• Adler, R. F., and Coauthors, 2003: The Version-2 Global Precipitation Climatology Project (GPCP) monthly precipitation analysis (1979–present). J. Hydrometeor., 4, 11471167.

    • Search Google Scholar
    • Export Citation
  • Barnier, B., , L. Siefridt, , and P. Marchesiello, 1995: Thermal forcing for a global ocean circulation model using a three-year climatology of ECMWF analysis. J. Mar. Syst., 6, 363380.

    • Search Google Scholar
    • Export Citation
  • Delcroix, T., , and M. J. McPhaden, 2002: Interannual sea surface salinity and temperature changes in the western Pacific warm pool during 1992–2000. J. Geophys. Res., 107, 8002, doi:10.1029/2001JC000862.

    • Search Google Scholar
    • Export Citation
  • Fine, R. A., , R. Lukas, , F. M. Bingham, , M. J. Warner, , and R. H. Gammon, 1994: The western equatorial Pacific: A water mass crossroads. J. Geophys. Res., 99 (C12), 25 06325 080.

    • Search Google Scholar
    • Export Citation
  • Fine, R. A., , K. A. Maillet, , K. F. Sullivan, , and D. Willey, 2001: Circulation and ventilation flux of the Pacific Ocean. J. Geophys. Res., 106 (C10), 22 15922 178.

    • Search Google Scholar
    • Export Citation
  • Fukumori, I., , T. Lee, , B. Cheng, , and D. Menemnlis, 2004: The origin, pathway, and destination of Niño-3 water estimated by a simulated passive tracer and its adjoint. J. Phys. Oceanogr., 34, 582604.

    • Search Google Scholar
    • Export Citation
  • Gao, S., , T. Qu, , and I. Fukumori, 2011: Effects of mixing on the subduction of South Pacific waters. Dyn. Atmos. Oceans, 51, 4554.

  • Gent, P. R., , and J. C. McWilliams, 1990: Isopycnal mixing in ocean circulation models. J. Phys. Oceanogr., 20, 150155.

  • Grenier, M., , S. Cravatte, , B. Blanke, , C. Menkes, , A. Koch-Larrouy, , F. Durand, , A. Melet, , and C. Jeandel, 2011: From the western boundary currents to the Pacific Equatorial Undercurrent: Modeled pathways and water mass evolutions. J. Geophys. Res., 116, C12044, doi:10.1029/2011JC007477.

    • Search Google Scholar
    • Export Citation
  • Gu, D., , and S. G. H. Philander, 1997: Interdecadal climate fluctuations that depend on exchange between the tropics and extratropics. Science, 275, 805807.

    • Search Google Scholar
    • Export Citation
  • Hanawa, K., , and L. D. Talley, 2001: Mode waters. Ocean Circulation and Climate, G. Sieldler, J. Church, and J. Gould, Eds., Academic Press, 373–386.

  • Huang, R. X., , and B. Qiu, 1998: The structure of the wind-driven circulation in the subtropical South Pacific Ocean. J. Phys. Oceanogr., 28, 11731186.

    • Search Google Scholar
    • Export Citation
  • Johnson, G. C., , and M. J. McPhaden, 1999: Interior pycnocline flow from the subtropical to the equatorial Pacific Ocean. J. Phys. Oceanogr.,29, 3073–3089.

  • Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77, 437371.

  • Kessler, W. S., 2006: The circulation of the eastern tropical Pacific: A review. Prog. Oceanogr., 69, 181217.

  • Kessler, W. S., , and B. A. Taft, 1987: Dynamic heights and zonal geostrophic transports in the central tropical Pacific during 1979–1984. J. Phys. Oceanogr., 17, 97122.

    • Search Google Scholar
    • Export Citation
  • Large, W. G., , J. C. McWilliams, , and S. C. Doney, 1994: Oceanic vertical mixing: A review and a model with a nonlocal boundary-layer parameterization. Rev. Geophys., 32, 363403.

    • Search Google Scholar
    • Export Citation
  • Lindstrom, E., , R. Lukas, , R. Fine, , E. Firing, , S. J. Godfrey, , G. Meyers, , and M. Tsuchiya, 1987: The western equatorial Pacific Ocean circulation study. Nature, 330, 533537.

    • Search Google Scholar
    • Export Citation
  • Lukas, R., , and E. Lindstrom, 1991: The mixed layer of the western equatorial Pacific Ocean. J. Geophys. Res., 96, 33433358.

  • Luo, J.-J., , and T. Yamagata, 2001: Long-term El Niño–Southern Oscillation (ENSO)-like variation with special emphasis on the South Pacific. J. Geophys. Res., 106 (C10), 22 21122 227.

    • Search Google Scholar
    • Export Citation
  • Maes, C., 2008: On the ocean salinity stratification observed at the eastern edge of the equatorial Pacific warm pool. J. Geophys. Res., 113, C03027, doi:10.1029/2007JC004297.

    • Search Google Scholar
    • Export Citation
  • Marshall, J. C., , A. Adcroft, , C. Hill, , L. Perelman, , and C. Heisey, 1997: A finite-volume, incompressible Navier Stokes model for studies of the ocean on parallel computers. J. Geophys. Res., 102 (C3), 57535766.

    • Search Google Scholar
    • Export Citation
  • McCreary, J. P., , and P. Lu, 1994: On the interaction between the subtropical and the equatorial oceans: The subtropical cell. J. Phys. Oceanogr., 24, 466497.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , and J. Picaut, 1990: El Niño–Southern Oscillation displacement of the western equatorial Pacific warm pool. Science, 250, 13851388.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , and D. Zhang, 2002: Slowdown of the meridional overturning circulation in the upper Pacific Ocean. Nature, 415, 603608.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , S. E. Zebiak, , and M. H. Glantz, 2006: ENSO as an integrating concept in earth. Science, 314, 17401745.

  • O'Connor, B. M., , R. A. Fine, , K. A. Maillet, , and D. B. Olson, 2002: Formation rates of subtropical underwater in the Pacific Ocean. Deep-Sea Res., 49, 15711590.

    • Search Google Scholar
    • Export Citation
  • O'Connor, B. M., , R. A. Fine, , and D. B. Olson, 2005: A global comparison of subtropical underwater formation rate. Deep-Sea Res., 52, 15691590.

    • Search Google Scholar
    • Export Citation
  • Palmer, T. N., , and D. A. Mansfield, 1984: Response of two atmospheric general circulation models to sea surface temperature anomalies in the tropical east and west Pacific. Nature, 310, 483.

    • Search Google Scholar
    • Export Citation
  • Picaut, J., , M. Loualalen, , C. Menkes, , T. Delcroix, , and M. J. McPhaden, 1996: Mechanism of the zonal displacement of the Pacific warm pool: Implications for ENSO. Science, 274, 14861489.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , H. Mitsudera, , and T. Yamagata, 1999: A climatology of the circulation and water mass distribution near the Philippine coast. J. Phys. Oceanogr., 29, 14881505.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , S. Gao, , I. Fukumori, , R. A. Fine, , and E. J. Lindstrom, 2008: Subduction of South Pacific waters. Geophys. Res. Lett., 35, L02610, doi:10.1029/2007GL032605.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , S. Gao, , I. Fukumori, , R. A. Fine, , and E. J. Lindstrom, 2009: Origin and pathway of equatorial 13°C Water in the Pacific identified by a simulated passive tracer and its adjoint. J. Phys. Oceanogr., 39, 18361853.

    • Search Google Scholar
    • Export Citation
  • Redi, M., 1982: Oceanic isopycnal mixing by coordinate rotation. J. Phys. Oceanogr., 12, 11541158.

  • Reynolds, R. W., , and T. M. Smith, 1994: Improved global sea surface temperature analyses using optimum interpolation. J. Climate, 7, 929948.

