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Observed subseasonal variability of oceanic barrier and compensated
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layers
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Hailong Liu, Semyon A. Grodsky, and James A. Carton
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Submitted to Journal of Climate
December 10, 2008
Department of Atmospheric and Oceanic Science
University of Maryland, College Park, MD 20742
Corresponding author: Semyon Grodsky (senya@atmos.umd.edu)
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Abstract
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1960-2007 is explored for evidence of subseasonal variability and the causes of these
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components of mixed layer variability. The strongest variability from monthly climatology is
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found in the subpolar latitudes of the North Atlantic in winter, where it reaches 100m. A
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temperature-salinity compensated layer in the eastern North Atlantic varies interannually in
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conjunction with a North Atlantic Oscillation-like pattern of anomalous sea level pressure. Thus
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thickening/thinning of this layer occurs in conjunction with the meteorological changes
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associated with strengthening/weakening of the Azores sea level high pressure center.
A monthly gridded analysis of barrier layer/compensated layer thickness covering the period
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In the winter subpolar North Pacific a salinity stratified barrier layer exists. While further south
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along the Kuroshio extension a compensated layer exists. The thickness of both of these layers
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varies from year to year by up to 60m. A significant component of this variability is a long-term
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trend of shrinkage of the barrier layer in the subpolar gyre and growth of the compensated layer
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along the Kuroshio Extension. This trend is shown to result from trends in meteorological
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variables. In the subpolar gyre the barrier layer has shrunk in response to a strengthening of the
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Aleutian low pressure system, the resulting strengthening of dry northerly winds, and a decrease
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of precipitation. In the Kuroshio Extension the compensated layer thickness has increased in
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response to this pressure system strengthening and related amplification of the midlatitude
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westerly winds, the resulting increase of net surface heat loss, and its effect on the temperature
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and salinity of the upper ocean water masses.
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The tropical Pacific and Atlantic both have barrier layers with variability that is less than 20m
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but is still important in influencing mixed layer properties. In the tropical Pacific west of 160oE,
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the barrier layer thickness changes by approximately 5m in response to 10 unit change in the
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South Oscillation Index. It thickens during La Ninas as a result of the abundant rainfall and thins
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during dry El Ninos. But east of 160oE variations in zonal advection of salt becomes as important
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as local rainfall in influencing the thickness of the barrier layer.
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1. Introduction
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The ocean mixed layer is a nearsurface layer of fluid with quasi-uniform properties such as
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temperature, salinity, and density. The thickness of this mixed layer and its time rate of change
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both strongly influence the ocean’s role in air-sea interaction. However, the thickness of the
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nearsurface layer of quasi-uniform temperature, MLT, may differ from the thickness of the
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nearsurface layer of quasi-uniform density, MLD. MLT may be deeper than MLD when positive
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salinity stratification forms a barrier layer (BL=MLT-MLD) isolating the shallower and deeper
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levels of the mixed layer as was originally found in the western equatorial Pacific (Lukas and
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Lindstrom, 1991). Elsewhere MLT may be shallower than MLD when negative salinity
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stratification compensates for positive temperature stratification (or the reverse situation) to form
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a Compensated Layer (CL=MLD-MLT) (Stommel and Fedorov, 1967; Weller and
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Plueddemann, 1996). Changes in the seasonal thicknesses of BLs and CLs from one year to the
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next may cause corresponding changes in the role of the mixed layer in air-sea interaction by
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altering the effective depth of the mixed layer or the temperature of water at the mixed layer base
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(e.g., Ando and McPhaden, 1997). Here we examine the global historical profile data set
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covering the period 1960-2007 for corresponding year-to-year changes in the BL/CL thickness
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distribution.
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Four studies; Sprintall and Tomczak (1992), Tomczak and Godfrey (1994), de Boyer Montegut et
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al. (2007), and Mignot et al. (2007); have provided an observational description of the seasonal
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cycle of BL/CL distribution over much of the global ocean. BLs are a persistent feature of the
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tropics as well as high latitudes during winter. In high latitudes BLs occur where freshening in
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the near-surface is produced by excess precipitation over evaporation, river discharge, or ice
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melting (de Boyer Montegut et al., 2007) and may be most evident in regions where upward
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Ekman pumping ( wEk ) acts against the effects of vertical mixing such as occurs in the subpolar
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gyre (Kara et al., 2000). BLs are also observed equatorward of the salty subtropical gyres where
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subduction leads to a shallow subsurface salinity maximum (Sato et al., 2006). The North Pacific
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provides an example of this, the result of southward subduction of salty North Pacific
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Subtropical Mode Water below fresher tropical surface water (as originally suggested by
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Sprintall and Tomczak, 1992). BLs also occur in the Southern Ocean south of the Polar Front as
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a result of near surface freshening and weak thermal stratification (e.g. de Boyer Montegut et al.,
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2004).
