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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 111, C12006, doi:10.1029/2006JC003680, 2006
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Surface wind response to oceanic fronts
Qingtao Song,1 Peter Cornillon,1 and Tetsu Hara1
Received 2 May 2006; revised 11 August 2006; accepted 24 August 2006; published 8 December 2006.
[1] The response of surface winds to ocean fronts characterized by sharp gradients in both
sea surface temperature (SST) and ocean currents was analyzed using scatterometer
(NSCAT and QuikSCAT) wind data and Gulf Stream path positions in conjunction with
simulations made with the Pennsylvania State University (PSU)-National Center for
Atmospheric Research (NCAR) Mesoscale Model (MM5). All match-ups, between
each scatterometer pass and the Gulf Stream path, were visually examined and only those
for which the wind field was free of atmospheric fronts or large curvature over a
reasonably straight segment of the Gulf Stream were selected. Ten match-ups met these
criteria for the period studied from 16 September 1996 to 29 June 1997 for NSCAT
and from 24 July 1999 to 31 December 2000 for QuikSCAT. Changes in the modeled
surface wind field across the front in each of the ten cases agree well with changes in the
observed winds. Our findings suggest that the perturbation pressure gradient resulting
from the thermal forcing by the front accounts for the decrease in wind speed when
moving from warm to cold water and the increase observed in the converse. In the cases
examined, the adjustment of the surface wind to the front occurred as a result of the
vertical motion induced by horizontal divergence/convergence and advection in the marine
atmospheric boundary layer (MABL). The dynamical forcing associated with strong
surface currents is also shown to modify scatterometer-derived winds. Finally the numerical
simulations suggest that the dynamical and thermal effects are very nearly additive.
Citation: Song, Q., P. Cornillon, and T. Hara (2006), Surface wind response to oceanic fronts, J. Geophys. Res., 111, C12006,
doi:10.1029/2006JC003680.
1. Introduction
[2] Modification of surface winds within the Marine
Atmospheric Boundary Layer (MABL) by sea surface
temperature (SST) gradients has been observed and modeled for more than two decades in different parts of the
world ocean including the eastern equatorial Pacific [e.g.,
Chelton et al., 2001; de Szoeke and Bretherton, 2004;
Chelton, 2005], the North Atlantic [Sweet et al., 1981;
Businger and Shaw, 1984; Friehe et al., 1991; Park and
Cornillon, 2002], the Kuroshio Extension east of Japan
[Nonaka and Xie, 2003], the Agulhas Current south of
Africa [e.g., Lee-Thorp et al., 1999; O’Neill et al., 2005],
the Southern Ocean [O’Neill et al., 2003], the Arabic Sea
[Vecchi et al., 2004] and the Azores region [e.g., Giordani et
al., 1998]. An understanding of this one-way forcing from
the ocean to the atmosphere has progressed from the early
work of Sweet et al. [1981], who observed intensification of
surface winds over the warm water of the Gulf Stream, to
the recent work of Park et al. [2006], who observed
significant modification of surface winds over Gulf Stream
rings. Results from numerous observations indicate that the
structures of temperature, wind and turbulent fluxes in the
1
Graduate School of Oceanography, University of Rhode Island,
Narragansett, Rhode Island, USA.
Copyright 2006 by the American Geophysical Union.
0148-0227/06/2006JC003680$09.00
MABL differ significantly from one side of an SST front to
the other, with an increase (decrease) in wind speed as the
wind blows from cold (warm) to warm (cold) water across
the front [Friehe et al., 1991].
[3] While observations have yielded conclusive evidence
for a positive correlation between surface winds and underlying SST fronts [Xie, 2004; Chelton et al., 2004; Chelton
and Wentz, 2005], the mechanisms for the surface wind
response to oceanic forcing have yet to be clearly articulated.
The mechanisms have been studied both observationally
and in numerical simulations. Lindzen and Nigam [1987]
used a one-dimensional (1D) planetary boundary layer
(PBL) model to study the surface wind response to tropical
instability waves (TIW) in the eastern equatorial Pacific for
time scales greater than one month. They concluded that
much of the horizontal structure in the MABL wind field is
driven by horizontal pressure gradients that develop in
response to the boundary layer baroclinicity induced by
the underlying SST gradients. They argue that near the
equator the wind should be particularly strong directly over
large SST gradients. In the following, we refer to this as the
sea level pressure (SLP) driving mechanism.