    • Search Google Scholar
    • Export Citation
  • Rodgers, K. B., , B. Blanke, , G. Mades, , O. Aumont, , P. Ciais, , and J.-C. Dutay, 2003: Extratropical sources of equatorial Pacific upwelling in an OGCM. Geophys. Res. Lett., 30, 1084, doi:10.1029/2002GL016003.

    • Search Google Scholar
    • Export Citation
  • Roemmich, D., , and B. Cornuelle, 1992: The subtropical mode waters of the South Pacific Ocean. J. Phys. Oceanogr., 22, 11781187.

  • Singh, A., , T. Delcroix, , and S. Cravatte, 2011: Contrasting the flavors of El Niño–Southern Oscillation using sea surface salinity observations. J. Geophys. Res., 116, C06016, doi:10.1029/2010JC006862.

    • Search Google Scholar
    • Export Citation
  • Slutz, R. J., , S. J. Lubker, , J. D. Hiscox, , S. D. Woodruff, , R. L. Jenne, , D. H. Joseph, , P. M. Steurer, , and J. D. Elms, 1985: Comprehensive ocean–atmosphere data set; Release 1. NOAA Environmental Research Laboratories, Climate Research Program Rep. NTIS PB86105723, 268 pp.

  • Tsuchiya, M., , and L. D. Talley, 1996: Water property distributions along an eastern Pacific hydrographic section at 135°W. J. Mar. Res., 54, 541564.

    • Search Google Scholar
    • Export Citation
  • Tsuchiya, M., , R. Lukas, , R. A. Fine, , E. Firing, , and E. Lindstrom, 1989: Source waters of the Pacific Equatorial Undercurrent. Prog. Oceanogr.,23, 101–147.

  • White, W. B., , and D. R. Cayan, 1998: Quasi-periodicity and global symmetries in interdecadal upper ocean temperature variability. J. Geophys. Res., 103 (C10), 21 33521 354.

    • Search Google Scholar
    • Export Citation
  • Yu, L., 2007: Global variations in oceanic evaporation (1958–2005): The role of the changing wind speed. J. Climate, 20, 53765390.

  • Zhang, R.-H., , L. M. Rothstein, , and A. J. Busalacchi, 1998: Origin of upper-ocean warming and El Niño change on decadal scales in the tropical Pacific Ocean. Nature, 391, 879883.

    • Search Google Scholar
    • Export Citation
  • View in gallery

    Annual-mean SSS (psu) (shading) with winter (September) sea surface density (kg m−3) superimposed (white contours) from (a) WOA09 and (b) ECCO in the South Pacific. The horizontal SSS maximum is marked by the black box in these, and the following, figures.

  • View in gallery

    Annual-mean surplus of evaporation over precipitation (EP) (shading) from (a) OAFlux and GPCP and (b) ECCO with surface wind stress (N m−2; black arrows, reference arrow provided) and its curl (N m−3; white contours, contour interval: 10−7 N m−3) from NCEP superimposed in the subtropical South Pacific.

  • View in gallery

    Winter (September) MLD (m; shading) with SSS (psu; white contours) from (a) WOA09 and (b) ECCO. The 1-yr Lagrangian trajectories of water particles are also presented, in which the solid circles indicate the start points of water particles at the base of winter mixed layer.

  • View in gallery

    (a),(b) Subduction rate as a result of (c),(d) vertical pumping and (e),(f) lateral induction in the South Pacific between 10° and 40°S (m yr−1) from (left) WOA09 and (right) ECCO and its components. Black contours in top panels represent annual-mean SSS.

  • View in gallery

    Subduction volume (Sv) against (a),(c) surface density and (b),(d) salinity from (top) WOA09 and (bottom) ECCO in the subtropical South Pacific (10°–40°S). The light shading represents the subduction of salinity maximum water in the area as indicated by the black box in Fig. 1.

  • View in gallery

    Vertically integrated passive tracer (ATU m−2) distribution at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, in the Pacific. The black box indicates the SSS maximum region.

  • View in gallery

    Distribution of passive tracer concentration (1013 ATU) in temperature and salinity space at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, in the Pacific.

  • View in gallery

    (a) Time evolution and (b) spatial distribution of tracer-weighted density of the water. (c),(d) As in (a),(b), but for tracer-weighted salinity. Here, (a),(c) are the spatial averages (solid line) with std dev (gray) over the entire domain, and (b),(d) are the temporal averages (shading) over the entire period.

  • View in gallery

    Spatial distribution of passive tracer (105 ATU m−2) that has entered the equatorial region (shading) with mean isopycnals (kg m−3) superimposed (contours) at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, through 8°S. The interior and western boundary pathways are indicated by the high concentrations centered at 160°W and 150°E, respectively.

  • View in gallery

    (a) Time evolution of passive tracer (heavy solid black) that has entered the equatorial region through 8°S against year and (b) its distribution against longitude at year 13. The light solid and dashed lines in (a) represent components through the western boundary (west of 160°E) and interior (east of 160°E) pathways. The units on the vertical axis are 1014 ATU in (a) and 108 ATU m−1 in (b).

  • View in gallery

    Distribution of passive tracer (10−3 ATU m−2) at 160°E between 5°S and 5°N (shading) at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, superimposed with the zonal velocity in the upper 300 m. Positive values represent eastward flow.

  • View in gallery

    Distribution of passive tracer (ATU m−2) that has entered the mixed layer in equatorial Pacific between 5°S and 5°N at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, with annual-mean SSS (white contours) superimposed.

  • View in gallery

    Time evolution of passive tracer (1014 ATU) that has entered the mixed layer in the equatorial Pacific (heavy solid) and its component east (dashed) and west (solid) of the date line.

All Time Past Year Past 30 Days
Abstract Views 0 0 0
Full Text Views 24 24 10
PDF Downloads 26 26 10

Subduction of South Pacific Tropical Water and Its Equatorward Pathways as Shown by a Simulated Passive Tracer

View More View Less
  • 1 International Pacific Research Center, School of Ocean and Earth Science and Technology, University of Hawaii at Manoa, Honolulu, Hawaii
  • 2 Rosenstiel School of Marine and Atmospheric Science, University of Miami, Miami, Florida
© Get Permissions
Full access

Abstract

This study investigates the subduction of South Pacific Tropical Water (SPTW) and its equatorward pathways using a simulated passive tracer of the consortium Estimating the Circulation & Climate of the Ocean (ECCO). The results show that approximately 5.8 Sv (1 Sv ≡ 106 m3 s−1) of the SPTW is formed in the subtropical South Pacific Ocean within the density range between 24.0 and 25.0 kg m−3, of which about 87% is due to vertical pumping and 13% is due to lateral induction, comparing reasonably well with estimates from climatological data. Once subducted, most SPTW spreads in the subtropical South Pacific. Because of the presence of mixing, some portion of the water is transformed, and its tracer-weighted density steadily increases from an initial value of 24.4 to nearly 25.0 kg m−3 after 13 years of integration. Approximately 42% of the water makes its way into the equatorial Pacific, either through the western boundary or interior pathway. The two equatorward pathways are essentially of equal importance. A large (~70%) portion of the SPTW entering the equatorial region resurfaces in the central equatorial Pacific. The potential impacts of the resurfacing SPTW on the equatorial thermocline and surface stratification are discussed.

School of Ocean and Earth Science and Technology Contribution Number 8934 and International Pacific Research Center Contribution Number IPRC-982.

Current affiliation: Institute of Oceanology, Chinese Academy of Sciences, Qingdao, China.