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BLs also occur seasonally in the tropics in regions of high rainfall and river discharge such as the
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Arabian Sea and Bay of Bengal, where layers as thick as 20-60m have been observed (Thadathil
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et al., 2008). Similarly, BLs occur in the western Equatorial Pacific under the high precipitation
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regions of the Intertropical Convergence Zone and South Pacific Convergence Zone (Lukas and
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Lindstrom, 1991) and in the western tropical Atlantic (Pailler et al., 1999; Ffield, 2007).
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Much less is known about subseasonal variations in BL/CL. In their examination of mooring
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time series Ando and McPhaden (1997) show that BLs do have interannual variability in the
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central and eastern Pacific, and conclude that the major driver is precipitation variability
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associated with El Nino. At 0oN 140oW, for example, the BL thickness increased from 10m to
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40m in response to the enhanced rains of the 1982-3 El Nino. In contrast, Mignot et al. (2007)
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suggest that changes in zonal advection of high salinity water in response to El Nino winds are
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important in regulating BLs in midbasin. In the west an observational study by Cronin and
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McPhaden (2002) is able to document the response of the mixed layer to intense westerly wind
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bursts and accompanying precipitation and show how these lead to both the formation and
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erosion of BLs.
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CLs in contrast may result from excess evaporation over precipitation, such as occurs in the
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subtropical gyres, or by differential advection where it leads to cooler fresher surface water
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overlying warmer saltier subsurface water (Yeager and Large, 2007). de Boyer Montegut et al.
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(2004) summarize several additional possible mechanisms of CL formation, such as subduction-
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induced advection, Ekman transport, slantwise convection and density adjustment. CLs are most
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prominent in the eastern subpolar North Atlantic and in the Southern Ocean (de Boyer Montegut
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et al., 2007). In the eastern North Atlantic a CL is formed by transport of the warm and salty
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North Atlantic Current above fresher colder subpolar water. Further east the North Atlantic
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Current splits into a northern branch comprising the Norwegian and Irminger Currents, and the
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southward Canary Current, all of which also develop CLs.
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In this study we build on previous observational examinations of the seasonal cycle of BL/CL
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development, to explore year-to-year variability. This study is made possible by the extensive 7.9
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million hydrographic profile data set contained in the World Ocean Database 2005 (Boyer et al.,
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2006) supplemented by an additional 0.4 million profiles collected as part of the Argo observing
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program. We focus our attention primarily on the Northern Hemisphere because of its higher
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concentration of historical observations.
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2. Data and methods
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This study is based on the combined set of temperature and salinity vertical profiles archived in
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the World Ocean Database 2005 for the period 1960-2004 and Argo floats from 1997 to 2007.
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The first dataset was adopted from the study of variability of the oceanic mixed layer by Carton
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et al. (2008) where issues of data quality control and processing are detailed.
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Mixed layer depth is defined here following Carton et al. (2008) which in turn combines the
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approaches of Kara et al. (2000) and de Boyer Montegut et al. (2004). The mixed layer depth is
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the depth at which the change in temperature or density from its value at the reference depth of
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10m exceeds a specified value (for temperature: T | 0.20 C | ). This reference depth is
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sufficiently deep to avoid aliasing by the diurnal signal, but shallow enough to give a reasonable
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approximation of monthly SST. The value of T | 0.20 C | is chosen following de Boyer
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Montegut et al. (2004) as a compromise between the need to account for the accuracy of mixed
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layer depth retrievals and the need to avoid sensitivity of the results to measurement error. The
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absolute temperature difference instead of the negative temperature difference is used following
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Kara et al. (2000) in order to accommodate for temperature inversions that are widespread at
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high latitudes. The specified change in density used to define the density-based mixed layer
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depth follows the variable density criterion (e.g. Monterey and Levitus, 1997) to be locally
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compatible with the specified temperature value, (i.e.   ( / T )  0.20 C ). In this study the
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thickness (or width) of either a barrier layer or compensated layer is defined as a difference of
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isothermal mixed layer depth and isopycnal mixed layer depth, MLT-MLD. As a result of these
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definitions a positive MLT-MLD difference indicates the presence of a BL while a negative
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difference indicates the presence of a CL. We compute BL/CL thicknesses for each profile which
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are then passed through a subjective quality control to eliminate outliers and binned into 2o×2o×1
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month bins without attempt to fill in empty bins.
