="a-plus-plus">t</em> is time, <span class="a-plus-plus inline-equation id-i-eq2"><span class="a-plus-plus equation-source format-t-e-x">${\user2{u}}_{\rm g}$</span></span> is the upper-ocean geostrophic velocity, <span class="a-plus-plus inline-equation id-i-eq57"><span class="a-plus-plus equation-source format-t-e-x">$T_{\rm s}$</span></span> is the sea surface temperature (SST), and <span class="a-plus-plus inline-equation id-i-eq58"><span class="a-plus-plus equation-source format-t-e-x">$\nabla$</span></span> is the horizontal gradient operator), indicating that thick mixed-layer water is subducted there parallel to the SST contour. This is a generalization of the past result that the MLD front coincides with the curve <span class="a-plus-plus inline-equation id-i-eq3"><span class="a-plus-plus equation-source format-t-e-x">${\user2{u}}_{\rm g} \cdot \nabla T_{\rm s} = 0$</span></span> when the forcing is steady. Irreversible subduction at the MLD front is limited to about 1 month, where the beginning of the irreversible subduction period agrees with the first coincidence of the MLD front and <span class="a-plus-plus inline-equation id-i-eq59"><span class="a-plus-plus equation-source format-t-e-x">${\rm D}T_{\rm s}/{\rm D}t =0$</span></span> in late winter, and the end of the period roughly corresponds to the disappearance of the MLD front in early spring. Subduction volume at the MLD front during this period is similar to that during 1 year in the steady-forcing model. Since the cooling of the deep mixed-layer water occurs only in winter and SST can not fully catch up with the seasonally varying reference temperature of restoring, the cooling rate of SST is reduced and the zonal gradient of the SST in the northwestern subtropical gyre is a little altered in the seasonal-forcing case. These effects result in slightly lower densities of subducted water and the eastward shift of the MLD front.<br/>
    Pan A. J., Q. Y. Liu, and Z. Y. Liu, 2008: Formation mechanism of the "Stability Gap" and the North Pacific central mode water. Chinese Journal of Geophysics, 51( 1), 77- 87. (in Chinese)10.1002/cjg2.11954204b447-ca8f-4a85-9092-195362973783982fa326560e41d3ceed9842f7a68a8fhttp%3A%2F%2Fonlinelibrary.wiley.com%2Fdoi%2F10.1002%2Fcjg2.1195%2Fcitedbyhttp://en.cnki.com.cn/Article_en/CJFDTotal-DQWX200801012.htmLocal feature of the formation region (165E-160W,38N-42N) of the North Pacific Central Mode Water (NPCMW) is first put forward from data analysis, and for which, the external atmospheric forcing (solar shortwave radiation, net heat flux and wind stress curl) could not give acceptable explanation. Further analysis on the seasonal variability of the upper ocean stratification shows that a special weak zone of the ocean stratification in the upper ocean (75 m)-the "stability gap" is detected in (165E-160W,38N-42N) in autumn (September-October). As "Precondition Mechanism", the "stability gap" provides a reliable answer for the "local feature" of the formation of the NPCMW. Based on a heat balance equation of the upper ocean mixed layer, diagnostic analysis suggests that the formation of the "stability gap" is the cooperative product of the surface heat flux forcing, vertical entrainment, Ekman advection and geostrophic advction. Among which, the latitudinal differences of the surface heat flux forcing, the cold Ekman advection and the warm geostrophic advection play the crucial roles on determining the critical eastern and western bound of the "stability gap".
    Qiu B., R. X. Huang, 1995: Ventilation of the North Atlantic and North Pacific: Subduction versus obduction. J. Phys. Oceanogr., 25, 2374- 2390.2fa79fd7ce9aac653661a58184a0be93http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F1995jpo....25.2374q/s?wd=paperuri%3A%28b80d133fabb514fdf3e3383bda174649%29&filter=sc_long_sign&tn=SE_xueshusource_2kduw22v&sc_vurl=http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F1995jpo....25.2374q&ie=utf-8
    Sato K., T. Suga, 2009: Structure and modification of the South Pacific eastern subtropical mode water. J. Phys. Oceanogr., 39, 1700- 1714.10.1175/2008JPO3940.1f59a0a7d3666709922da986ba122942fhttp%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2009JPO....39.1700Shttp://adsabs.harvard.edu/abs/2009JPO....39.1700SUsing all available temperature and salinity profiles obtained by Argo floats from July 2004 to June 2007, this study investigated the structure and modification of the South Pacific Eastern Subtropical Mode Water (SPESTMW). Based on the observed characteristics of the vertical minima of potential vorticity over the subtropical South Pacific, SPESTMW is defined as water with potential vorticity magnitude less than 2.5 01- 10-10 m-1 s-1 and thickness exceeding 40 m. It is found between 350009/500S and 1600009/7000W and has a temperature of 130009/2600C, salinity greater than 34.0, and density of 24.509/25.8 kg m-3 at its core. This study confirmed that vertical changes in temperature and salinity tend to compensate for each other in terms of density changes, resulting in favorable salt fingering conditions, as previously reported. By analyzing many profiles of Argo data in spring immediately after the SPESTMW formation period, its temperature and salinity are vertically uniform in the formation region, but large vertical gradients of temperature and salinity are found downstream from that region, even in the SPESTMW core. Consequently, the low potential vorticity signature of SPESTMW spread much wider than its signature as a thermostad. The Argo data also captured the seasonal changes of the vertical gradients of temperature and salinity at the SPESTMW core; these gradients increased as the seasons progressed, even in the formation region. Therefore, SPESTMW is truly vertically uniform water (i.e., thermostad, halostad, and pycnostad simultaneously) only immediately after the formation period. Afterward, it is only pycnostad. This seasonal evolution is related to temperature and salinity diffusion due to salt fingering in a manner similar to the rapid modification of interannual anomalies as shown by previous research. The temperature and salinity near the SPESTMW core and lower region decreased soon after its formation.
    Williams R. G., 1991: The role of the mixed layer in setting the potential vorticity of the main thermocline. J. Phys. Oceanogr., 21, 1803- 1814.10.1175/1520-0485(1991)021<1803:TROTML>2.0.CO;2f3264ce549c80ecf19d7da305ce3a768http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F1991JPO....21.1803Whttp://adsabs.harvard.edu/abs/1991JPO....21.1803WAbstract A steady ventilation model is used to assess the effect of the mixed layer on the structure of the main thermocline; the potential vorticity is found in a subtropical gyre after imposing the thickness and density of the mixed layer, the Ekman pumping, and the hydrography on the eastern boundary. The modeled potential vorticity becomes comparable in value to observations in the North Atlantic if the mixed layer deepens poleward as is observed in winter. The isopycnal gradients in potential vorticity are reduced on the denser ventilated surfaces if the mixed-layer outcrops deviate from latitude circles and, more realistically, sweep southward along the eastern boundary; the age of the subducted fluid is also in reasonable agreement with observations of the tritium-helium age by Jenkins. This study suggests that ventilation may form much of the extensive region of nearly uniform potential vorticity observed on the = 26.75 surface in the North Atlantic, with lateral mixing by eddies being required only in the unventilated pool on the western side of the gyre.
    Wong A. P. S., G. C. Johnson, 2003: South Pacific eastern subtropical mode water. J. Phys. Oceanogr., 33( 7), 1493- 1509.ab3a6bf0d644d4cf082c87066a01d7a8http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2003JPO....33.1493W/s?wd=paperuri%3A%289ea36b14be1c1e9376426ad380c08244%29&filter=sc_long_sign&tn=SE_xueshusource_2kduw22v&sc_vurl=http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2003JPO....33.1493W&ie=utf-8
    Xie P. P., P.A. Arkin, 1997: Global precipitation: A 17-year monthly analysis based on gauge observations, satellite estimates, and numerical model outputs. Bull. Amer. Meteor. Soc., 78, 2539- 2558.10.1175/1520-0477(1997)078<2539:GPAYMA>2.0.CO;2b11e9f6a-e0a7-42f9-a97b-a26069da3c8d3039680de89ffc852c3d9b6b72a9b3dbhttp%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F1997BAMS...78.2539Xrefpaperuri:(f637afb56e50553202efc8c31489db4c)http://adsabs.harvard.edu/abs/1997BAMS...78.2539XAbstract Gridded fields (analyses) of global monthly precipitation have been constructed on a 2.5 latitude ongitude grid for the 17-yr period from 1979 to 1995 by merging several kinds of information sources with different characteristics, including gauge observations, estimates inferred from a variety of satellite observations, and the NCEPCAR reanalysis. This new dataset, which the authors have named the CPC Merged Analysis of Precipitation (CMAP), contains precipitation distributions with full global coverage and improved quality compared to the individual data sources. Examinations showed no discontinuity during the 17-yr period, despite the different data sources used for the different subperiods. Comparisons of the CMAP with the merged analysis of Huffman et al. revealed remarkable agreements over the global land areas and over tropical and subtropical oceanic areas, with differences observed over extratropical oceanic areas. The 17-yr CMAP dataset is used to investigate the annual and interannual variability in large-scale precipitation. The mean distribution and the annual cycle in the 17-yr dataset exhibit reasonable agreement with existing long-term means except over the eastern tropical Pacific. The interannual variability associated with the El Ni09o鈥揝outhern Oscillation phenomenon resembles that found in previous studies, but with substantial additional details, particularly over the oceans. With complete global coverage, extended period and improved quality, the 17-yr dataset of the CMAP provides very useful information for climate analysis, numerical model validation, hydrological research, and many other applications. Further work is under way to improve the quality, extend the temporal coverage, and to refine the resolution of the merged analysis.
    Xie S. P., L. X. Xu, Q. Y. Liu, and F. Kobashi, 2011: Dynamical role of mode water ventilation in decadal variability in the central subtropical gyre of the North Pacific. J. Climate, 24, 1212- 1225.10.1175/2010JCLI3896.1132699c5078a1bfc35667ede9e7a7029http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2011JCli...24.1212Xhttp://adsabs.harvard.edu/abs/2011JCli...24.1212XAbstract Decadal variability in the interior subtropical North Pacific is examined in the Geophysical Fluid Dynamics Laboratory coupled model (CM2.1). Superimposed on a broad, classical subtropical gyre is a narrow jet called the subtropical countercurrent (STCC) that flows northeastward against the northeast trade winds. Consistent with observations, the STCC is anchored by mode water characterized by its low potential vorticity (PV). Mode water forms in the deep winter mixed layer of the Kuroshio yashio Extension (KOE) east of Japan and flows southward riding on the subtropical gyre and preserving its low-PV characteristic. As a thick layer of uniform properties, the mode water forces the upper pycnocline to shoal, and the associated eastward shear results in the surface-intensified STCC. On decadal time scales in the central subtropical gyre (15-35N, 170E-130W), the dominant mode of sea surface height variability is characterized by the strengthening and weakening of the STCC because of variations in mode water ventilation. The changes in mode water can be further traced upstream to variability in the mixed layer depth and subduction rate in the KOE region. Both the mean and anomalies of STCC induce significant sea surface temperature anomalies via thermal advection. Clear atmospheric response is seen in wind curls, with patterns suggestive of positive coupled feedback. In oceanic and coupled models, northeast-slanted bands often appear in anomalies of temperature and circulation at and beneath the surface. The results of this study show that such slanted bands are characteristic of changes in mode water ventilation. Indeed, this natural mode of STCC variability is excited by global warming, resulting in banded structures in sea surface warming.