Corresponding author address: Dr. Tangdong Qu, International Pacific Research Center, SOEST, University of Hawaii at Manoa, 1680 East-West Road, Honolulu, HI 96822. E-mail: tangdong@hawaii.edu

Abstract

This study investigates the subduction of South Pacific Tropical Water (SPTW) and its equatorward pathways using a simulated passive tracer of the consortium Estimating the Circulation & Climate of the Ocean (ECCO). The results show that approximately 5.8 Sv (1 Sv ≡ 106 m3 s−1) of the SPTW is formed in the subtropical South Pacific Ocean within the density range between 24.0 and 25.0 kg m−3, of which about 87% is due to vertical pumping and 13% is due to lateral induction, comparing reasonably well with estimates from climatological data. Once subducted, most SPTW spreads in the subtropical South Pacific. Because of the presence of mixing, some portion of the water is transformed, and its tracer-weighted density steadily increases from an initial value of 24.4 to nearly 25.0 kg m−3 after 13 years of integration. Approximately 42% of the water makes its way into the equatorial Pacific, either through the western boundary or interior pathway. The two equatorward pathways are essentially of equal importance. A large (~70%) portion of the SPTW entering the equatorial region resurfaces in the central equatorial Pacific. The potential impacts of the resurfacing SPTW on the equatorial thermocline and surface stratification are discussed.

School of Ocean and Earth Science and Technology Contribution Number 8934 and International Pacific Research Center Contribution Number IPRC-982.

Current affiliation: Institute of Oceanology, Chinese Academy of Sciences, Qingdao, China.

Corresponding author address: Dr. Tangdong Qu, International Pacific Research Center, SOEST, University of Hawaii at Manoa, 1680 East-West Road, Honolulu, HI 96822. E-mail: tangdong@hawaii.edu

1. Introduction

The South Pacific Tropical Water (SPTW), also referred to as the Subtropical Underwater (STUW), is identified by a shallow salinity maximum (>35.6 psu) centered at 20°S, 125°W (cf. O'Connor et al. 2005). Its density ranges between about 24.6 and 25.4 kg m−3, slightly overlying the eastern subtropical mode water in the South Pacific (e.g., Tsuchiya and Talley 1996; O'Connor et al. 2002). The SPTW is formed by subduction within the subtropical gyre. By its definition, the subduction of the SPTW consists of two components. One is the contribution due to vertical pumping at the base of the mixed layer, and the other is related to the slope of the mixed layer base that has been termed lateral induction by previous studies [Huang and Qui (1998), and references therein].

The subduction of the SPTW is largely determined by the basin-scale wind in the South Pacific, either through vertical pumping or lateral induction. O'Connor et al. (2002) examined the SPTW subduction rate in two periods, 1988–92 and 1992–96, and found that from the first to second period there was a shift to a higher annual subduction rate from 25 to 40 m yr−1. Vertical pumping and lateral induction contributed almost equally to the subduction rate during the first period, while vertical pumping greatly exceeded lateral induction during the second period. The increase in vertical pumping during the second period was consistent with stronger wind stress curl because of the occurrence of El Niño–Southern Oscillation (ENSO) in the central/eastern parts of the subtropical South Pacific (White and Cayan 1998). This result clearly demonstrated a connection between the SPTW subduction rate and ENSO (e.g., O'Connor et al. 2002, 2005).

The subduction of the SPTW, in turn, provides important water sources for the subtropical cell in the South Pacific, which involves water subducting in the subtropics, equatorward flow in the upper thermocline, upwelling at the equator, and poleward return flow in the surface layer that closes the circulation (e.g., McCreary and Lu 1994). Some portion of the SPTW is believed to resurface and eventually affect the equatorial thermocline and sea surface temperature in the equatorial Pacific (e.g., Lindstrom et al. 1987; Tsuchiya et al. 1989; Rodgers et al. 2003), thereby playing a potentially important role in modulating ENSO and possibly longer time-scale variability (e.g., Gu and Philander 1997; Zhang et al. 1998; McPhaden and Zhang 2002). Compared with its North Pacific counterpart, the SPTW has a higher subduction rate, and its impacts on the equatorial thermocline and sea surface temperature are expected to be faster, because of the water's significant interior pathway toward the equator (e.g., Johnson and McPhaden 1999; Fine et al. 2001; O'Connor et al. 2002, 2005). However, because of the lack of sufficient observations, the detailed pathways of this influence in the upper thermocline remain poorly understood.

Taking advantage of rapid advances in ocean modeling, this study investigates the subduction of the SPTW and its equatorward pathways in the South Pacific, using a simulated passive tracer of the consortium for Estimating the Circulation & Climate of the Ocean (ECCO) [Fukumori et al. (2004), and references therein]. The focus of this study is on the climatological pathways of the SPTW. The time dependence of these pathways will be investigated by a separate study. The rest of the paper is organized as follows. A brief description of the model and method of analysis is presented in section 2. The salinity maximum simulated by the model is described and compared with observations in section 3. The subduction of the SPTW and the processes that control it are discussed in section 4. The pathways of the SPTW are examined in section 5. The resurfacing of the SPTW in the equatorial Pacific is discussed in section 6. Results are summarized and discussed in section 7.

2. Data and model description

The ECCO model, based on the Massachusetts Institute of Technology (MIT) General Circulation Model (Marshall et al. 1997), is nearly global, extending from 80°S to 80°N. Horizontal grid spacing is 1° globally, except within 20° of the equator, in which meridional grid spacing is gradually reduced to 0.3° within 10° of the equator. Vertical grid spacing is 10 m within 150 m of the surface, gradually increasing to 400 m toward the bottom of the domain. The vertical mixing scheme of Large et al. (1994) is employed for realistic simulation of near-surface mixing processes. Mixing effects of mesoscale eddies are represented using the Redi (1982) isoneutral mixing scheme and the Gent and McWilliams (1990) parameterization.

The model is initially at rest with climatological temperature and salinity, and spun up for 10 years, forced by time-mean seasonal wind stress and heat flux climatologies based on Comprehensive Ocean–Atmosphere Data Set (COADS) (Slutz et al. 1985). Following spin up, the model is forced from 1980 to the present by wind stress, heat flux, and evaporation minus precipitation estimates of the National Centers for Environmental Prediction (NCEP)–National Center for Atmospheric Research (NCAR) reanalysis project (Kalnay et al. 1996). A time-mean correction to the NCEP fluxes is applied in which the 1980–97 NCEP means are replaced by corresponding 1945–93 means of COADS because the time-mean NCEP fluxes are too weak in comparison to other estimates. Sea surface temperature is relaxed to observed temperatures (Reynolds and Smith 1994) with a spatially varying relaxation coefficient computed from the NCEP–NCAR product using the method employed by Barnier et al. (1995). The equivalent of a freshwater flux is implemented by relaxing surface salinity to climatological values with a 60-day relaxation coefficient. Results from a simulation for the period 1993–2006 are analyzed in the following sections.

The “offline” simulated passive tracer method of Fukumori et al. (2004) will be used to quantify the pathways of the SPTW and its connection with the equatorial circulation. In the context of a GCM, the temporal evolution of a passive tracer is dictated by the same advection–diffusion equation as temperature and salinity. If a particular patch of water is uniformly initialized by a passive tracer with no other sources or sinks, the subsequent movement of this tracer indicates the pathway of the initial patch of water. In particular, the relative magnitude of the tracer at a given location to the value of the initial patch describes the concentration of this initial water mass at the location in question. The simulated passive tracer provides unambiguous means to identify the pathways of the SPTW exactly as simulated by models, whereas the traditional advective particle methods of tracing water masses do not account for mixing (e.g., Fukumori et al. 2004; Qu et al. 2009; Gao et al. 2011).