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In order to quantify the relative impact of temperature and salinity stratification within BLs and
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CLs we introduce a bulk Turner Angle, defined following Yeager and Large (2007) as:
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Tub  tan 1 [(T  S ) /(T  S )] , where    1 / T and    1 / S are the
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expansion coefficients for temperature, T , and salinity, S . In this study the changes in
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temperature and salinity T and S are defined between the top, zt  min( MLT , MLD) , and
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the bottom, zb  max( MLT , MLD) , of either a BL or CL based on analysis of individual vertical
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profiles. The bulk Turner angle is then evaluated from spatially binned values of T and S .
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Fig. 1 shows the correspondence between values of T , S , Tub , and the presence of BLs and
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CLs. Angles | Tub | <45o correspond to BLs stabilized by both, temperature and salinity ( T  0 ,
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S <0). A perfect BL stabilized by salinity and homogeneous in T corresponds to Tub =-45o,
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while Tub =45o corresponds to a pure thermal stratification. Angles greater than 45o correspond
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to the most frequently occurring CLs where positive temperature stratification compensates for
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negative salinity stratification (mixed layer is saltier than the thermocline). Less frequently
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occurring CLs are observed at high latitudes below cool and fresh mixed layers (-90o< Tub < -
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72o). The transition point is associated with the density ratio R  T / S  0.5 or
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Tub  tan 1 ( 3)  72 0 . Thus bulk Turner angle provides an alternate way of displaying BL/CL
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distribution.
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We explore the role that surface forcing plays in regulating mixed layer properties through
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comparison of the BL/CL distribution to fluxes from the NCEP-NCAR reanalysis of Kalnay et
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al. (1996). Satellite QuikSCAT scatterometer winds (see Liu, 2002), which begin in mid-1999,
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are used to characterize the finer scale spatial patterns of wEk . To better characterize precipitation
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in the tropics, we also examine the Climate Prediction Center Merged Analysis of Precipitation
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(CMAP) of Xie and Arkin (1997), which covers the period 1979 -present.
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Limitations in the historical observational coverage limit our ability to explore the mechanisms
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giving rise to subseasonal BL/CL variability. In order to address these mechanisms we repeat our
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analysis on output from a 140-year long control simulation by the NOAA/Geophysical Fluid
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Dynamics Laboratory CM2.1 global coupled climate model forced with natural and
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anthroprogenic forcings beginning in year 1860 (after a 220 year spin-up period, Delworth et al.,
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2006). This coupled simulation maintains full freshwater and heat flux balances at the air-sea
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interfaces. The ocean model component is based on Modular Ocean Model V4 numerics with
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10m vertical resolution in the upper 220 m and including vertical mixing with mixing rates
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determined based on the K-profile parameterization (KPP) mixed layer scheme of Large et al.
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(1994). Comparison of the simulated and observed mixed layer shows encouraging
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correspondence in both the geographic position and seasonal timing of BL/CL variability.
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3. Results
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We begin with a brief review of features of the mean and seasonal distribution.