    Xu L. X., S. P. Xie, J. L. McClean, Q. Y. Liu, and H. Sasaki, 2014: Mesoscale eddy effects on the subduction of North Pacific mode waters. J. Geophys. Res.-Oceans,119, 4867-4886, doi: 10.1002/2014JC009861.10.1002/2014JC009861c36cb47bf8cc6fc905ae1856eba81c45http%3A%2F%2Fonlinelibrary.wiley.com%2Fdoi%2F10.1002%2F2014JC009861%2Fabstracthttp://onlinelibrary.wiley.com/doi/10.1002/2014JC009861/abstractAbstract Mesoscale eddy effects on the subduction of North Pacific mode waters are investigated by comparing observations and ocean general circulation models where eddies are either parameterized or resolved. The eddy-resolving models produce results closer to observations than the noneddy-resolving model. There are large discrepancies in subduction patterns between eddy-resolving and noneddy-resolving models. In the noneddy-resolving model, subduction on a given isopycnal is limited to the cross point between the mixed layer depth (MLD) front and the outcrop line whereas in eddy-resolving models and observations, subduction takes place in a broader, zonally elongated band within the deep mixed layer region. Mesoscale eddies significantly enhance the total subduction rate, helping create remarkable peaks in the volume histogram that correspond to North Pacific subtropical mode water (STMW) and central mode water (CMW). Eddy-enhanced subduction preferentially occurs south of the winter mean outcrop. With an anticyclonic eddy to the west and a cyclonic eddy to the east, the outcrop line meanders south, and the thermocline/MLD shoals eastward. As eddies propagate westward, the MLD shoals, shielding the water of low potential vorticity from the atmosphere. The southward eddy flow then carries the subducted water mass into the thermocline. The eddy subduction processes revealed here have important implications for designing field observations and improving models.
    Yu L. S., 2007: Global variations in oceanic evaporation (1958-2005): The role of the changing wind speed. J. Climate, 20( 21), 5376- 5390.10.1175/2007JCLI1714.17c5dd972-fc32-4abb-8652-3efe3da52001528d04042991cf1eef016e690ab7ecbahttp%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2007JCli...20.5376Yrefpaperuri:(116174cc574f783c4fcd02bf6f4eddc1)http://adsabs.harvard.edu/abs/2007JCli...20.5376YAbstract Global estimates of oceanic evaporation (Evp) from 1958 to 2005 have been recently developed by the Objectively Analyzed Airea Fluxes (OAFlux) project at the Woods Hole Oceanographic Institution (WHOI). The nearly 50-yr time series shows that the decadal change of the global oceanic evaporation (Evp) is marked by a distinct transition from a downward trend to an upward trend around 1977-78. Since the transition, the global oceanic Evp has been up about 11 cm yr 611 (6510%), from a low at 103 cm yr 611 in 1977 to a peak at 114 cm yr 611 in 2003. The increase in Evp was most dramatic during the 1990s. The uncertainty of the estimates is about 卤2.74 cm yr 611 . By utilizing the newly developed datasets of Evp and related airea variables, the study investigated the cause of the decadal change in oceanic Evp. The decadal differences between the 1990s and the 1970s indicates that the increase of Evp in the 1990s occurred over a global scale and had spatially coherent structures. Larger Evp is most pronounced in two key regionsne is the paths of the global western boundary currents and their extensions, and the other is the tropical Indo-Pacific warm water pools. It is also found that Evp was enlarged primarily during the hemispheric wintertime (defined as the mean of Decemberebruary for the northern oceans and June ugust for the southern oceans). Despite the dominant upward tendency over the global basins, a slight reduction in Evp appeared in such regions as the subtropical centers of the Evp maxima as well as the eastern equatorial Pacific and Atlantic cold tongues. An empirical orthogonal function (EOF) analysis was performed for the yearly winter-mean time series of Evp and the related airea variables [i.e., wind speed ( U ) and airea humidity differences ( dq )]. The analysis suggested a dominant role of the wind forcing in the decadal change of both Evp and dq . It is hypothesized that wind impacts Evp in two ways. The first way is direct: the greater wind speed induces more evaporation by carrying water vapor away from the evaporating surface to allow the airea humidity gradients to be reestablished at a faster pace. The second way is indirect: the enhanced surface wind strengthens the wind-driven subtropical gyre, which in turn drives a greater heat transport by the western boundary currents, warms up SST along the paths of the currents and extensions, and causes more evaporation by enlarging the airea humidity gradients. The EOF analysis performed for the time series of the global annual-mean Evp fields showed that the first three EOF modes account for nearly 50% of the total variance. The mode 1 variability represents the upward trend in Evp after 1978 and is attributable to the increased U , and the mode 2 variability explains much of the downward trend in Evp before 1978 and is correlated to the global dq variability. The EOF mode 3 of Evp captures the interannual variability of Evp on time scales of the El Ni09oouthern Oscillation, with the center of action over the eastern equatorial Pacific.
    Yu L. S., R. A. Weller, 2007: Objectively analyzed air-sea heat fluxes for the global ice-free oceans (1981-2005). Bull. Amer. Meteor. Soc., 88( 4), 527- 539.10.1175/BAMS-88-4-527e1d266f6-da9d-4447-9ea8-46e3e0cd317e13f000d1e05099b83ea6218aa70a28f8http%3A%2F%2Fadsabs.harvard.edu%2Fabs%2F2007BAMS...88..527Yrefpaperuri:(6110649686cf0990540db0110f7812f8)http://adsabs.harvard.edu/abs/2007BAMS...88..527YAbstract A 25-yr (1981-2005) time series of daily latent and sensible heat fluxes over the global ice-free oceans has been produced by synthesizing surface meteorology obtained from satellite remote sensing and atmospheric model reanalyses outputs. The project, named Objectively Analyzed Air ea Fluxes (OAFlux), was developed from an initial study of the Atlantic Ocean that demonstrated that such data synthesis improves daily flux estimates over the basin scale. This paper introduces the 25-yr heat flux analysis and documents variability of the global ocean heat flux fields on seasonal, interannual, decadal, and longer time scales suggested by the new dataset. The study showed that, among all the climate signals investigated, the most striking is a long-term increase in latent heat flux that dominates the data record. The globally averaged latent heat flux increased by roughly 9 W m 612 between the low in 1981 and the peak in 2002, which amounted to about a 10% increase in the mean value over the 25-yr period. Positive linear trends appeared on a global scale, and were most significant over the tropical Indian and western Pacific warm pool and the boundary current regions. The increase in latent heat flux was in concert with the rise of sea surface temperature, suggesting a response of the atmosphere to oceanic forcing.
    Yu L. S., X. Z. Jin, and R. A. Weller, 2006: Role of net surface heat flux in seasonal evolutions of sea surface temperature in the tropical Atlantic Ocean. J. Climate, 19, 6153- 6169.10.1175/JCLI3970.11f7ff90c-7a5e-46c0-8cfd-c234c48225d3026bea503a13f7778f3a9cfcf13a6182http%3A%2F%2Fwww.researchgate.net%2Fpublication%2F49245594_Role_of_net_surface_heat_flux_in_the_seasonal_evolution_of_sea_surface_temperature_in_the_Atlantic_Oceanrefpaperuri:(33c69410c2aad0530942cb4216409dee)http://www.researchgate.net/publication/49245594_Role_of_net_surface_heat_flux_in_the_seasonal_evolution_of_sea_surface_temperature_in_the_Atlantic_OceanABSTRACT The present study used a new net surface heat flux (Qnet) product obtained from the Objective Analyzed Air&ndash;Sea Fluxes (OAFlux) project and the International Satellite Cloud Climatology Project (ISCCP) to examine two specific issues&mdash;one is to which degree Qnet controls seasonal variations of sea surface temperature (SST) in the tropical Atlantic Ocean (20&ndash;20, east of 60), and the other is whether the physical relation can serve as a measure to evaluate the physical representation of a heat flux product. To better address the two issues, the study included the analysis of three additional heat flux products: the Southampton Oceanographic Centre (SOC) heat flux analysis based on ship reports, and the model fluxes from the National Centers for Environmental Prediction&ndash;National Center for Atmospheric Research (NCEP&ndash;NCAR) reanalysis and the 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40). The study also uses the monthly subsurface temperature fields from the World Ocean Atlas to help analyze the seasonal changes of the mixed layer depth (hMLD). The study showed that the tropical Atlantic sector could be divided into two regimes based on the influence level of Qnet. SST variability poleward of 5 and 10 is dominated by the annual cycle of Qnet. In these regions the warming (cooling) of the sea surface is highly correlated with the increased (decreased) Qnet confined in a relatively shallow (deep) hMLD. The seasonal evolution of SST variability is well predicted by simply relating the local Qnet with a variable hMLD. On the other hand, the influence of Qnet diminishes in the deep Tropics within 5 and 10 and ocean dynamic processes play a dominant role. The dynamics-induced changes in SST are most evident along the two belts, one of which is located on the equator and the other off the equator at about 3 in the west, which tilts to about 10 near the northwestern African coast. The study also showed that if the degree of consistency between the correlation relationships of Qnet, hMLD, and SST variability serves as a measure of the quality of the Qnet product, then the Qnet from OAFlux + ISCCP and ERA-40 are most physically representative, followed by SOC. The NCEP&ndash;NCAR Qnet is least representative. It should be noted that the Qnet from OAFlux + ISCCP and ERA-40 have a quite different annual mean pattern. OAFlux + ISCCP agrees with SOC in that the tropical Atlantic sector gains heat from the atmosphere on the annual mean basis, where the ERA-40 and the NCEP&ndash;NCAR model reanalyses indicate that positive Qnet occurs only in the narrow equatorial band and in the eastern portion of the tropical basin. Nevertheless, seasonal variances of the Qnet from OAFlux + ISCCP and ERA-40 are very similar once the respective mean is removed, which explains why the two agree with each other in accounting for the seasonal variability of SST. In summary, the study suggests that an accurate estimation of surface heat flux is crucially important for understanding and predicting SST fluctuations in the tropical Atlantic Ocean. It also suggests that future emphasis on improving the surface heat flux estimation should be placed more on reducing the mean bias. Author Posting. American Meteorological Society 2006. This article is posted here by permission of American Meteorological Society for personal use, not for redistribution. The definitive version was published in Journal of Climate 19 (2006): 6153&ndash;6169, doi:10.1175/JCLI3970.1. This study is support by the NOAA CLIVAR Atlantic under Grant NA06GP0453 and NOAA Climate observations and Climate Change and Data Detection under Grant NA17RJ1223.
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Manuscript received: 30 April 2015
Manuscript revised: 20 August 2015
Manuscript accepted: 27 August 2015
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Role of Horizontal Density Advection in Seasonal Deepening of the Mixed Layer in the Subtropical Southeast Pacific