3. Salinity maximum

The SPTW has been studied by many earlier studies (e.g., Tsuchiya and Talley 1996; O'Connor et al. 2002, 2005). At the sea surface, the horizontal extent of salinity maximum in the subtropical South Pacific has a density range between about 24.0 and 25.0 kg m−3 in winter (Figs. 1a,b). An average density of 24.48 kg m−3 is observed over the area where sea surface salinity (SSS) is higher than 36.2 psu. The SPTW is also identified as a vertical salinity maximum in the upper subtropical South Pacific (e.g., O'Connor et al. 2005). The average density of vertical salinity maximum in the region (marked by the black box in Fig. 1) is 24.44 kg m−3, which is nearly identical to the value of winter horizontal surface salinity maximum. In general, the observed characteristics of the SPTW are well simulated by ECCO (Fig. 1b). As in the observation, the simulated annual-mean surface salinity maximum lies at about 20°S, 125°W. In winter, the density of this SSS maximum averaged over the black box (Fig. 1) is 24.37 kg m−3, in reasonable agreement with observations. The high-salinity signature of the SPTW forms a marked contrast with lower SSS in the tropical and subpolar regions (Fig. 1b).

Fig. 1.
Fig. 1.

Annual-mean SSS (psu) (shading) with winter (September) sea surface density (kg m−3) superimposed (white contours) from (a) WOA09 and (b) ECCO in the South Pacific. The horizontal SSS maximum is marked by the black box in these, and the following, figures.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

Surface freshwater flux is a key process influencing the ocean's SSS. The region's surface freshwater flux is dominated by the excess of evaporation over precipitation (EP). The EP is predominantly positive in the subtropical South Pacific, except for a southeastward-directed precipitation band associated with the South Pacific Convergence Zone (SPCZ). Previous studies have shown that the meandering and variability of the SPCZ greatly affects the SSS in the subtropical South Pacific (e.g., Luo and Yamagata 2001). The EP attains its maximum in the eastern subtropical gyre, lying to the northeastward side of the SSS maximum. This displacement reflects the influence of ocean dynamics in governing the salinity maximum. Careful examination of surface wind stress indicates that strong southeasterly wind prevails in much of the EP maximum region (Fig. 2a). With a southwestward-directed SSS gradient (Fig. 1), the southeasterly wind produces a freshwater advection toward the EP maximum region that balances some of the EP received from the atmosphere, thus leading to a shift of SSS maximum to its southwestward side.

Fig. 2.
Fig. 2.

Annual-mean surplus of evaporation over precipitation (EP) (shading) from (a) OAFlux and GPCP and (b) ECCO with surface wind stress (N m−2; black arrows, reference arrow provided) and its curl (N m−3; white contours, contour interval: 10−7 N m−3) from NCEP superimposed in the subtropical South Pacific.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The EP distribution in the South Pacific is also well simulated by the model (Fig. 2b). Averaged over the area indicated by the black box (Fig. 1), the model's annual-mean EP (73.2 cm yr−1) is only slightly smaller than that (89.4 cm yr−1) derived from observations [objectively analyzed air–sea fluxes (OAFlux)–Global Precipitation Climatology Project (GPCP); Yu 2007; Adler et al. 2003]. Some of this discrepancy is due to the relaxation of SSS to climatological values in the model. Averaged over the SSS maximum region, the contribution associated with the SSS relaxation is about −4.6 cm yr−1, which is significantly smaller than the total value simulated by the model.

4. Subduction

The salinity maximum water lies in the eastern/central subtropical gyre, where vertical Ekman pumping is predominantly downward (Fig. 2). The downward Ekman pumping provides a favorable condition for the formation of the SPTW. Lateral induction, defined as the rate at which the water is swept beneath the shoaling mixed layer base by horizontal currents [Huang and Qiu (1998), and references therein], is another process which has been shown to be important in modulating the SPTW subduction rate (e.g., O'Connor et al. 2005; Qu et al. 2008). By its definition, lateral induction is sensitive to the ocean's winter mixed layer depth (MLD) configuration, that is, the angle between the winter MLD gradient and ocean current. Before proceeding to the estimate of the SPTW subduction rate, we first examine the winter MLD distribution simulated by the model and compare it with observations. The definition of the MLD is based on a variable potential density criterion (Lukas and Lindstrom 1991), which determines the depth as where sq is equal to the potential density near the sea surface plus an increment of 0.1 kg m−3.

The MLD in the South Pacific reaches its seasonal maximum in September (Fig. 3). In general, the simulated winter MLD shows the same pattern as observations, despite some quantitative discrepancies. It is generally shallow (<60 m) along the northern rim of the subtropical gyre, and gradually gets deeper toward higher latitudes. In the southern subtropical region, near 40°S, the winter MLD can reach as deep as 200 m, in response to the enhanced surface cooling. A local maximum (~150 m) of MLD is visible in the eastern subtropical gyre, lying around the eastern edge of the SSS maximum (Fig. 1). This local MLD maximum is a consequence of surface buoyancy forcing associated with the large excess of evaporation over precipitation (Fig. 2), combined with the effect of downward Ekman pumping (e.g., O'Connor et al. 2002, 2005).

Fig. 3.
Fig. 3.

Winter (September) MLD (m; shading) with SSS (psu; white contours) from (a) WOA09 and (b) ECCO. The 1-yr Lagrangian trajectories of water particles are also presented, in which the solid circles indicate the start points of water particles at the base of winter mixed layer.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The MLD in the South Pacific, as in many other parts of the global ocean, exhibits a large annual excursion. After reaching its seasonal maximum in winter, it starts to shoal in middle and late spring (figure not shown). The rapid shoaling of MLD allows winter mixed layer water in the salinity maximum region to be subducted into the main thermocline and carried away by horizontal circulation in the subtropical gyre. The total amount of salinity maximum water that enters the main thermocline per unit horizontal area during the 1-yr period is defined as the water's annual subduction rate, what has been done for South Pacific waters by several earlier studies (e.g., Huang and Qiu 1998; Qu et al. 2008).

Following Huang and Qiu (1998), we calculate the annual subduction rate Sann by tracing water parcels released at the base of winter mixed layer for one year in a Lagrangian framework. That is,
e1
where T represents the time period of integration (one year), t1 and t2 are the end of the first and second winter, respectively, hm is the winter MLD, and is the vertical velocity at the base of mixed layer. Both hm and are values following water parcels. The first term on the right-hand side of Eq. (1), an integrated quantity of along the pathways of water parcels, represents the contribution from vertical pumping, most of which is due to Ekman pumping. The second term represents the contribution from lateral induction. See Huang and Qiu (1998) and references therein for more details.

The calculation of subduction rate is based on the model's long-term winter-mean MLD and monthly velocity fields, as well as the NCEP climatological wind stress used to force the model. To validate the model results, we also calculate the subduction rate using the World Ocean Atlas 2009 (WOA09) combined with the NCEP wind stress climatology. In this calculation, geostrophic velocities along isopycnal surfaces are used (e.g., Qu et al. 2008), assuming a zero-velocity surface at 1500-db depth. Results both from the model and WOA09 are presented in Fig. 4. From this figure, one can see that, despite some quantitative discrepancies, the simulated annual subduction rate shows a pattern that resembles that from the observations in almost all the details. The annual subduction rate is highest (>110 m) near the eastern edge of the SSS maximum, corresponding to the formation of SPTW and eastern subtropical mode water (e.g., Hanawa and Talley 2001; O'Connor et al. 2005). As one can see from the 1-yr Lagrangian trajectories of water particles released at the base of winter mixed layer (Fig. 3), this is the area where most lateral induction takes place as a result of enhanced MLD horizontal gradients. Another high (~100 m) subduction rate is confined in a small region north of New Zealand, which is associated with the formation of western subtropical mode water (e.g., Roemmich and Cornuelle 1992). This spatial distribution is also shown by earlier studies (e.g., Huang and Qiu 1998; Qu et al. 2008). But, with a focus on the subtropical South Pacific, we are able to illustrate more detailed features of the SPTW subduction rate (Fig. 4).

Fig. 4.
Fig. 4.