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3.1. Time mean and seasonal patterns Global seasonal patterns of BL/CL display many features
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revealed by previous analyses (de Boyer Montegut et al., 2007). Throughout the year there are
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persistent BLs in the tropics in areas of high precipitation (Figs. 2a, 2b). BLs are thick under the
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Intertropical Convergence Zone and the South Pacific Convergence Zone. BLs are generally
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thickest in the western side of the tropical Pacific and Atlantic Oceans reflecting higher levels of
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rain as well as (in the case of the Atlantic) Amazon river discharge. BLs are thickest on the
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eastern side of the tropical Indian Ocean due to the presence of the Java and Sumatra high
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precipitation area (Qu and Meyers, 2005). Rainfall in the southern Intertropical Convergence
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Zone (Grodsky and Carton, 2003) may contribute to freshening of the mixed layer along 10oS in
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the South Atlantic during austral winter. In midlatitude BL/CLs occur mainly in winter and early
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spring. In boreal winter BLs exceeding 60 m are observed in the North Pacific subpolar gyre
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(Fig. 2a). Similarly thick BLs occur in the Atlantic Ocean north of the Gulf Stream. In both
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locations the BLs appear coincident with a seasonal increase of precipitation.
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Sea surface salinity (SSS) increases drastically across the Gulf Stream front leading to a switch
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from the BL regime north of the front to a CL regime south of the front (Fig. 2a). Thick CLs
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(thicker than 30m) are also observed along the Gulf Stream due to cross-frontal transport of low
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salinity water. And even thicker CLs (thicker than 60m) are observed further northeast along the
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path of the North Atlantic Current where its warm, salty water overlies cooler, fresher water.
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Interestingly, despite the presence of warm and salt western boundary currents in both the
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Atlantic and Pacific Oceans, the winter CLs are much less pronounced in the North Pacific than
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the North Atlantic. Explanation for the basin-to-basin difference likely lies in the higher surface
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salinity of the Atlantic (Fig. 2a) and consequently larger values of S (Fig. 3a).
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CLs are also evident in the southern subtropical gyres of the Pacific and Atlantic Oceans as well
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as the South Indian and Southwest Pacific Oceans (Fig. 2b) south of the 30oS SSS. The presence
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of CLs in this region reflects the northward advection of cold and fresh water which subducts
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under the water of the SSS maximum (Sprintall and Tomczak, 1993; note correspondence
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between CL in Fig. 2b and S in Fig. 3b).
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A similar mechanism is used to explain BL formation in subtropical gyres (e.g. Sato et al., 2006).
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In the southern Indian Ocean BLs north of 30oS form as salty water from the region of the SSS
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maximum subducts northward under relatively fresh surface water (Fig. 2b). Similar meridional
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dipole-like patterns with CLs to the south and BLs to the north of local subtropical SSS maxima
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are seen during austral winter in the South Pacific and the South Atlantic in the regions of
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downward wEk (Fig. 2b). In the Northern Hemisphere the areas of CLs in the North Pacific and
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Atlantic Oceans extend well north of the downward wEk regions (Fig. 2a) suggesting that
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horizontal transport of warm salt waters by the western boundary currents plays a role.
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Spatial patterns of BL/CL (Fig. 2) are in close correspondence with the spatial patterns of the
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vertical changes of salinity, S , (Figs. 3a, 3b). As expected, the barrier layers are distinguished
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by a stable salinity stratification, S  0 , where salinity increases downward below the mixed
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layer. In contrast, compensated layers have unstable salinity stratification, S  0 . As discussed
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above, regions of fresh mixed layer trace major areas of precipitation (like the Intertropical
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Convergence Zone) and river runoff (the Bay of Bengal). Different type of BL is observed on the
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equatorward flanks of subtropical salinity maxima. In these areas the ocean accumulates salt due
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to an excess of evaporation over precipitation. Here the equatorward propagation of subducted
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water produces meridional dipole-like BL/CL and S structures that are most pronounced
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during local winter in the Southern Hemisphere (Figs. 2b, 3b). Here barrier layers are observed
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to the north of SSS maxima where mixed layer tops saltier water below while compensated
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layers are observed to the south where mixed layer is saltier than thermocline.
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The spatial patterns of the bulk Turner angle (Figs. 3e,f) indicate that majority of CL cases are
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associated with warm, salty water overlaying colder, fresher water (thus Tub >45o). CLs widen
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during winter leading to higher values of Tub in winter than summer. When cold, fresh water
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overlays warmer, saltier water a CL can also be formed where Tub <-72o. This latter type of
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density compensation is observed only in limited regions of the Labrador Sea and near Antarctica
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during winter.
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3.2 Subseasonal variability Subseasonal variability of BL/CL thickness is similar in amplitude to
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seasonal variability (compare Fig. 4 and Fig. 2). Subseasonal variability is stronger at higher
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latitudes reflecting weaker temperature stratification. The highest variability of BL/CL thickness
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of up to 100m occurs in the North Atlantic. It is also high in the western Atlantic and Pacific due
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to the impact of subseasonal variations in precipitation and river discharge.