[4] Wallace et al. [1989] and Hayes et al. [1989] went on
to further study the effect of SST gradients on surface winds
in the eastern equatorial Pacific. They found that on
seasonal and interannual timescales the prevailing southeasterly trade winds are accelerated over warm SST anomalies and decelerated over cold anomalies. The surface winds
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were strongest over warm water to the north of the strongest
SST gradients. Wallace et al. argued that if the SLP driving
mechanism was the dominant forcing, the strongest southerly winds would be observed directly over the oceanic
frontal zone rather than to the north of it. The positive
correlation between wind speed and SST supports the
hypothesis of Wallace et al. that, in the eastern equatorial
Pacific, modification of boundary layer shear rather than
changes in the atmospheric SLP gradient is the dominant
process affecting surface winds on the time scales studied.
They argued that the stronger surface wind over warmer
water is due to efficient turbulent convection that mixes
momentum down to the surface (referred as the momentummixing mechanism).
[5] While many observational analyses support the momentum-mixing mechanism in the eastern tropical Pacific
[Chelton et al., 2001; Polito et al., 2001], recent findings
from Tropical Atmosphere Ocean (TAO) buoy observations
[Cronin et al., 2003] and numerical simulations [Small et
al., 2003, 2005] of surface wind response to TIW demonstrate the importance of the pressure gradient force in
driving spatial changes of the cross-equatorial flow.
[6] Away from the tropics, Sweet et al. [1981] documented the modifying effects of air-sea interaction and
low-level wind speeds, sea state and the distribution of
low-level air temperature during a flight over the Gulf
Stream’s north wall. The momentum-mixing mechanism
was proposed to explain enhancement of low-level
winds over warm water. Recent studies of the extratropical
MABL response to a discontinuity in SST are consistent
with the momentum-mixing mechanism [Park and
Cornillon, 2002; Nonaka and Xie, 2003; O’Neill et al.,
2003; Vecchi et al., 2004].
[7] Wai and Stage [1989] made use of a two-dimensional
(2D) MABL model to analyze the dynamical modification
of the MABL structure by an oceanic front. Their analysis
of the cold-to-warm SST case showed that, in addition to
enhancement of vertical mixing in the MABL when the air
mass is advected from the cold to the warm side, the
pressure gradient force is important in the momentum
budget, similar to the SLP-driving mechanism. Warner et
al. [1990] and Doyle and Warner [1992] performed threedimensional (3D) and 2D numerical modeling experiments
to study the impact of SST resolution on mesoscale coastal
processes, and they found that the MABL structure is very
sensitive to the SST distribution. From their simulations, a
secondary circulation, ascending over the warm water and
descending over the cold water across SST fronts suggests
that there might exist other physical processes with the
MABL in extratropical regions.
[8] Most previous studies using satellite data have
focused on seasonal and interannual variability. However,
changes in SST also exhibit pronounced fluctuations in the
atmosphere on shorter temporal scales of order days [Friehe
et al., 1991; Park and Cornillon, 2002]. The atmosphere
has most of its energy in scales of several hundred kilometers and longer [Weller, 1991]. However, traditional in
situ observations (using aircraft and ship) have shown that
the spatial variation of the MABL, forced by mesoscale
variability in SST, is characterized by scales from a few
kilometers to several tens of kilometers [Guymer et al.,
1983; Friehe et al., 1991; Lee-Thorp et al., 1999].
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[9] The focus of this study is on the response of surface
winds in the MABL to mesoscale fluctuations of SST.
Although the one-way forcing from ocean to atmosphere
at large scale has been documented in a number of studies
(see the review by Xie [2004]), knowledge of the mesoscale
response of surface winds to strong oceanic fronts is limited.
The limitation is due to limited observations near oceanic
fronts by traditional aircraft and ship-based measurements.
Moreover, synoptic variability in the atmosphere and strong
atmospheric forcing at midlatitudes, discussed in previous
field studies, obscure the signatures of oceanic forcing.
[10] This study begins with a quantification of the principal mesoscale characteristics of the MABL for different
wind conditions in the vicinity of the Gulf Stream. The
modification of the MABL forced by the Gulf Stream is
examined with scatterometer wind data in conjunction with
Advanced Very High Resolution Radiometer (AVHRR)
SST data. Cases are selected based on a set of filters chosen
to ensure that oceanic forcing is the dominant process
modifying the observed wind field. This simplifies interpretation of the atmospheric response to the oceanic forcing.
The observed modifications of the MABL are then analyzed
with the Pennsylvania State University (PSU)-National
Center for Atmospheric Research (NCAR) Mesoscale
Model (MM5) configured such that the mean surface wind
stress upwind of the SST front equals that observed by the
scatterometer and the mean vertical profiles on the upwind
and downwind boundaries of the study domain equal those
for the corresponding day in the National Centers for
Environmental Prediction (NCEP) reanalysis. The main
objective of this analysis is to characterize the responses
of the MABL and the physical processes in determining
these responses.