(a),(b) Subduction rate as a result of (c),(d) vertical pumping and (e),(f) lateral induction in the South Pacific between 10° and 40°S (m yr−1) from (left) WOA09 and (right) ECCO and its components. Black contours in top panels represent annual-mean SSS.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The total simulated subduction volume over the subtropical South Pacific (10°–40°S) is 34.1 Sv (1 Sv ≡ 106 m3 s−1), which compares reasonably well with the estimate (40.5 Sv) from WOA09. Most of the annual subduction rate resulting from vertical pumping occurs in the eastern subtropical gyre (Fig. 4), where its magnitude is typically larger than 50 m yr−1, because of widespread downward Ekman pumping (Fig. 2). The annual subduction rate as a result of lateral induction is relatively weak in much of the region studied. Only in a small region north of New Zealand, it becomes dominant over vertical pumping, as a result of large winter MLD gradients (Fig. 3). Averaged over the SSS maximum region, approximately 87% of the simulated subduction rate is due to vertical pumping, and the contribution from lateral induction is minor (13%). The relative importance between vertical pumping and lateral induction is also estimated using WOA09, combined with the NCEP wind stress climatology, and the result (71% versus 29%) shows a reasonably good agreement with that from the model.

Against winter surface density, the annual subduction rate shows a maximum near 25.0–25.5 kg m−3 (Fig. 5), most of which corresponds with the formation of eastern subtropical mode water (e.g., Tsuchiya and Talley 1996; Hanawa and Talley 2001). Subduction of the salinity maximum water (>35.5 psu), as marked by the black box in Fig. 1, occurs in a density range roughly between 24.0 and 25.0 kg m−3. This density range is slightly lower but broader than that (24.6–25.4 kg m−3) selected by O'Connor et al. (2002, 2005). Based on drifter and tracer data, O'Connor et al. (2002) provided the first estimate of the SPTW subduction rate. According to their calculation, approximately 6–7 Sv of high-salinity (35.6–36.4 psu) water is formed and subducted to the thermocline in the subtropical South Pacific. Our analysis of model results yields a mean SPTW subduction rate estimate of 5.8 Sv within the density range of 24.0–25.0 kg m−3. If we choose the same density range (24.6–25.4 kg m−3) as O'Connor et al. (2002), the estimate (6.3 Sv) is slightly larger. These estimates from the model compare reasonably well with those (7.5 and 5.9 Sv) within the same density ranges (24.0–25.0 and 24.6–25.4 kg m−3) from WOA09. If defined by its highest (>36.2 psu) salinity, the SPTW subduction rate from the model falls slightly to 5.4 Sv, which also shows a good agreement with the estimate (6.0 Sv) from WOA09.

Fig. 5.
Fig. 5.

Subduction volume (Sv) against (a),(c) surface density and (b),(d) salinity from (top) WOA09 and (bottom) ECCO in the subtropical South Pacific (10°–40°S). The light shading represents the subduction of salinity maximum water in the area as indicated by the black box in Fig. 1.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

5. Pathways

Spreading of subducted water property anomalies to remote locations where they can resurface and affect air–sea interaction is one hypothetical teleconnection mechanism for climate variability (e.g., McCreary and Lu 1994; Gu and Philander 1997; McPhaden and Zhang 2002). As an important component of this teleconnection, water property anomalies in the SPTW can be advected equatorward to feed the Equatorial Undercurrent (EUC) and eventually carried eastward to upwell in the eastern equatorial Pacific, as already suggested by several earlier studies (e.g., Lindstrom et al. 1987; Tsuchiya et al. 1989). To demonstrate the SPTW pathways and in particular its interior and western boundary pathways toward the equatorial Pacific, we conduct a passive tracer integration, based on the model's offline circulation. The passive tracer is initially released with a unit value (in arbitrary tracer units per volume, ATU m−3) in the mixed layer of the SSS maximum region marked by the black box in Fig. 1 and then integrated forward in time using velocity and mixing tensors of the simulation averaged at 10-day intervals from September 1993 to August 2006. Additional integrations starting from different initial times (September 1993, September 1994, etc.) were also conducted to test the sensitivity of the results to the initial conditions. Despite some quantitative discrepancies in the passive tracer distribution, the results regarding the SPTW pathways remain the same, suggesting that our results reasonably represent the water's climatological pathways.

As noted by earlier studies (e.g., O'Connor et al. 2005), the intrusion of the SPTW toward the equatorial Pacific is rather fast (Fig. 6). Some of the subducted SPTW approaches the New Guinea coast in about 2 years. From there, some of the water joins the EUC as part of the equatorial current system (details shown in section 6), and some continues to flow northward along the western boundary (see section 5b). Water remaining in the subtropical South Pacific recirculates in the gyre circulation. The SPTW first appears in the eastern equatorial Pacific in about 3 years. A small portion of the water extends to the midlatitudes of the North Pacific, and reaches the Kuroshio extension region in about 7 years, suggesting an interhemisphere exchange of waters. Being part of the basin-scale circulation, the SPTW extends over nearly the entire Pacific basin at the end of the integration (year 13), though its highest concentrations still remain in the subtropical South Pacific (Fig. 6).

Fig. 6.
Fig. 6.

Vertically integrated passive tracer (ATU m−2) distribution at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, in the Pacific. The black box indicates the SSS maximum region.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

a. Water mass transformation

Passive tracer distribution in temperature and salinity space further illustrates how the subducted SPTW is transformed by mixing (Fig. 7). When passive tracer is released in the subtropical South Pacific (year 0), most of the tracer-tagged water is confined between σθ = 24.0 and 25.5 kg m−3, with its temperature and salinity lying at 20°–25°C and 35.5–36.5 psu, respectively. As the water is traced forward in time, it gets mixed with its surroundings. By the end of year 2, both temperature and salinity of the water have changed, leading to a slight increase in its density ranging from about 24.5 to 25.6 kg m−3. As time progresses, the high tracer concentration (>1.0 × 1013 ATU) becomes narrowly confined, and some SPTW gradually loses its high-salinity signature (Fig. 7). By the end of integration (year 13), the water's temperature, salinity, and density have reached 15°–20°C, 35.2–35.8 psu, and 25.3–26.0 kg m−3, respectively, clearly demonstrating the effects of mixing on the water mass transformation.

Fig. 7.
Fig. 7.

Distribution of passive tracer concentration (1013 ATU) in temperature and salinity space at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, in the Pacific.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

To further illustrate when and where the water property changes take place, we calculate the tracer-weighted potential density and salinity of the SPTW and their standard deviations from the mean values (Fig. 8). It is clear that the tracer-weighted density of the water changes steadily from its initial value of about 24.4 to nearly 25.0 kg m−3 at the end of integration (Fig. 8a). Its associated standard deviation is relatively small (~0.6 kg m−3) in the first 5 years. The value then increases gradually to more than 1.0 kg m−3 in the later period of the integration. The relatively small standard deviation in the first 5 years is consistent with the water's narrow horizontal extent (Fig. 6). When some portion of the water reaches the western boundary and penetrates into higher latitudes of the basin along the western boundary, it is transformed by strong surface cooling and mixing, and the wide spreading of the water results in a large density range or standard deviation (Fig. 8a). The mean spatial distribution of tracer-weighted density shows that the SPTW spreads over nearly the entire basin of the Pacific and some portion of it can reach as high density as 27.0 kg m−3 in the southern subtropical South Pacific (Fig. 8b). This probably explains why most of the SPTW at later years is denser than its initial value. On spatial average, the tracer-weighted salinity of the SPTW decreases from about 36.1 to 35.3 psu during the period of integration (Fig. 8c), but its standard deviation is always smaller than 0.3 psu, suggesting that the large standard deviation in density (Fig. 8a) is mostly due to temperature changes. On temporal average, the water that reaches tropical and subpolar regions loses salinity to its surroundings (Fig. 8d), contributing directly to the freshening of the SPTW.

Fig. 8.
Fig. 8.