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We next consider BL/CL thickness separated by season and roughly 15-year averaging periods
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(Fig. 5). During the first period 1960-1975 thick BLs are evident during local winter in the North
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Pacific, tropical western Pacific and western Atlantic, northern Indian Ocean, and Southern
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Ocean (the latter being evident even in austral summer). CLs during this early period appear
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primarily in the eastern North Atlantic. Little can be said about the existence of BLs in the
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Southern Ocean in austral winter due to the lack of data during this period. By the last period,
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1991-2007 several changes are evident. CLs have appeared in the subtropical North Pacific
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during winter and have strengthened in the eastern North Atlantic. Strong CLs are also evident
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on the northern side of the Circumpolar Current during austral winter (in fact these may have
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existed earlier but simply not been observed). By this last period the extensive winter BLs noted
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in the North Pacific in earlier decades had declined. We next examine monthly time series
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describing four regions: (1) BLs in the subpolar North Pacific, NP/BL box (2) CLs in the
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subtropical North Pacific, NP/CL box; (3) CLs in the eastern North Atlantic; and (4) BLs in the
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western equatorial Pacific (the regions outlined in Fig. 5c).
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The monthly time series of the northern subpolar North Pacific BL region and the subtropical CL
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region both show a long-term trend towards thinner BLs and thicker CLs interrupted by
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occasional reversals (Figs. 6a, 6b). Indeed, the subtropical CL region actually supported a 10-
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20m thick BL prior to 1980s. One direct cause of this change from BL to CL seems to be the
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gradual deepening of the late winter-spring mixed layer in the central North Pacific noted by
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Polovina et al. (1995) and Carton et al. (2008). This observed 20 m deepening into the cooler,
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fresher sub-mixed layer water has the effect of strengthening density compensation (the ‘spice
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injection’ mechanism is discussed by Yeager and Large, 2007). Carton et al. (2008) attribute the
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cause of mixed layer deepening to changes in the atmospheric forcing associated with the
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deepening of the Aleutian sea level pressure low after 1976. These changes led to strengthening
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of the midlatitude westerlies and the ocean surface heat loss in the North Pacific, hence the
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deepening of the mixed layer. The deepening of the mixed layer has opposite impacts on the
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width of CLs and BLs. Deepening widens CLs by injecting saltier water from the mixed layer
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into fresher thermocline. In contrast, stronger atmospheric forcing normally destroys near-
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surface BLs by enhancing mixing. These mechanisms likely explain the narrowing of BLs and
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the widening of CLs in the North Pacific during recent decades (Figs. 6a, 6b).
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The BL and CL thicknesses in the NP/BL and NP/CL boxes both seem to be associated with
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wind changes resulting from changes in the Aleutian pressure low (Fig. 7). Widening of CLs
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(that corresponds to the inverse of regression patterns shown in Fig. 7a) is linked to
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strengthening of westerly winds and latent heat loss (hence, deeper mixed layers) in the NP/CL
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box. Winds also strengthen in the NP/BL box to the north. This wind strengthening may reduce
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the BL due to enhanced mixing. But, in addition the strengthening dry northerly winds to the east
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of the low that decreases moisture transport from the south and thus reduces precipitation that is
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vital to maintain the BL (Fig. 7b). Limitations in the historical data record limit the extent to
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which these relationships can be identified based on data analysis. To overcome this data
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limitation we examine the processes regulating the mixed layer of the CM2.1 coupled model.
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We begin by comparing the simulated and observed winter/spring BL/CL distributions. In the
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simulation as in the observations a BL forms in the North Pacific north of approximately 35oN
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and a CL south of the BL region (compare Fig. 2a, Fig. 8a). The southern CL region can be
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divided into two zones. The first CL zone lies in the Kuroshio-Oyashio confluence at
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approximately 40oN in the northwest. The second CL zone extends across the North Pacific
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along approximately 30oN. This latter zone corresponds to the NP/CL region previously defined
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and is collocated with the subtropical SSS maximum and is bounded to the north by the northern
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edge of the downward wEk area (Fig. 8b). Subduction to the north of the subtropical SSS
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maximum produces negative salinity stratification in the region of SSS maximum while
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subduction at the SSS maximum produces positive salinity stratification to the south (see
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examples in Fig. 9).