  • 1. Physical Oceanography Laboratory/Qingdao Collaborative Innovation Center of Marine Science and Technology, Key Laboratory of Ocean-Atmosphere Interaction and Climate in Universities of Shandong, Ocean University of China, Qingdao 266100

Abstract: The mechanisms behind the seasonal deepening of the mixed layer (ML) in the subtropical Southeast Pacific were investigated using the monthly Argo data from 2004 to 2012. The region with a deep ML (more than 175 m) was found in the region of (22°-30°S, 105°-90°W), reaching its maximum depth (∼200 m) near (27°-28°S, 100°W) in September. The relative importance of horizontal density advection in determining the maximum ML location is discussed qualitatively. Downward Ekman pumping is key to determining the eastern boundary of the deep ML region. In addition, zonal density advection by the subtropical countercurrent (STCC) in the subtropical Southwest Pacific determines its western boundary, by carrying lighter water to strengthen the stratification and form a "shallow tongue" of ML depth to block the westward extension of the deep ML in the STCC region. The temperature advection by the STCC is the main source for large heat loss from the subtropical Southwest Pacific. Finally, the combined effect of net surface heat flux and meridional density advection by the subtropical gyre determines the northern and southern boundaries of the deep ML region: the ocean heat loss at the surface gradually increases from 22°S to 35°S, while the meridional density advection by the subtropical gyre strengthens the stratification south of the maximum ML depth and weakens the stratification to the north. The freshwater flux contribution to deepening the ML during austral winter is limited. The results are useful for understanding the role of ocean dynamics in the ML formation in the subtropical Southeast Pacific.

1. Introduction
  • The ocean mixed layer (ML) plays an important role in air-sea interactions and climate variability. The ML is characterized by its vertical homogeneity in temperature, salinity and density. The ML depth (MLD) determines the transfer of water mass, momentum, and energy between ocean and atmosphere (de Boyer MontèGut et al., 2004). The MLD is important for the subduction process from the surface layer to greater depths (Qiu and Huang, 1995; Xie et al., 2011; Liu and Huang, 2012), and the subduction process plays an important role in climate variability (Williams, 1991; Deser et al., 1996; Sato and Suga, 2009; Liu and Wang, 2014).

    Several mechanisms have been proposed to explain ML formation. Besides Ekman pumping and net surface heat flux, horizontal density advection can also induce change in stratification and MLD (de Boyer MontèGut et al., 2004). The Ekman drift in the upper fraction of the ML slides a different water mass over the lower ML, leading to vertical density convection, which, as a possible mechanism for vertical convection, can explain a strong horizontally density-compensated front south of Australia in winter. Analysis of the seasonal variability of the upper-ocean stratification shows that a specific region with weak stratification in the upper ocean ("stability gap") detected in the North Pacific central mode water formation region also provides a reliable answer for the "local feature" of the deep ML formation during winter. In addition, cold Ekman advection and warm geostrophic advection also play crucial roles in determining the eastern and western boundaries of the "stability gap" (Pan et al., 2008).