(a) Time evolution and (b) spatial distribution of tracer-weighted density of the water. (c),(d) As in (a),(b), but for tracer-weighted salinity. Here, (a),(c) are the spatial averages (solid line) with std dev (gray) over the entire domain, and (b),(d) are the temporal averages (shading) over the entire period.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

b. Equatorward intrusion

The SPTW enters the equatorial region both through the interior and western boundary pathways (e.g., Fine et al. 1994). Johnson and McPhaden (1999) showed a net thermocline equatorward flow of about 14 Sv across 8°S to the east of 165°E. The passive tracer integration confirms the existence of the two pathways, through which the SPTW can be conveyed into the equatorial region. To quantify the relative importance of the two pathways, we conduct another passive tracer integration using the offline circulation from September 1993 to August 2006. The only difference between this passive tracer integration and the one described above is that water entering the equatorial region is trapped. Specifically, when a tracer-tagged water parcel crosses 8°S, it is considered to have entered the equatorial region, and no tracking of this water parcel will be further conducted. Computationally, the tracer in the region to the north of 8°S is tabulated and reset to zero after each model time step. The time-integrated tracer content represents the amount of the initially tagged SPTW that has entered the equatorial region across 8°S.

The results of this passive tracer integration show that the tracer-tagged SPTW starts to appear in the equatorial region in about 1 year (Fig. 9). The water does so through its interior pathway, which allows the water to reach the eastern equatorial Pacific in about 3 years. The western boundary pathway becomes evident in about 5 years, and gradually gets enhanced afterward, consistent with the tracer's horizontal extent (Fig. 6). At the end (year 13) of integration, the accumulated passive tracer transport per unit area within the western boundary can be as large as 32.2 × 105 ATU m−2 at 8°S (Fig. 9), which is greater than the value (5.9 × 105 ATU m−2) averaged between 150° and 170°W by at least a factor of 5. In space, the interior pathway is widespread, with its core expending from about 150° to 170°W. Most of the water entering the equatorial region via the interior pathway lies between the 24.0 and 25.0 kg m−3 density surfaces (Fig. 9). In contrast, the western boundary pathway is narrowly confined to about 150°E in longitude, coinciding with the northward-flowing New Guinea Coastal Undercurrent (NGCUC; Lindstrom et al. 1987). This western boundary pathway, however, has a much broader vertical extension than the interior pathway. By year 13, a significant portion of the water entering the equatorial region via the western boundary pathway extends below 26.5 kg m−3 or deeper, presumably because of enhanced diapycnal mixing there.

Fig. 9.
Fig. 9.

Spatial distribution of passive tracer (105 ATU m−2) that has entered the equatorial region (shading) with mean isopycnals (kg m−3) superimposed (contours) at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, through 8°S. The interior and western boundary pathways are indicated by the high concentrations centered at 160°W and 150°E, respectively.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The time evolution of passive tracer transport through the two pathways is shown in Fig. 10. From this figure, one can see that most of the tracer-tagged SPTW enters the equatorial region through its interior pathway in the first half of the integration (Fig. 10a). As noted above, the water starts to penetrate toward the equator through its western boundary pathway in about 5 years. Then, the relative importance of western boundary versus interior pathway increases with time (Fig. 10a). By the end of integration (year 13) (Fig. 10b), approximately 42% of the total subducted SPTW has crossed 8°S into the equatorial region, of which about 61% does so through the interior pathway and 39% through the western boundary pathway. This result confirms the importance of interior pathway in the South Pacific (e.g., Johnson and McPhaden 1999). Given the trend of relative importance between the western boundary and interior pathway shown in Fig. 10a, one can easily get an impression that the two pathways are of similar importance.

Fig. 10.
Fig. 10.

(a) Time evolution of passive tracer (heavy solid black) that has entered the equatorial region through 8°S against year and (b) its distribution against longitude at year 13. The light solid and dashed lines in (a) represent components through the western boundary (west of 160°E) and interior (east of 160°E) pathways. The units on the vertical axis are 1014 ATU in (a) and 108 ATU m−1 in (b).

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

6. Resurfacing

It has been known for decades that the SPTW contributes to the EUC in the equatorial Pacific (e.g., Lindstrom et al. 1987; Tsuchiya et al. 1989). To demonstrate this in the model, Fig. 11 shows the distribution of passive tracer at 160°E between 5°S and 5°N. From this figure, one can see that the tracer-tagged SPTW starts to appear in the western equatorial Pacific in about 2 years (Fig. 11c). As already shown in Fig. 9c, most of this water enters the western equatorial Pacific through the interior pathway. The signature of western boundary pathway at this longitude becomes markedly evident in about 7 years (Figs. 9f and 11f) and reaches its maximum strength around the end of year 9 (Figs. 9g and 11g). Water entering the western equatorial Pacific through the western boundary pathway directly contributes to the EUC, as indicated by a high concentration core of passive tracer coinciding with the EUC at year 9 and afterward (Figs. 11g–i). Contribution from the interior pathway is relatively weak during the later period of integration (Fig. 10a). It is also interesting to note that the SPTW mostly feeds the southern part of the EUC in the western equatorial Pacific (e.g., Fig. 11i). This result seems to be consistent with the recent modeling study of Grenier et al. (2011). What feeds the northern part of the EUC is another interesting issue to address, and we will leave it for future studies.

Fig. 11.
Fig. 11.

Distribution of passive tracer (10−3 ATU m−2) at 160°E between 5°S and 5°N (shading) at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, superimposed with the zonal velocity in the upper 300 m. Positive values represent eastward flow.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The SPTW, upon reaching the equatorial Pacific, may directly affect the properties of water upwelled at the equator (e.g., Lindstrom et al. 1987; Tsuchiya et al. 1989). To further demonstrate where the SPTW reaches the surface mixed layer, we conduct an additional passive tracer integration, which is essentially the same as the one described in section 5, except for the surface mixed layer in the equatorial Pacific (5°S–5°N). When a tracer-tagged water parcel reaches the surface mixed layer in the equatorial region, no tracking of this water parcel will be further conducted. As such, the time-integrated tracer content represents the amount of initially tagged SPTW that has entered the surface mixed layer in the equatorial Pacific.

As one would expect, some of the tracer-tagged SPTW reaches the surface mixed layer in the equatorial Pacific within a year (Fig. 12), through its interior pathway in the central part of the basin. As the western boundary pathway starts to become significant, high tracer concentrations appear in the central equatorial Pacific, roughly, between 150° and 170°W around the end of year 5. As discussed above (Fig. 11), a significant portion of the SPTW reaches the equatorial region via the NGCUC (e.g., Lindstrom et al. 1987), and from there it flows into the central equatorial Pacific in the EUC. At the end of integration (year 13), about 2.3 × 1014 ATU or 30% of the initially released passive tracer has entered the surface mixed layer in the equatorial Pacific, accounting for more than 70% of the total SPTW that has reached the equatorial region through 8°S. Considering that some of the passive tracer lying below the surface mixed layer at the end of year 13 may resurface at a later time, this result implies that the overwhelming majority of the SPTW entering the equatorial Pacific will eventually reach the surface mixed layer there.

Fig. 12.
Fig. 12.

Distribution of passive tracer (ATU m−2) that has entered the mixed layer in equatorial Pacific between 5°S and 5°N at years (a) 0, (b) 1, (c) 2, (d) 3, (e) 5, (f) 7, (g) 9, (h) 11, and (i) 13, with annual-mean SSS (white contours) superimposed.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

The total amount of resurfaced passive tracer in the equatorial region can be further partitioned into two parts (Fig. 13). One is in the western equatorial Pacific (west of the international date line) and the other in the central and eastern equatorial Pacific (east of the international date line). It appears that most of the tracer-tagged SPTW enters the surface mixed layer in the central equatorial Pacific. As water entering the surface mixed layer in the western equatorial Pacific is supplied primarily by the western boundary current, no water does so in the first 4 years. Afterward, the amount of tracer entering the surface mixed layer in the western equatorial Pacific increases slowly, but it can only account for about 9% of the total resurfaced passive tracer at the end of integration (year 13). Note that the resurfacing of the SPTW is asymmetric, greater in the Southern Hemisphere than that in the Northern Hemisphere. For the central part of the basin, this asymmetry is probably related to the interior pathway of the South Pacific, while for the eastern part of the basin, this asymmetry reflects significant differences of ocean circulation between the two hemispheres [e.g., Kessler (2006), and references therein].