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We next explore year to year changes in the NP/CL region by contrasting March 1953 when the
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simulated CL is the thickest of the 140 year record (25m) with March 1956 when the CL has
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been replaced by a 10m thick BL (Fig. 8c). The striking difference between these two years
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seems to be the different surface forcing. In 1953 downward wEk extends throughout the
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subtropics leading to a deep mixed layer (Fig. 9a). During this year the isohaline contours along
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180oE indicate an exaggeration of the normal situation in which there is southward advection of
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high salinity water below the fresh nearsurface layer. These conditions lead to a CL layer north
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of 25oN and BL layer to the south. In contrast, in 1956 wEk was upward north of 28oN trapping
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the freshwater introduced through precipitation in the surface layer and thus causing a shallow
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BL to form. The lag correlations in Fig. 8d confirm that CLs in the NP/CL region thicken in
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years when the mixed layer is anomalously deep, in line with the spice injection mechanism of
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Yeager and Large (2007), and BLs in the NP/BL thin as a result of the enhanced mixing
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associated with deep mixed layers. Thus deepening of the subtropical mixed layer may explain
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the trends observed in Figs. 6a, 6b.
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Correspondence between temporal changes of BL/CL width and atmospheric pressure patterns
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seen in both, observations and simulations in the north Pacific suggests that similar
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correspondence with local atmospheric patterns may take place in the north Atlantic. But the
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time series of variations in CL region of the North Atlantic is striking in showing primarily year-
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to-year variations (Fig. 6c) even though the winter-spring meteorology of this region does
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exhibit decadal variations frequently referred to as the North Atlantic Oscillation (Hurrell, 1995).
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We suspect that some of this variability is the result of the smaller spatial scales of variability
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relative to the North Pacific which has been insufficiently sampled by the historical observing
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systems. During the past ten years when the data coverage has been increasingly extensive (due
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to implementation of the Argo floats) CL thickness in the North Atlantic CL region has
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undergone 15m variations from year-to-year with thick layers in 1999, 2002-3, and 2006-7 and
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thinner layers during the other years (Fig. 10a). This pattern of CL variation is related to the
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strength of the subtropical high in sea level pressure (Fig. 10a). Indeed, a regression of sea level
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pressure, latent heat flux, and winds on CL thickness in the North Atlantic CL box shows basin-
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scale relationships (Fig. 10b). Vertical expansion of CLs (more negative CL width) corresponds
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to NAO-like pattern of the atmospheric pressure with stronger Azores high and deeper Iceland
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low. This pressure pattern strengthens winds and latent heat losses in the east while weakens
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them in the central and western subtropics (Fig. 10b). Anomalously strong winds and net surface
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heat losses suggest deepening of the mixed layer that, in turn, expands width of CLs by the spice
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injection mechanism. This is in line with 2002-3 and 2006-7 data when anomalously thick CLs
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in the NA/CL box occur in phase with anomalously deep mixed layers. But in 1999 this
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relationship is opposite (Fig. 10a). This may suggest that alternative mechanisms including salt
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advection play a role. Indeed, in the north-east it is clear that advection plays a central role in
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variability of the mixed layers (Nilsen and Falck, 2006 show this process at work in the
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Norwegian Sea).
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Finally, we consider year-to-year variability of the BL region of the tropical Pacific. BL thickness
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in the western Pacific west of 160oE, which is generally 10-20m, varies in thickness coherently
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with the value of the Southern Oscillation Index (Fig. 11c). The eastward shift of precipitation
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during El Nino decreases precipitation west of 160oE by 20% or 1 mm/dy (in response to a 20
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unit decrease of the SOI) and as a result BLs shrink by 20% or 2-5m (Figs. 11a, 11b). Thus, in
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this region variations in BL thickness respond primarily to surface freshwater flux. East of 160oE
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BLs are thinner and occur at much shallower depths so that they are affected more strongly by
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solar radiation, advection and entrainment.