    Besides the Antarctic circumpolar region and the North Atlantic deep-water formation region, it is clear that there are other regions with local maximum MLD (>100 m) in the subtropical Southeast Pacific and South Atlantic, respectively, during austral winter, based on individual conductivity-temperature-depth (CTD) profiles (de Boyer MontèGut et al., 2004, Fig. 5). On the north side of the local maximum MLD in the Southeast Pacific, there is a strong MLD front and an obvious subduction process in austral winter, where the South Pacific eastern subtropical mode water (SPESTMW) forms (Wong and Johnson, 2003). After the SPESTMW is formed, it moves northwestward towards the equator, driven by the eastern component of the South Pacific subtropical gyre, and eventually joins the south equatorial current (Nishikawa and Kubokawa, 2012). As the climate warms, the SPESTMW, in sharp contrast with the response to the North Pacific mode water, tends to extend southwestward and is significantly increased in volume, which mainly depends on an intensification of the southeasterly trade wind (Luo et al., 2011). A comparison of the MLD spatial patterns from a series of numerical ocean model experiments suggested that it is the intensified southeasterly trade wind, via generating a stronger buoyancy flux from the ocean to the atmosphere, that results in a deeper ML in the subtropical Southeast Pacific in a warmer climate (Luo et al., 2011).

    It is well known that the MLD is mainly determined by vertical convection and turbulent mixing of the water mass due to wind stress and heat exchange at the air-sea surface (Kara et al., 2003). By using high-quality CTD sections and Argo profile data collected between 1991 and 1996, (Wong and Johnson, 2003) indicated that the destabilizing of the salinity gradient in the SPESTMW contributes to its formation, with its major subduction region east of 130°W and north of 30°S. The Argo profile data also captured the seasonal changes of the vertical gradients of temperature and salinity at the SPESTMW core with density of 24.5-25.8 kg m-3 (Sato and Suga, 2009). Although based on few CTD profiles, the seasonal variation of the MLD in the subtropical Southeast Pacific has been shown, but CTD data alone were insufficient and the formation mechanism with respect to a deep ML in subtropical Southeast Pacific has not been mentioned (de Boyer MontèGut et al., 2004). On the other hand, observations of net surface heat flux in the Southeast Pacific have been used for examining the heat budget of the upper 250 m in the ocean (Colbo and Weller, 2007), indicating that the equatorward heat transport compensates for nearly half of the heat balance, while horizontal eddy heat flux divergence accounts for the rest, with Ekman transport and pumping being negligible. These studies did not directly estimate horizontal density advection, because the authors only focused on the heat budget.

    In order to understand the seasonal deepening process of the ML in the subtropical Southeast Pacific, we set out to answer the following questions in the present study: When and where does the local maximum of the MLD in the subtropical Southeast Pacific appear? And what determines the location of the deep ML? As is well known, density is usually greater in the south than it is in the north, and the surface wind is southeasterly, in the subtropical Southeast Pacific. The southwestward Ekman flow and the eastward subtropical countercurrent (STCC) advecting low-density water are not conducive to ML deepening during austral winter. So, we hypothesized that the northwestward subtropical gyre current plays an important role in the seasonal deepening of the ML, because it should advect heavier water northwards from the south and weaken the upper-ocean stratification there.

    In the remainder of the paper, we introduce the data in section 2. We show the location of the maximum MLD and seasonal evolution of the MLD, and discuss the relative importance of each factor in determining the maximum ML location, in section 3. Section 4 summarizes the study's key findings.

1. Introduction
  • The ocean mixed layer (ML) plays an important role in air-sea interactions and climate variability. The ML is characterized by its vertical homogeneity in temperature, salinity and density. The ML depth (MLD) determines the transfer of water mass, momentum, and energy between ocean and atmosphere (de Boyer MontèGut et al., 2004). The MLD is important for the subduction process from the surface layer to greater depths (Qiu and Huang, 1995; Xie et al., 2011; Liu and Huang, 2012), and the subduction process plays an important role in climate variability (Williams, 1991; Deser et al., 1996; Sato and Suga, 2009; Liu and Wang, 2014).

    Several mechanisms have been proposed to explain ML formation. Besides Ekman pumping and net surface heat flux, horizontal density advection can also induce change in stratification and MLD (de Boyer MontèGut et al., 2004). The Ekman drift in the upper fraction of the ML slides a different water mass over the lower ML, leading to vertical density convection, which, as a possible mechanism for vertical convection, can explain a strong horizontally density-compensated front south of Australia in winter. Analysis of the seasonal variability of the upper-ocean stratification shows that a specific region with weak stratification in the upper ocean ("stability gap") detected in the North Pacific central mode water formation region also provides a reliable answer for the "local feature" of the deep ML formation during winter. In addition, cold Ekman advection and warm geostrophic advection also play crucial roles in determining the eastern and western boundaries of the "stability gap" (Pan et al., 2008).

    Besides the Antarctic circumpolar region and the North Atlantic deep-water formation region, it is clear that there are other regions with local maximum MLD (>100 m) in the subtropical Southeast Pacific and South Atlantic, respectively, during austral winter, based on individual conductivity-temperature-depth (CTD) profiles (de Boyer MontèGut et al., 2004, Fig. 5). On the north side of the local maximum MLD in the Southeast Pacific, there is a strong MLD front and an obvious subduction process in austral winter, where the South Pacific eastern subtropical mode water (SPESTMW) forms (Wong and Johnson, 2003). After the SPESTMW is formed, it moves northwestward towards the equator, driven by the eastern component of the South Pacific subtropical gyre, and eventually joins the south equatorial current (Nishikawa and Kubokawa, 2012). As the climate warms, the SPESTMW, in sharp contrast with the response to the North Pacific mode water, tends to extend southwestward and is significantly increased in volume, which mainly depends on an intensification of the southeasterly trade wind (Luo et al., 2011). A comparison of the MLD spatial patterns from a series of numerical ocean model experiments suggested that it is the intensified southeasterly trade wind, via generating a stronger buoyancy flux from the ocean to the atmosphere, that results in a deeper ML in the subtropical Southeast Pacific in a warmer climate (Luo et al., 2011).

    It is well known that the MLD is mainly determined by vertical convection and turbulent mixing of the water mass due to wind stress and heat exchange at the air-sea surface (Kara et al., 2003). By using high-quality CTD sections and Argo profile data collected between 1991 and 1996, (Wong and Johnson, 2003) indicated that the destabilizing of the salinity gradient in the SPESTMW contributes to its formation, with its major subduction region east of 130°W and north of 30°S. The Argo profile data also captured the seasonal changes of the vertical gradients of temperature and salinity at the SPESTMW core with density of 24.5-25.8 kg m-3 (Sato and Suga, 2009). Although based on few CTD profiles, the seasonal variation of the MLD in the subtropical Southeast Pacific has been shown, but CTD data alone were insufficient and the formation mechanism with respect to a deep ML in subtropical Southeast Pacific has not been mentioned (de Boyer MontèGut et al., 2004). On the other hand, observations of net surface heat flux in the Southeast Pacific have been used for examining the heat budget of the upper 250 m in the ocean (Colbo and Weller, 2007), indicating that the equatorward heat transport compensates for nearly half of the heat balance, while horizontal eddy heat flux divergence accounts for the rest, with Ekman transport and pumping being negligible. These studies did not directly estimate horizontal density advection, because the authors only focused on the heat budget.

    In order to understand the seasonal deepening process of the ML in the subtropical Southeast Pacific, we set out to answer the following questions in the present study: When and where does the local maximum of the MLD in the subtropical Southeast Pacific appear? And what determines the location of the deep ML? As is well known, density is usually greater in the south than it is in the north, and the surface wind is southeasterly, in the subtropical Southeast Pacific. The southwestward Ekman flow and the eastward subtropical countercurrent (STCC) advecting low-density water are not conducive to ML deepening during austral winter. So, we hypothesized that the northwestward subtropical gyre current plays an important role in the seasonal deepening of the ML, because it should advect heavier water northwards from the south and weaken the upper-ocean stratification there.

    In the remainder of the paper, we introduce the data in section 2. We show the location of the maximum MLD and seasonal evolution of the MLD, and discuss the relative importance of each factor in determining the maximum ML location, in section 3. Section 4 summarizes the study's key findings.

2. Data
  • When studying the ML and mode water subduction in the North Pacific, the use of reanalysis data produces different results to Argo data because eddy-resolving model results are closer than non-eddy resolving model results to Argo observations (Xu et al., 2014). The temperature and salinity data used in this study were from Argo profiling floats in the Pacific Ocean from January 2004 to December 2012. Each float descended to a preprogrammed parking depth (typically 1000 m), drifted freely at that depth, and then ascended to the surface at a predetermined interval (usually 10 days) after descending to the maximum pressure (2000 m). These data were collected and quality-controlled before being made freely available by the International Argo Program and the national programs that contributed to it (http://www.argo.ucsd.edu, http://argo.jcommops.org). Only profiles with a quality flag of "1" and "2", indicating "good data" and "probably good data" respectively, were used in this study, resulting in a total of 20 571 profiles in the Pacific region (40°S-20°N, 120°E-60°W) and 1425 profiles in the subtropical Southeast Pacific (20°-40°S, 90°-120°W). Further quality control was conducted by the China Argo Real-time Data Center (http://www.argo.org.cn), including interpolating the profile data into vertical standard depths (of 48) by the Akima interpolation method, and then averaging in each 1°× 1° bin. Good coverage of Argo floats ensured the viability of our study in this region. The potential temperature and density (called temperature and density henceforth) referred to the surface are calculated from the interpolated temperature and salinity data.