Fig. 13.
Fig. 13.

Time evolution of passive tracer (1014 ATU) that has entered the mixed layer in the equatorial Pacific (heavy solid) and its component east (dashed) and west (solid) of the date line.

Citation: Journal of Physical Oceanography 43, 8; 10.1175/JPO-D-12-0180.1

7. Summary and discussion

Based on the results from a simulated passive tracer, this study has provided a comprehensive description of the subduction of the SPTW and its equatorward pathways in the South Pacific. It has been shown that approximately 5.8 Sv of the SPTW is formed in the density range between 24.0 and 25.0 kg m−3, of which about 87% is due to vertical pumping and 13% due to lateral induction. If defined by its high-salinity (>36.2 psu) signature, the subduction rate of the SPTW falls slightly to 5.4 Sv. All these estimates show a reasonable agreement with those derived from WOA09 and other observations (e.g., O'Connor et al. 2002, 2005). Note that the dominance of vertical pumping over lateral induction presented in this study differs from what was suggested by O'Connor et al. (2002). Based on the World Ocean Circulation Experiment drifter array and chlorofluorocarbon-12 data during the periods 1988–92 and 1992–96, O'Connor et al. (2002) found that vertical pumping greatly exceeded lateral induction during the second but not the first period. This discrepancy between the two studies could result from different representation of the data. The details need to be investigated further by research.

Once subducted, the SPTW first spreads in the subtropical South Pacific. As time goes by, the tracer-tagged high-salinity water can reach almost everywhere of the Pacific basin, but most of it still remains in the southern subtropical gyre. A significant portion (42%) of the water makes its way into the equatorial region, either through its interior or western boundary pathway. With a significant time delay, the western boundary pathway is relatively weak at the beginning, but becomes equally important as the interior pathway by the end of integration. Another important difference between the two pathways is that the western boundary pathway seems to extend much deeper than the interior pathway, suggesting enhanced diapycnal mixing there.

An important advantage of this study is the use of passive tracer that involves both advection and mixing of the SPTW. Different from those earlier studies tracking the SPTW along isopycnal surfaces (e.g., Qu et al. 1999; O'Connor et al. 2002), this study shows that the SPTW changes its density both in time and space. Because of the presence of mixing, the tracer-weighted density of the SPTW steadily increases from an initial value of about 24.4 to nearly 25.0 kg m−3 at the end of integration. This density change reflects the strong influence of gyre circulation, which allows the SPTW to spread over nearly the entire Pacific basin. Though a significant portion of the tracer-tagged water becomes the relatively warm and freshwater that resurfaces at the equator, most of the SPTW remains in the subtropical South Pacific within the depth range of the thermocline or below. The mixing of the SPTW with the relatively cold and dense water at the high latitudes leads to an increase in its tracer-weighted density.

Approximately 30% of the tracer-tagged SPTW resurfaces in the equatorial Pacific after 13 years of integration. The tracer-tagged SPTW enters the surface mixed layer over a broad longitude band from the international date line to the eastern boundary of the Pacific, with its highest concentration roughly coinciding with the western edge of the cold tongue near 140°–160°W. The resurfacing of the relatively cold, saline SPTW may alter the surface stratification and sea surface temperature along the equator. We have already known that the eastern Pacific cold tongue is often inaccurately simulated in ocean models, with the sea surface temperature always being too cold and/or extending too far from the eastern boundary. The causes of these model biases are likely related to inaccurate representation of the ocean's communication between the thermocline and surface mixed layer, namely, the models' parameterizations of entrainment processes. The resurfacing of the SPTW may represent one of the entrainment processes that need to be investigated further by research.

The resurfacing of the SPTW may also affect the western Pacific warm pool. The western Pacific warm pool, characterized by sea surface temperatures warmer than 28°–29°C, lies around the equator to the west of the international date line and is believed to be an important component of the world's climate system (e.g., Palmer and Mansfield 1984; McPhaden et al. 2006). Associated with the eastern edge of the warm pool is a zonal SSS front separating the fresh western Pacific water from relatively saline central Pacific water (e.g., Delcroix and McPhaden 2002; Maes 2008). The SSS front has been shown to change its position from year to year along the equator, with a strong ENSO signal in its variability (e.g., Singh et al. 2011). Earlier studies have related the zonal displacement of this SSS front to large-scale processes both in the atmosphere and ocean (e.g., Kessler and Taft 1987; McPhaden and Picaut 1990; Picaut et al. 1996). Here, we emphasize the potential importance of subsurface ocean processes. The resurfacing of the SPTW in the central equatorial Pacific may alter the thermocline structure and surface stratification and directly contribute to the formation of the SSS front and its zonal displacement. The details will be investigated by a separate study.

Acknowledgments

This research was supported by NSF through Grants OCE11-30050 and OCE10-29793. T. Qu was also supported by NSF through Grant OCE10-29704 and by NASA as part of the Aquarius Science Team investigation through Grant NNX12AG02G. The authors are grateful to I. Fukumori for long-term collaboration, to N. Schneider for useful discussion on the topic, and to two anonymous reviewers for valuable comments on an earlier version of the manuscript.

REFERENCES

  • Adler, R. F., and Coauthors, 2003: The Version-2 Global Precipitation Climatology Project (GPCP) monthly precipitation analysis (1979–present). J. Hydrometeor., 4, 11471167.

    • Search Google Scholar
    • Export Citation
  • Barnier, B., , L. Siefridt, , and P. Marchesiello, 1995: Thermal forcing for a global ocean circulation model using a three-year climatology of ECMWF analysis. J. Mar. Syst., 6, 363380.

    • Search Google Scholar
    • Export Citation
  • Delcroix, T., , and M. J. McPhaden, 2002: Interannual sea surface salinity and temperature changes in the western Pacific warm pool during 1992–2000. J. Geophys. Res., 107, 8002, doi:10.1029/2001JC000862.

    • Search Google Scholar
    • Export Citation
  • Fine, R. A., , R. Lukas, , F. M. Bingham, , M. J. Warner, , and R. H. Gammon, 1994: The western equatorial Pacific: A water mass crossroads. J. Geophys. Res., 99 (C12), 25 06325 080.

    • Search Google Scholar
    • Export Citation
  • Fine, R. A., , K. A. Maillet, , K. F. Sullivan, , and D. Willey, 2001: Circulation and ventilation flux of the Pacific Ocean. J. Geophys. Res., 106 (C10), 22 15922 178.

    • Search Google Scholar
    • Export Citation
  • Fukumori, I., , T. Lee, , B. Cheng, , and D. Menemnlis, 2004: The origin, pathway, and destination of Niño-3 water estimated by a simulated passive tracer and its adjoint. J. Phys. Oceanogr., 34, 582604.

    • Search Google Scholar
    • Export Citation
  • Gao, S., , T. Qu, , and I. Fukumori, 2011: Effects of mixing on the subduction of South Pacific waters. Dyn. Atmos. Oceans, 51, 4554.

  • Gent, P. R., , and J. C. McWilliams, 1990: Isopycnal mixing in ocean circulation models. J. Phys. Oceanogr., 20, 150155.

  • Grenier, M., , S. Cravatte, , B. Blanke, , C. Menkes, , A. Koch-Larrouy, , F. Durand, , A. Melet, , and C. Jeandel, 2011: From the western boundary currents to the Pacific Equatorial Undercurrent: Modeled pathways and water mass evolutions. J. Geophys. Res., 116, C12044, doi:10.1029/2011JC007477.