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4. Conclusions
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This study examines subseasonal changes of barrier and compensated layers (BLs and CLs)
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based on analysis of profiles of temperature and salinity covering the years 1960-2007. Because
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of data limitations we focus mainly on the Northern Hemisphere and tropics. The processes that
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regulate subseasonal variability of BL/CL thickness are similar to those which regulate their
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seasonal appearance: fluctuations in surface freshwater flux, Ekman pumping, and horizontal
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advection. Thus, the spatial distribution of subseasonal variability reflects aspects of the
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subseasonal variability of these forcing terms.
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The most striking interannual variability of BL width occurs in the rainy tropical oceans with
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deep thermoclines such as the Arabian Sea and Bay of Bengal in the Indian Ocean and the
378
western tropical Pacific under the Intertropical Convergence Zone and the South Pacific
379
Convergence Zone. Precipitation in these regions varies strongly interannually. During high
380
precipitation years the mixed layers in these regions show capping fresh layers and thick BLs. In
381
contrast, during low precipitation years mixed layer salinities increase and BL thickness
382
decreases. Thus, in the western Equatorial Pacific between 120oE and 160oE, the BL thickens by
383
2m to 5m during La Nina while during the El Nino the BL thins by a similar amount. In the
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central and eastern equatorial Pacific variations of BL thickness are not nearly as well correlated
385
with the phase of ENSO. We think this happens because in the central basin salt advection plays
386
an important role in regulating BL thickness as well, while further east the mixed layer shallows
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and mixed layer entrainment variability becomes increasingly important.
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In the subtropics and midlatitudes during late winter-spring we find alternating regions of CLs
390
and BLs in the seasonal climatology. The northern tropics of both the Pacific and Atlantic (the
391
southern edge of the subtropical gyres) show broad regions of BLs where salty subtropical
392
surface water formed further north has subducted, advected equatorward, and affected the water
393
properties of the winter mixed layer. Within the evaporative subtropical North Pacific and eastern
394
North Atlantic we find CLs resulting from mixed layers with positive temperature stratification
395
but negative salinity stratification. In the eastern subpolar North Atlantic this CL merges with a
396
subpolar CL resulting from negative temperature stratification but positive salinity stratification.
397
In the North Pacific, strengthening of the Aleutian pressure low during successive winters, thus
398
strengthening the midlatitude westerly winds has led to deeper mixed layers, cooler SSTs, and a
399
long-term increase in the thickness of this CL. Further north in the North Pacific a supolar BL
400
region is evident, which does not have a counterpart in the North Atlantic. The same changes in
401
meteorology which include strengthening of the Aleutian pressure low have led to an increase in
402
dry northerly winds which in turn cause a thinning of this BL from ~40m before 1980s to ~ 20m
403
afterwards.
404
405
In the North Atlantic the broad area of CL also undergoes fluctuations in thickness associated
406
with surface meteorology. In this basin the fluctuations in winter thickness also seem to vary in
17
407
response to meteorological changes associated with changes of the Azores sea level high
408
pressure. Estimation of the timescale of variability is more difficult in this basin as the analysis is
409
more noisy. Here our examination focuses on the changes that are apparent in the shorter 1997-
410
2007 period when the data coverage is better. During this period the thickness of the winter-
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spring CL in the northeastern Atlantic undergoes 4-5 year timescale fluctuations from 45m to
412
60m in phase with the strengthening/weakening of the Azores high.
413
414
Determining the basin-scale BL/CL structure tests the limits of the historical observing system.
415
Further progress in understanding BL/CL variability and its role in air-sea interactions will likely
416
require further exploration of models that provide reasonable simulations of observed variability.
417
418
Acknowledgements We gratefully acknowledge the Ocean Climate Laboratory of the National
419
Oceanographic Data Center/NOAA and the Argo Program data upon which this work is based.
420
Support for this research has been provided by the National Science Foundation (OCE0351319)
421
and the NASA Ocean Programs.
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18
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Figure captions.
Figure 1. Climatological cold season barrier layer/compensated layer thickness versus the
496
bulk Turner angle evaluated using temperature ( T ) and salinity ( S ) difference between the
497
top and bottom of a barrier or compensated layer. Vertical bars show the mean and the standard
498
deviation for consecutive 22.50 interval of the Turner angle. Grey dots are the cold season data
499
that aggregate January-March data from the Northern Hemisphere and July-September data from
500
the Southern Hemisphere. The Turner angle range -720 to 450 corresponds to BL. CL occurs
501
outside this interval.