    The monthly net surface heat flux data from 2004 to 2009, at 1°× 1° resolution, were acquired from the Objectively Analyzed Air-Sea Fluxes (OAFlux) project at the Woods Hole Oceanographic Institution (Yu et al., 2006). The objective analysis method used in OAFlux combines optimal satellite measurements and model reanalysis data (Yu and Weller, 2007; Yu, 2007). According to (Large and Yeager, 2008) and (Liu et al., 2010), the net surface heat flux of OAFlux is overestimated by about 5-10 W m-2 in the tropical region. Therefore, we first removed the extra heat flux from the shortwave radiation (keeping only 94.5% of its shortwave radiation) before we analyzed the net heat flux (Liu et al., 2010).

    Figure 1.  Climatology of monthly-mean MLD (showing contours of 50, 75, 100, 125, 150 and 175 m) and surface wind (arrows; m s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September.

    The OAFlux monthly 1°× 1° evaporation data (in cm yr-1) from January 2004 to December 2009 were also used. Meanwhile, the monthly precipitation data (in mm d-1), on a 2.5°× 2.5° grid, were from the CMAP dataset, which merges gauge data with five kinds of satellite estimates, from January 2004 to December 2009 (Xie and Arkin, 1997).

    The Quick Scatterometer (QuikSCAT) winds were used to calculate the wind stress curl and Ekman pumping velocity. QuikSCAT consists of weekly-mean scalar ocean surface wind speed, ocean surface wind direction, and a rain flag/ collocated radiometer rain rate combination value, at a resolution of 0.25°× 0.25°. For averaged QuikSCAT data, wind speeds are scalar-averaged, while wind directions are vector-averaged.

    The monthly-mean dynamic topography from 2004 to 2012 was obtained from the Archiving, Validation, and Interpretation of Satellite Oceanographic (AVISO) data (AVISO, 2008), whose horizontal resolution is 0.25°× 0.25°.

    The MLD was defined as the depth at which the ocean potential density is different from the 5-m density by 0.125 kg m-3, following (de Boyer MontèGut et al., 2004). The geostrophic current and sea surface height were calculated using the monthly-mean Argo temperature and salinity data relative to 1200 m.

    Figure 2.  Climatology of Ekman pumping velocity (shading; positive upward; m s$^-1$) and MLD (showing contours of 50, 75, 100, 125, 150 and 175 m) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September.

    Figure 3.  Climatology of monthly-mean (a-f) net surface heat flux (shading; positive downward; W m$^-2$), (g-l) freshwater flux (shading; positive means evaporation is larger than precipitation), and MLD (as in Fig. 2) in (a, g) April, (b, h) May, (c, i) June, (d, j) July, (e, k) August, and (f, l) September.

2. Data
  • When studying the ML and mode water subduction in the North Pacific, the use of reanalysis data produces different results to Argo data because eddy-resolving model results are closer than non-eddy resolving model results to Argo observations (Xu et al., 2014). The temperature and salinity data used in this study were from Argo profiling floats in the Pacific Ocean from January 2004 to December 2012. Each float descended to a preprogrammed parking depth (typically 1000 m), drifted freely at that depth, and then ascended to the surface at a predetermined interval (usually 10 days) after descending to the maximum pressure (2000 m). These data were collected and quality-controlled before being made freely available by the International Argo Program and the national programs that contributed to it (http://www.argo.ucsd.edu, http://argo.jcommops.org). Only profiles with a quality flag of "1" and "2", indicating "good data" and "probably good data" respectively, were used in this study, resulting in a total of 20 571 profiles in the Pacific region (40°S-20°N, 120°E-60°W) and 1425 profiles in the subtropical Southeast Pacific (20°-40°S, 90°-120°W). Further quality control was conducted by the China Argo Real-time Data Center (http://www.argo.org.cn), including interpolating the profile data into vertical standard depths (of 48) by the Akima interpolation method, and then averaging in each 1°× 1° bin. Good coverage of Argo floats ensured the viability of our study in this region. The potential temperature and density (called temperature and density henceforth) referred to the surface are calculated from the interpolated temperature and salinity data.

    The monthly net surface heat flux data from 2004 to 2009, at 1°× 1° resolution, were acquired from the Objectively Analyzed Air-Sea Fluxes (OAFlux) project at the Woods Hole Oceanographic Institution (Yu et al., 2006). The objective analysis method used in OAFlux combines optimal satellite measurements and model reanalysis data (Yu and Weller, 2007; Yu, 2007). According to (Large and Yeager, 2008) and (Liu et al., 2010), the net surface heat flux of OAFlux is overestimated by about 5-10 W m-2 in the tropical region. Therefore, we first removed the extra heat flux from the shortwave radiation (keeping only 94.5% of its shortwave radiation) before we analyzed the net heat flux (Liu et al., 2010).

    Figure 1.  Climatology of monthly-mean MLD (showing contours of 50, 75, 100, 125, 150 and 175 m) and surface wind (arrows; m s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September.

    The OAFlux monthly 1°× 1° evaporation data (in cm yr-1) from January 2004 to December 2009 were also used. Meanwhile, the monthly precipitation data (in mm d-1), on a 2.5°× 2.5° grid, were from the CMAP dataset, which merges gauge data with five kinds of satellite estimates, from January 2004 to December 2009 (Xie and Arkin, 1997).

    The Quick Scatterometer (QuikSCAT) winds were used to calculate the wind stress curl and Ekman pumping velocity. QuikSCAT consists of weekly-mean scalar ocean surface wind speed, ocean surface wind direction, and a rain flag/ collocated radiometer rain rate combination value, at a resolution of 0.25°× 0.25°. For averaged QuikSCAT data, wind speeds are scalar-averaged, while wind directions are vector-averaged.

    The monthly-mean dynamic topography from 2004 to 2012 was obtained from the Archiving, Validation, and Interpretation of Satellite Oceanographic (AVISO) data (AVISO, 2008), whose horizontal resolution is 0.25°× 0.25°.

    The MLD was defined as the depth at which the ocean potential density is different from the 5-m density by 0.125 kg m-3, following (de Boyer MontèGut et al., 2004). The geostrophic current and sea surface height were calculated using the monthly-mean Argo temperature and salinity data relative to 1200 m.

    Figure 2.  Climatology of Ekman pumping velocity (shading; positive upward; m s$^-1$) and MLD (showing contours of 50, 75, 100, 125, 150 and 175 m) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September.

    Figure 3.  Climatology of monthly-mean (a-f) net surface heat flux (shading; positive downward; W m$^-2$), (g-l) freshwater flux (shading; positive means evaporation is larger than precipitation), and MLD (as in Fig. 2) in (a, g) April, (b, h) May, (c, i) June, (d, j) July, (e, k) August, and (f, l) September.

3. Seasonal deepening of the ML
3. Seasonal deepening of the ML
  • The climatology was defined using the average during the period 2004-12. Figure 1 shows the seasonal ML deepening processes in the subtropical Southeast Pacific from austral fall to spring. During austral fall (April and May), the MLD is only about 50-75 m (Figs. 1a and b) in the whole subtropical South Pacific. The ML deepens gradually in the approximate area (20°-30°S, 120°-90°W) during June to August, which is spatially non-uniform (Figs. 1c-e). The deep ML (>175 m) is located in the area (22°-32°S, 105°-90°W), reaching its maximum (200 m) near (27°S, 100°W) in September (early austral spring) (Fig. 1f). There is also a "shallow tongue" of the MLD (zonal band with shallow MLD) in the area (20°-30°N, 180°-120°W) (Figs. 2d-f). Similar features of the MLD seasonal cycle have been reported before, based on float observation data (Wong and Johnson, 2003) and limited CTD data (de Boyer MontèGut et al., 2004). After September, the MLD becomes shallower, since the deep winter ML is quickly replaced by a shallow seasonal thermocline as the surface temperature rises. As indicated by (Sato and Suga, 2009), the northern part of the deepest MLD in the subtropical Southeast Pacific is the formation region of the South Pacific eastern subtropical mode water.