    • Search Google Scholar
    • Export Citation
  • Gu, D., , and S. G. H. Philander, 1997: Interdecadal climate fluctuations that depend on exchange between the tropics and extratropics. Science, 275, 805807.

    • Search Google Scholar
    • Export Citation
  • Hanawa, K., , and L. D. Talley, 2001: Mode waters. Ocean Circulation and Climate, G. Sieldler, J. Church, and J. Gould, Eds., Academic Press, 373–386.

  • Huang, R. X., , and B. Qiu, 1998: The structure of the wind-driven circulation in the subtropical South Pacific Ocean. J. Phys. Oceanogr., 28, 11731186.

    • Search Google Scholar
    • Export Citation
  • Johnson, G. C., , and M. J. McPhaden, 1999: Interior pycnocline flow from the subtropical to the equatorial Pacific Ocean. J. Phys. Oceanogr.,29, 3073–3089.

  • Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77, 437371.

  • Kessler, W. S., 2006: The circulation of the eastern tropical Pacific: A review. Prog. Oceanogr., 69, 181217.

  • Kessler, W. S., , and B. A. Taft, 1987: Dynamic heights and zonal geostrophic transports in the central tropical Pacific during 1979–1984. J. Phys. Oceanogr., 17, 97122.

    • Search Google Scholar
    • Export Citation
  • Large, W. G., , J. C. McWilliams, , and S. C. Doney, 1994: Oceanic vertical mixing: A review and a model with a nonlocal boundary-layer parameterization. Rev. Geophys., 32, 363403.

    • Search Google Scholar
    • Export Citation
  • Lindstrom, E., , R. Lukas, , R. Fine, , E. Firing, , S. J. Godfrey, , G. Meyers, , and M. Tsuchiya, 1987: The western equatorial Pacific Ocean circulation study. Nature, 330, 533537.

    • Search Google Scholar
    • Export Citation
  • Lukas, R., , and E. Lindstrom, 1991: The mixed layer of the western equatorial Pacific Ocean. J. Geophys. Res., 96, 33433358.

  • Luo, J.-J., , and T. Yamagata, 2001: Long-term El Niño–Southern Oscillation (ENSO)-like variation with special emphasis on the South Pacific. J. Geophys. Res., 106 (C10), 22 21122 227.

    • Search Google Scholar
    • Export Citation
  • Maes, C., 2008: On the ocean salinity stratification observed at the eastern edge of the equatorial Pacific warm pool. J. Geophys. Res., 113, C03027, doi:10.1029/2007JC004297.

    • Search Google Scholar
    • Export Citation
  • Marshall, J. C., , A. Adcroft, , C. Hill, , L. Perelman, , and C. Heisey, 1997: A finite-volume, incompressible Navier Stokes model for studies of the ocean on parallel computers. J. Geophys. Res., 102 (C3), 57535766.

    • Search Google Scholar
    • Export Citation
  • McCreary, J. P., , and P. Lu, 1994: On the interaction between the subtropical and the equatorial oceans: The subtropical cell. J. Phys. Oceanogr., 24, 466497.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , and J. Picaut, 1990: El Niño–Southern Oscillation displacement of the western equatorial Pacific warm pool. Science, 250, 13851388.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , and D. Zhang, 2002: Slowdown of the meridional overturning circulation in the upper Pacific Ocean. Nature, 415, 603608.

    • Search Google Scholar
    • Export Citation
  • McPhaden, M. J., , S. E. Zebiak, , and M. H. Glantz, 2006: ENSO as an integrating concept in earth. Science, 314, 17401745.

  • O'Connor, B. M., , R. A. Fine, , K. A. Maillet, , and D. B. Olson, 2002: Formation rates of subtropical underwater in the Pacific Ocean. Deep-Sea Res., 49, 15711590.

    • Search Google Scholar
    • Export Citation
  • O'Connor, B. M., , R. A. Fine, , and D. B. Olson, 2005: A global comparison of subtropical underwater formation rate. Deep-Sea Res., 52, 15691590.

    • Search Google Scholar
    • Export Citation
  • Palmer, T. N., , and D. A. Mansfield, 1984: Response of two atmospheric general circulation models to sea surface temperature anomalies in the tropical east and west Pacific. Nature, 310, 483.

    • Search Google Scholar
    • Export Citation
  • Picaut, J., , M. Loualalen, , C. Menkes, , T. Delcroix, , and M. J. McPhaden, 1996: Mechanism of the zonal displacement of the Pacific warm pool: Implications for ENSO. Science, 274, 14861489.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , H. Mitsudera, , and T. Yamagata, 1999: A climatology of the circulation and water mass distribution near the Philippine coast. J. Phys. Oceanogr., 29, 14881505.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , S. Gao, , I. Fukumori, , R. A. Fine, , and E. J. Lindstrom, 2008: Subduction of South Pacific waters. Geophys. Res. Lett., 35, L02610, doi:10.1029/2007GL032605.

    • Search Google Scholar
    • Export Citation
  • Qu, T., , S. Gao, , I. Fukumori, , R. A. Fine, , and E. J. Lindstrom, 2009: Origin and pathway of equatorial 13°C Water in the Pacific identified by a simulated passive tracer and its adjoint. J. Phys. Oceanogr., 39, 18361853.

    • Search Google Scholar
    • Export Citation
  • Redi, M., 1982: Oceanic isopycnal mixing by coordinate rotation. J. Phys. Oceanogr., 12, 11541158.

  • Reynolds, R. W., , and T. M. Smith, 1994: Improved global sea surface temperature analyses using optimum interpolation. J. Climate, 7, 929948.

    • Search Google Scholar
    • Export Citation
  • Rodgers, K. B., , B. Blanke, , G. Mades, , O. Aumont, , P. Ciais, , and J.-C. Dutay, 2003: Extratropical sources of equatorial Pacific upwelling in an OGCM. Geophys. Res. Lett., 30, 1084, doi:10.1029/2002GL016003.

    • Search Google Scholar
    • Export Citation
  • Roemmich, D., , and B. Cornuelle, 1992: The subtropical mode waters of the South Pacific Ocean. J. Phys. Oceanogr., 22, 11781187.

  • Singh, A., , T. Delcroix, , and S. Cravatte, 2011: Contrasting the flavors of El Niño–Southern Oscillation using sea surface salinity observations. J. Geophys. Res., 116, C06016, doi:10.1029/2010JC006862.

    • Search Google Scholar
    • Export Citation
  • Slutz, R. J., , S. J. Lubker, , J. D. Hiscox, , S. D. Woodruff, , R. L. Jenne, , D. H. Joseph, , P. M. Steurer, , and J. D. Elms, 1985: Comprehensive ocean–atmosphere data set; Release 1. NOAA Environmental Research Laboratories, Climate Research Program Rep. NTIS PB86105723, 268 pp.

  • Tsuchiya, M., , and L. D. Talley, 1996: Water property distributions along an eastern Pacific hydrographic section at 135°W. J. Mar. Res., 54, 541564.

    • Search Google Scholar
    • Export Citation
  • Tsuchiya, M., , R. Lukas, , R. A. Fine, , E. Firing, , and E. Lindstrom, 1989: Source waters of the Pacific Equatorial Undercurrent. Prog. Oceanogr.,23, 101–147.

  • White, W. B., , and D. R. Cayan, 1998: Quasi-periodicity and global symmetries in interdecadal upper ocean temperature variability. J. Geophys. Res., 103 (C10), 21 33521 354.

    • Search Google Scholar
    • Export Citation
  • Yu, L., 2007: Global variations in oceanic evaporation (1958–2005): The role of the changing wind speed. J. Climate, 20, 53765390.

  • Zhang, R.-H., , L. M. Rothstein, , and A. J. Busalacchi, 1998: Origin of upper-ocean warming and El Niño change on decadal scales in the tropical Pacific Ocean. Nature, 391, 879883.

    • Search Google Scholar
    • Export Citation
Save