502
Figure 2. Climatological (a) January-March and (b) July-September barrier layer width
503
(positive) and compensated layer width (negative). Climatological sea surface salinity (  35 psu
504
solid contours, below 35 psu dashed contours) and downward wind-driven vertical velocity
505
(hatching) are overlain. Salinity data are from Boyer et al. (2006).
506
Figure 3. (a,b) Salinity ( S ) and (c,d) temperature ( T ) difference between the top z t =
507
min[MLT, MLD] and the bottom zb =max[MLT, MLD] of barrier/compensated layer, (e,f) bulk
508
Turner angle calculated from S and T between the same two depths. (left) January - March
509
(JFM) values, (right) July - September (JAS) values. Turner angles in the range from -72o to 45o
510
corresponds to barrier layers, while compensated layers occur outside this range
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Figure 4. Standard deviation (STD) of monthly anomalous BL/CL thickness.
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Figure 5. Quasi-decadal mean barrier layer (positive) and compensated layer (negative)
514
thickness in (left) northern winter and (right) austral winter. Rectangles in (c) show locations of
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the North Pacific barrier layer box (NP/BL 140oE-140oW, 50o-60oN), North Pacific compensated
516
layer box (NP/CL 140oE-160oW, 25o-35oN), equatorial Pacific barrier layer box (EP/BL 120oE-
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517
160oE, 5oS-5oN), and North Atlantic compensated layer box (NA/CL 30oW-10oW, 40o-60oN)
518
Figure 6. Box averaged BL/CL width in (a) North Pacific barrier layer region, (b) North
519
Pacific compensated layer region, and (c) North Atlantic compensated layer region. Thin lines
520
are monthly average values, bold lines are January-March average values. JFM values are shown
521
if at least 10 data points are available. Box locations are shown in Fig. 5c.
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Figure 7. Time regression of JFM barrier layer thickness in (a) North Pacific compensated
523
layer box on latent heat flux (Wm-2/m, shading), winds (ms-1/m, arrows), and sea level pressure
524
(mbar/m); (b) North Pacific barrier layer box on precipitation (mm h-1/m). See also Fig.5c for
525
box locations. Atmospheric parameters are from the NCEP/NCAR Reanalysis.
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Figure 8. CM2.1 model simulations in the north Pacific. (a) March climatological BL/CL
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width with climatological sea surface height (SSH, CI=0.2m) and sea surface salinity (SSS, bold
528
contours) overlain. CL and BL boxes are drawn in red and blue, respectively. (b) Zonally
529
averaged March climatological Ekman pumping (positive – downward is shaded gray). (c) Time
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series of the box-averaged CL width. Two contrasting model years with thick (1953) and thin
531
(1956) CL are marked by ‘o’. (d) Lagged correlation of anomalous (blue) box-averaged CL
532
width and MLD, (red) box-averaged BL width and MLD. (e) Time regression of March box-
533
averaged CL width on wind stress and Ekman pumping (mdy-1/m).
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Figure 9. Latitude-depth section of salinity along 180oE for March of the two contrasting
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model years with thick (1953) and thin (1956) CL. Ekman pumping (positive – downward is
536
shaded gray) is shown against the right-hand vertical axes.
537
Figure 10. (a) JFM compensated layer width (solid) and monthly mixed layer depth
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(compensated layer depth range between MLT and MLD is shaded) averaged in the North
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Atlantic compensated layer box, JFM anomalous mean sea level pressure (dashed) averaged
24
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40oW-30oW, 30oN-40oN. (b) Time regression of the 1997-2007 JFM compensated layer width on
541
latent heat flux (Wm-2/m, shaded), anomalous mean sea level pressure (mbar/m, contours), and
542
winds (ms-1/m, arrows).
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Figure 11. Lag regression of SOI index on 5oS-5oN averaged (a) anomalous barrier layer
544
width, (b) precipitation. Time mean BL width (solid), salinity (dashed), and precipitation are
545
shown against right hand vertical axes. Precipitation is Xie and Arkin (1997). (c) Time series of
546
annual running mean SOI (shaded) and anomalous BL width averaged in the western equatorial
547
Pacific box (120oE-160oE, 5oS-5oN). See Fig. 5c for the box location.
25