    Corresponding to the seasonal deepening of the ML, the sea surface wind is shown in Fig. 1. There is a trade wind near the equator and a westerly jet in the middle latitudes in the subtropical South Pacific. There is negative (downwards) Ekman pumping in the whole basin, and larger absolute values of negative Ekman pumping are located in the west and east of the subtropical South Pacific (about 20°S), respectively. The negative Ekman pumping is key to determining the eastern boundary of the subtropical gyre, where the seasonal thermocline outcrops and the ventilated thermocline location is controlled by the Ekman pumping. Considering the upwelling near the west coast of South America, the Ekman pumping is the major factor determining the eastern boundary of the deep ML region (Huang, 2010).

  • The climatology was defined using the average during the period 2004-12. Figure 1 shows the seasonal ML deepening processes in the subtropical Southeast Pacific from austral fall to spring. During austral fall (April and May), the MLD is only about 50-75 m (Figs. 1a and b) in the whole subtropical South Pacific. The ML deepens gradually in the approximate area (20°-30°S, 120°-90°W) during June to August, which is spatially non-uniform (Figs. 1c-e). The deep ML (>175 m) is located in the area (22°-32°S, 105°-90°W), reaching its maximum (200 m) near (27°S, 100°W) in September (early austral spring) (Fig. 1f). There is also a "shallow tongue" of the MLD (zonal band with shallow MLD) in the area (20°-30°N, 180°-120°W) (Figs. 2d-f). Similar features of the MLD seasonal cycle have been reported before, based on float observation data (Wong and Johnson, 2003) and limited CTD data (de Boyer MontèGut et al., 2004). After September, the MLD becomes shallower, since the deep winter ML is quickly replaced by a shallow seasonal thermocline as the surface temperature rises. As indicated by (Sato and Suga, 2009), the northern part of the deepest MLD in the subtropical Southeast Pacific is the formation region of the South Pacific eastern subtropical mode water.

    Corresponding to the seasonal deepening of the ML, the sea surface wind is shown in Fig. 1. There is a trade wind near the equator and a westerly jet in the middle latitudes in the subtropical South Pacific. There is negative (downwards) Ekman pumping in the whole basin, and larger absolute values of negative Ekman pumping are located in the west and east of the subtropical South Pacific (about 20°S), respectively. The negative Ekman pumping is key to determining the eastern boundary of the subtropical gyre, where the seasonal thermocline outcrops and the ventilated thermocline location is controlled by the Ekman pumping. Considering the upwelling near the west coast of South America, the Ekman pumping is the major factor determining the eastern boundary of the deep ML region (Huang, 2010).

  • Ocean heat loss is an important factor for ML deepening. The negative net surface heat flux (ocean losing heat) during April to August is located between 22°S and 50°S, which contributes to deepening the ML in the whole ocean basin (Figs. 3a-f). However, a larger absolute value of negative net surface heat flux with a zonal band pattern appears around the southwest of, rather than southeast of, the subtropical Pacific, and its pattern corresponds to a shallower ML, meaning there may be other processes supporting the loss of heat by the ocean in the subtropical Southwest Pacific. In the subtropical Southeast Pacific, the ocean surface heat loss increases gradually from 22°S to 35°S in May and June (Figs. 3b and c), and the deepest ML (over 175 m) is located in the area (22°-32°S, 105°-90°W). According to the above discussion, we can conclude that the deep ML is constricted to the south of 22°S because the ocean heat loss occurs south of 22°S (Figs. 3a-e) during May to August. Since the heat loss is larger in the south of the maximum MLD region than in the north (Figs. 3b-e), it means there are other factors blocking the convective mixing process in the relatively shallow MLD region when the heat loss increases during June, July and August.

    According to Figs. 1c-e, the northward Ekman current corresponding to the westerly jet south of 30°S carries denser ocean surface water northwards in the shallow surface (<50 m; not shown), which contributes to vertical mixing between 30°S and 40°S during June to August. Since the Ekman current is stronger in the south than it is in the north, it can be inferred that the Ekman current has a negative effect on the northward deepening of the MLD in the area (30°-40°S, 120°-90°W). On the other hand, the southwestward Ekman current corresponding to the trade wind north of 20°S carries the light water southwestwards, which blocks the deepening of the ML south of 20°S. Thus, according to the horizontal density advection by the Ekman currents, we are unable to explain why the local maximum of the MLD appears in the area (22°-32°S, 105°-90°W).

    As we know, the South Pacific convergence zone extends southeastward from Northeast Australia, which can induce a precipitation belt. In order to identify the freshwater flux contribution to ML deepening in austral winter, the freshwater flux (evaporation minus precipitation; E-P) is shown in Figs. 3g-l. We can see that negative E-P flux, where fresher ocean surface water weakens the vertical mixing process, corresponds well to a shallower ML (Figs. 3g-l). During austral winter (June to September), larger positive E-P flux, which gives denser ocean surface water and enhances the vertical mixing process, appears in the subtropical shallower ML region west of the dateline and north of 20°S (Figs. 3i-l). Therefore, the freshwater flux only contributes to determine the pattern of the ML during austral fall, and its contribution is limited during austral winter.

    3.2.1 Role of the STCC in zonal density advection

    The role of potential-density horizontal advection(PDHA) transported by the geostrophic currents is investigated in this section. PHDA is defined as \[ {\rm PDHA}=u_{\rm g}\dfrac{\partial\rho}{\partial x}+v_{\rm g}\dfrac{\partial \rho}{\partial y}~,\] where ρ is sea water potential density relative to the surface, u g and v g are the zonal and meridional components of the geostrophic current velocity, respectively. Note that negative (positive) PDHA means heavier (lighter) water replaces the local water via horizontal advection.

    Figure 4.  Map of MLD in September (as in Fig. 2f), standard deviation of surface steric height (shading; units: m) and geostrophic current (averaged within 0-100 m, relative to 1200 m, vectors in m s$^-1$) in the South Pacific Ocean.

    Figure 5.  Depth-latitude section of zonal PDHA along 130$^\circ$-140$^\circ$W (shading; light shading for "zero"; kg m$^-3$ s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September. Superimposed are the MLD (red line) and isopycnals (black dashed lines).

    Figure 6.  Depth-latitude section of meridional PDHA along 91$^\circ$-95$^\circ$W (shading; light shading for "zero"; kg m$^-3$ s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September. Superimposed are the MLD (red line) and isopycnals (black dashed lines).

    Figure 7.  As in Fig. 6 but along 104$^\circ$-107$^\circ$W.

    In order to show whether the zonal density advection by the STCC can induce the "shallow tongue" of the MLD, the geostrophic current (0-100 m) and the standard deviation of the surface steric height are shown in Fig. 4. The STCC between 20°S and 30°S west of 100°W takes the light water from the warm pool in the subtropical Southwest Pacific to the subtropical Southeast Pacific (positive zonal PDHA) in the upper 100 m between 22°S and 30°S, across 130°-140°W. It is worth noting that the larger PDHA is in the upper layer (Fig. 5), which corresponds to a shallower ML, inducing a more stable stratification and weakening the mixing process. In addition, the positive advection in the upper layer contributes to the larger heat loss west of 100°W than east of 100°W between 22°S and 30°S, causing the maximum of negative net surface heat flux to locate around the subtropical Southwest Pacific. During austral winter, the MLD is deeper than 80 m within 130°-140°W, which means the freshwater flux has a limited role (Figs. 5c-f).

    Therefore, ocean heat loss contributes to deepening the ML, but the zonal PDHA by the STCC west of 100°W constrains the ML deepening process west of the deep ML region by transporting lighter water in the upper layer and stabilizing the stratification. The STCC blocks the westward extension of the deep ML, and its advection is the main mechanism behind the formation of the MLD "shallow tongue". In addition, the heat advection by the STCC is the main source of heat that forms a zonal band of large surface heat loss in the subtropical Southwest Pacific.

    3.2.2 Meridional PDHA

    The effect of PDHA on ocean stratification in the deep ML region (22°-32°S, 105°-90°W) is investigated in this subsection. Since the ocean loses heat south of 22°S, the northern boundary of the deep ML region should be 22°S. However, the maximum heat loss region is south of the maximum ML region during June to August (Figs. 3c-e) and the role of the Ekman current cannot determine the deep ML region within 22°-32°S. This inconsistency has not been discussed before. Since the meridional gradient of density is larger than the zonal gradient, the absolute value of the zonal PDHA is much smaller than that of the meridional PDHA in the deep ML region. In order to find out what determines the southern boundary of the deep ML region, the zonal-mean meridional PDHA along the eastern section (90°-95°W) and that along the western section (104°-107°W) of the maximum MLD region are shown in Figs. 6 and 7, respectively.

    In the eastern section, the maximum MLD (180 m) is located at 27°S in September (Fig. 6f). The meridional PDHA between 22°S and 27°S along this section is negative (approximately -6× 10-4 kg m-3 s-1) in the upper isopycnal layers (<25.5 kg m-3; called the "upper layer" hereafter) and positive (<4× 10-4 kg m-3 s-1) in the lower isopycnal layers (≥25.5 kg m-3; called the "lower layer" hereafter), because of the opposite meridional PDHAs between the two layers (Fig. 6). The largest vertical gradient of the meridional PDHA is about 10-3 kg m-3 s-1 in July (Fig. 6d), which weakens the ocean stratification significantly and strengthens the vertical mixing between 22°S and 27°S, where the southwestward Ekman current corresponding to the trade wind north of 20°S can strengthen the ocean stratification. The meridional PDHA is mostly negative (-2× 10-4 to -4× 10-4 kg m-3 s-1) south of 27°S, with its large absolute values in the layer beneath the ML and its maximum at 35°S during June to September (Figs. 6c-f). This strengthens the ocean stratification and weakens the vertical mixing, even though the ocean loses more heat south of 27°S, where the Ekman current corresponding to the westerly jet south of 30°S weakens the ocean stratification. Therefore, the combined effect of meridional PDHA and ocean heat loss determines the northern and southern boundaries in the western section of the maximum MLD region and sets the deepest MLD at 27°S.

    Along the 104°-107°W section (western section), the maximum MLD is located between 27° and 28°S, and is about 190 m in September (Fig. 7f). During May to August, the vertical gradient of meridional PDHA along this section induces a weaker stratification with negative meridional PDHA in the upper layer and positive meridional PDHA in the lower layer north of 27°S (Figs. 7b-e), similar to those along the eastern section (Fig. 6). There is, however, an obvious difference between the eastern and western sections in that a positive meridional PDHA exists between 28°S and 32°S in all months along the western section, and its difference between the lower and upper layers is only about 4× 10-4 kg m-3 s-1 (Fig. 7). Although the vertical gradient is much less than that in the north, it also weakens the ocean stratification. This is the reason why the maximum ML (about 190 m) region is wider in the meridional direction along the western section than along the eastern section. The meridional PDHA is negative south of 32°S, and larger absolute values lie beneath the ML, strengthening the ocean stratification and weakening the mixing process. This sets the southern boundary of the deep ML region at 32°S (Fig. 7). It is worth mentioning that the positive PDHA around 30°S is presumed to be related to a southward eddy-induced current (Fig. 4). In other words, the combined effect of meridional PDHA and net surface heat flux determines the northern and southern boundaries in the eastern section of the maximum MLD region and sets the deepest ML location at 27°S. Along this section, the effect of the Ekman current on ocean stratification is also opposite to that of the geostrophic current.

    Based on the above analysis, we can conclude that there are two factors determining the northern and southern boundaries of the deep MLD region: the gradually increasing ocean heat loss from 22°S to 35°S, and the meridional PDHA, which induces a stable stratification south of the deep ML region. These two factors combine to determine the location of the maximum MLD region along 27°-28°S. The northern boundary of the deep ML region is 22°S, and the southern boundary is 32°S.

  • Ocean heat loss is an important factor for ML deepening. The negative net surface heat flux (ocean losing heat) during April to August is located between 22°S and 50°S, which contributes to deepening the ML in the whole ocean basin (Figs. 3a-f). However, a larger absolute value of negative net surface heat flux with a zonal band pattern appears around the southwest of, rather than southeast of, the subtropical Pacific, and its pattern corresponds to a shallower ML, meaning there may be other processes supporting the loss of heat by the ocean in the subtropical Southwest Pacific. In the subtropical Southeast Pacific, the ocean surface heat loss increases gradually from 22°S to 35°S in May and June (Figs. 3b and c), and the deepest ML (over 175 m) is located in the area (22°-32°S, 105°-90°W). According to the above discussion, we can conclude that the deep ML is constricted to the south of 22°S because the ocean heat loss occurs south of 22°S (Figs. 3a-e) during May to August. Since the heat loss is larger in the south of the maximum MLD region than in the north (Figs. 3b-e), it means there are other factors blocking the convective mixing process in the relatively shallow MLD region when the heat loss increases during June, July and August.

    According to Figs. 1c-e, the northward Ekman current corresponding to the westerly jet south of 30°S carries denser ocean surface water northwards in the shallow surface (<50 m; not shown), which contributes to vertical mixing between 30°S and 40°S during June to August. Since the Ekman current is stronger in the south than it is in the north, it can be inferred that the Ekman current has a negative effect on the northward deepening of the MLD in the area (30°-40°S, 120°-90°W). On the other hand, the southwestward Ekman current corresponding to the trade wind north of 20°S carries the light water southwestwards, which blocks the deepening of the ML south of 20°S. Thus, according to the horizontal density advection by the Ekman currents, we are unable to explain why the local maximum of the MLD appears in the area (22°-32°S, 105°-90°W).

    As we know, the South Pacific convergence zone extends southeastward from Northeast Australia, which can induce a precipitation belt. In order to identify the freshwater flux contribution to ML deepening in austral winter, the freshwater flux (evaporation minus precipitation; E-P) is shown in Figs. 3g-l. We can see that negative E-P flux, where fresher ocean surface water weakens the vertical mixing process, corresponds well to a shallower ML (Figs. 3g-l). During austral winter (June to September), larger positive E-P flux, which gives denser ocean surface water and enhances the vertical mixing process, appears in the subtropical shallower ML region west of the dateline and north of 20°S (Figs. 3i-l). Therefore, the freshwater flux only contributes to determine the pattern of the ML during austral fall, and its contribution is limited during austral winter.

    3.2.1 Role of the STCC in zonal density advection

    The role of potential-density horizontal advection(PDHA) transported by the geostrophic currents is investigated in this section. PHDA is defined as \[ {\rm PDHA}=u_{\rm g}\dfrac{\partial\rho}{\partial x}+v_{\rm g}\dfrac{\partial \rho}{\partial y}~,\] where ρ is sea water potential density relative to the surface, u g and v g are the zonal and meridional components of the geostrophic current velocity, respectively. Note that negative (positive) PDHA means heavier (lighter) water replaces the local water via horizontal advection.

    Figure 4.  Map of MLD in September (as in Fig. 2f), standard deviation of surface steric height (shading; units: m) and geostrophic current (averaged within 0-100 m, relative to 1200 m, vectors in m s$^-1$) in the South Pacific Ocean.

    Figure 5.  Depth-latitude section of zonal PDHA along 130$^\circ$-140$^\circ$W (shading; light shading for "zero"; kg m$^-3$ s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September. Superimposed are the MLD (red line) and isopycnals (black dashed lines).

    Figure 6.  Depth-latitude section of meridional PDHA along 91$^\circ$-95$^\circ$W (shading; light shading for "zero"; kg m$^-3$ s$^-1$) in (a) April, (b) May, (c) June, (d) July, (e) August, and (f) September. Superimposed are the MLD (red line) and isopycnals (black dashed lines).

    Figure 7.  As in Fig. 6 but along 104$^\circ$-107$^\circ$W.

    In order to show whether the zonal density advection by the STCC can induce the "shallow tongue" of the MLD, the geostrophic current (0-100 m) and the standard deviation of the surface steric height are shown in Fig. 4. The STCC between 20°S and 30°S west of 100°W takes the light water from the warm pool in the subtropical Southwest Pacific to the subtropical Southeast Pacific (positive zonal PDHA) in the upper 100 m between 22°S and 30°S, across 130°-140°W. It is worth noting that the larger PDHA is in the upper layer (Fig. 5), which corresponds to a shallower ML, inducing a more stable stratification and weakening the mixing process. In addition, the positive advection in the upper layer contributes to the larger heat loss west of 100°W than east of 100°W between 22°S and 30°S, causing the maximum of negative net surface heat flux to locate around the subtropical Southwest Pacific. During austral winter, the MLD is deeper than 80 m within 130°-140°W, which means the freshwater flux has a limited role (Figs. 5c-f).

    Therefore, ocean heat loss contributes to deepening the ML, but the zonal PDHA by the STCC west of 100°W constrains the ML deepening process west of the deep ML region by transporting lighter water in the upper layer and stabilizing the stratification. The STCC blocks the westward extension of the deep ML, and its advection is the main mechanism behind the formation of the MLD "shallow tongue". In addition, the heat advection by the STCC is the main source of heat that forms a zonal band of large surface heat loss in the subtropical Southwest Pacific.

    3.2.2 Meridional PDHA

    The effect of PDHA on ocean stratification in the deep ML region (22°-32°S, 105°-90°W) is investigated in this subsection. Since the ocean loses heat south of 22°S, the northern boundary of the deep ML region should be 22°S. However, the maximum heat loss region is south of the maximum ML region during June to August (Figs. 3c-e) and the role of the Ekman current cannot determine the deep ML region within 22°-32°S. This inconsistency has not been discussed before. Since the meridional gradient of density is larger than the zonal gradient, the absolute value of the zonal PDHA is much smaller than that of the meridional PDHA in the deep ML region. In order to find out what determines the southern boundary of the deep ML region, the zonal-mean meridional PDHA along the eastern section (90°-95°W) and that along the western section (104°-107°W) of the maximum MLD region are shown in Figs. 6 and 7, respectively.

    In the eastern section, the maximum MLD (180 m) is located at 27°S in September (Fig. 6f). The meridional PDHA between 22°S and 27°S along this section is negative (approximately -6× 10-4 kg m-3 s-1) in the upper isopycnal layers (<25.5 kg m-3; called the "upper layer" hereafter) and positive (<4× 10-4 kg m-3 s-1) in the lower isopycnal layers (≥25.5 kg m-3; called the "lower layer" hereafter), because of the opposite meridional PDHAs between the two layers (Fig. 6). The largest vertical gradient of the meridional PDHA is about 10-3 kg m-3 s-1 in July (Fig. 6d), which weakens the ocean stratification significantly and strengthens the vertical mixing between 22°S and 27°S, where the southwestward Ekman current corresponding to the trade wind north of 20°S can strengthen the ocean stratification. The meridional PDHA is mostly negative (-2× 10-4 to -4× 10-4 kg m-3 s-1) south of 27°S, with its large absolute values in the layer beneath the ML and its maximum at 35°S during June to September (Figs. 6c-f). This strengthens the ocean stratification and weakens the vertical mixing, even though the ocean loses more heat south of 27°S, where the Ekman current corresponding to the westerly jet south of 30°S weakens the ocean stratification. Therefore, the combined effect of meridional PDHA and ocean heat loss determines the northern and southern boundaries in the western section of the maximum MLD region and sets the deepest MLD at 27°S.

    Along the 104°-107°W section (western section), the maximum MLD is located between 27° and 28°S, and is about 190 m in September (Fig. 7f). During May to August, the vertical gradient of meridional PDHA along this section induces a weaker stratification with negative meridional PDHA in the upper layer and positive meridional PDHA in the lower layer north of 27°S (Figs. 7b-e), similar to those along the eastern section (Fig. 6). There is, however, an obvious difference between the eastern and western sections in that a positive meridional PDHA exists between 28°S and 32°S in all months along the western section, and its difference between the lower and upper layers is only about 4× 10-4 kg m-3 s-1 (Fig. 7). Although the vertical gradient is much less than that in the north, it also weakens the ocean stratification. This is the reason why the maximum ML (about 190 m) region is wider in the meridional direction along the western section than along the eastern section. The meridional PDHA is negative south of 32°S, and larger absolute values lie beneath the ML, strengthening the ocean stratification and weakening the mixing process. This sets the southern boundary of the deep ML region at 32°S (Fig. 7). It is worth mentioning that the positive PDHA around 30°S is presumed to be related to a southward eddy-induced current (Fig. 4). In other words, the combined effect of meridional PDHA and net surface heat flux determines the northern and southern boundaries in the eastern section of the maximum MLD region and sets the deepest ML location at 27°S. Along this section, the effect of the Ekman current on ocean stratification is also opposite to that of the geostrophic current.

    Based on the above analysis, we can conclude that there are two factors determining the northern and southern boundaries of the deep MLD region: the gradually increasing ocean heat loss from 22°S to 35°S, and the meridional PDHA, which induces a stable stratification south of the deep ML region. These two factors combine to determine the location of the maximum MLD region along 27°-28°S. The northern boundary of the deep ML region is 22°S, and the southern boundary is 32°S.

4. Conclusions and discussion
  • A qualitative investigation of the relative importance of Ekman pumping, net surface heat flux, freshwater flux, and horizontal density advection in the seasonal deepening of the ML in the subtropical Southeast Pacific was conducted in this study using the Argo profile data during 2004-12. During austral fall and winter, the ML deepens gradually around (20°-32°S, 120°-90°W). The deep ML (>175 m) is located in the area (22°-32°S, 105°-90°W), reaching its seasonal maximum (200 m) near (27°-28°S, 100°W) in September.

    The downward Ekman pumping has two local maximum regions in the west and east subtropical South Pacific, respectively. The eastern boundary of the downward Ekman pumping is key in determining the eastern boundary location of the deep ML region.

    The freshwater flux only contributes to the ML pattern during austral fall, and its contribution is limited during austral winter.

    The zonal PDHA by the STCC places a warm, freshwater cap in the upper layer, which strengthens the upper-ocean stratification and determines the western boundary of the deep ML region. The STCC blocks the westward extension of the deep ML region and forms an MLD "shallow tongue" along the STCC. This discovery implies a close relationship between the STCC and the ML in the subtropical Southeast Pacific, and explains the dynamic mechanism for the zonal band of large heat. The northern and southern boundaries of the deep ML region are determined by the combined effect of net surface heat flux and meridional PDHA in the subtropical Southeast Pacific. The gradual increase in ocean heat loss from 22°S to 33°S deepens the ML, while the meridional PDHA by the subtropical gyre strengthens the upper-ocean stratification south of 27°-28°S and weakens the stratification north of 27°-28°S, which sets the boundaries in the north and south.

    Although yielding far more profiles than historical CTD collections, the Argo coverage is not sufficient for a quantitative study on the seasonal variation of the ML. This is because, to date, there have not been any good-quality measurements of the seasonal variation of the vertical velocity and entrainment velocity at the bottom of the ML in the subtropical southeast Pacific Ocean. Unfortunately, both vertical velocity and entrainment velocity calculated based on the geostrophic current contain large errors. In addition, given the deficiency in vertical velocity and entrainment velocity observation, we cannot determine which ocean numerical models have the greater ability to simulate the mixing process successfully. Thus, the present-reported results are merely qualitative. Quantitative studies using long-term ocean observations and high-quality simulations from numerical models are necessary in the future.

4. Conclusions and discussion
  • A qualitative investigation of the relative importance of Ekman pumping, net surface heat flux, freshwater flux, and horizontal density advection in the seasonal deepening of the ML in the subtropical Southeast Pacific was conducted in this study using the Argo profile data during 2004-12. During austral fall and winter, the ML deepens gradually around (20°-32°S, 120°-90°W). The deep ML (>175 m) is located in the area (22°-32°S, 105°-90°W), reaching its seasonal maximum (200 m) near (27°-28°S, 100°W) in September.

    The downward Ekman pumping has two local maximum regions in the west and east subtropical South Pacific, respectively. The eastern boundary of the downward Ekman pumping is key in determining the eastern boundary location of the deep ML region.

    The freshwater flux only contributes to the ML pattern during austral fall, and its contribution is limited during austral winter.

    The zonal PDHA by the STCC places a warm, freshwater cap in the upper layer, which strengthens the upper-ocean stratification and determines the western boundary of the deep ML region. The STCC blocks the westward extension of the deep ML region and forms an MLD "shallow tongue" along the STCC. This discovery implies a close relationship between the STCC and the ML in the subtropical Southeast Pacific, and explains the dynamic mechanism for the zonal band of large heat. The northern and southern boundaries of the deep ML region are determined by the combined effect of net surface heat flux and meridional PDHA in the subtropical Southeast Pacific. The gradual increase in ocean heat loss from 22°S to 33°S deepens the ML, while the meridional PDHA by the subtropical gyre strengthens the upper-ocean stratification south of 27°-28°S and weakens the stratification north of 27°-28°S, which sets the boundaries in the north and south.

    Although yielding far more profiles than historical CTD collections, the Argo coverage is not sufficient for a quantitative study on the seasonal variation of the ML. This is because, to date, there have not been any good-quality measurements of the seasonal variation of the vertical velocity and entrainment velocity at the bottom of the ML in the subtropical southeast Pacific Ocean. Unfortunately, both vertical velocity and entrainment velocity calculated based on the geostrophic current contain large errors. In addition, given the deficiency in vertical velocity and entrainment velocity observation, we cannot determine which ocean numerical models have the greater ability to simulate the mixing process successfully. Thus, the present-reported results are merely qualitative. Quantitative studies using long-term ocean observations and high-quality simulations from numerical models are necessary in the future.

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