[11] In addition to the modification of the MABL by an
SST front outlined above, scatterometer winds are also
modified by strong oceanic currents [Cornillon and Park,
2001; Park and Cornillon, 2002; Kelly et al., 2001]. The
Gulf Stream brings subtropical warm water northward, thus
is characterized by a sharp SST front associated with a
strong surface current about 1.5 m s1 downstream of Cape
Hatteras [Rossby and Zhang, 2001]. Two interesting questions related to this are, first, to what extent ocean currents
affect the scatterometer observations (referred to as oceanic
dynamical forcing), and second, how this oceanic dynamical forcing is related to the oceanic frontal forcing. To
quantitatively address these questions, the bottom boundary
condition of the PSU-NCAR MM5 was modified to allow
for a surface ocean current and model runs were performed
with and without surface ocean currents.
2. Data and Data Processing
[12] The observational response to oceanic fronts was
investigated by using a combination of AVHRR SST data
and scatterometer (NSCAT and QuikSCAT) wind data. To
ensure the maximum coverage of the Gulf Stream the study
encompasses the geographical area bound by 30°N – 50°N
and 80°W – 45°W. The period studied, determined by the
availability of AVHRR data, is from 16 September 1996 to
29 June 1997 for NSCAT and from 24 July 1999 to
31 December 2000 for QuikSCAT. The AVHRR SST data
were obtained at 1 km resolution via OPeNDAP from
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Figure 1. Ten selected cases in the vicinity of the Gulf Stream. Thick lined rectangles define the
sampling area for the two cases discussed in detail. Arrows represent the direction in which the wind was
blowing upstream of the front for these cases. Thin lined boxes define the sampling area for the remaining
eight cases. Circles define the mean path of the Gulf Stream for the ten cases.
University of Rhode Island (URI). NSCAT and QuikSCAT
wind data were obtained at 25 km via OPeNDAP from the
National Aeronautics and Space Administration (NASA)Jet Propulsion Laboratory (JPL).
[13] To minimize the effect of cloud contamination, the
AVHRR SST fields were composited in two-day intervals
by retaining the warmest SST value at each pixel location
for each two day period. Cornillon et al. [1987] found that a
two-day compositing period is sufficiently long to provide a
moderately clear view of the Gulf Stream region while short
enough to avoid substantial errors due to the meandering of
the stream. The path of the Gulf Stream was then manually
digitized from two-day composite SST images obtained
from the AVHRR sensor. The points digitized generally
followed the maximum gradient contour thought to be
associated with the stream. The digitized path of the stream
is referred to as the path of the stream, the stream’s path or,
simply, the stream path in the following. The rms of the
stream’s path relative to the path defined by the 12°C
isotherm at 500 m is about 14 km [Cornillon and Watts,
1987].
[14] Only wind vectors with the best quality index were
used. QuikSCAT wind data have been found to be the most
accurate in areas of little to no rainfall and moderate wind
speed [Draper and Long, 2002]. Rain flagged winds are
considered suspect [Draper and Long, 2004] and therefore
eliminated in this study. Simulations and preliminary analysis of QuikSCAT data show that the scatterometer-derived
wind quality is degraded near nadir and near the swath edge
[Jet Propulsion Laboratory (JPL), 2001]. Because we are
interested in very accurate winds, we did not use winds
within 100 km of the edge of the scatterometer swath or
winds within 100 km of nadir for QuikSCAT.
[15] The focus of this study is on the response of the
MABL to an SST step in the ocean. Because this effect is
moderately weak compared with natural variation in the
atmosphere it was necessary to filter out observations for
which the natural variation in the wind field was large. Also,
to facilitate the numerical simulations, it was desirable to
consider cases for which the front, the Gulf Stream in this
case, was relatively straight and the upwind side of the front
was moderately free of large SST variations resulting from
rings and other oceanographic phenomena. A two step
screening process was used to achieve this. First, rough
AVHRR-scatterometer match-ups were identified for which
the scatterometer wind fields were slowly varying within
±250 km of a coincident (within 2 days), ‘straight’ Gulf
Stream path segment. Fifty-nine match-ups met these criteria. In the second step, a quality index was assigned to each
of these rough match-ups based on the following criteria:
(1) the homogeneity of the upstream wind field; (2) the
homogeneity of the SST field, both upwind and downwind
of the front; (3) the fraction of the SST field that was cloud
covered in the area of interest; and, (4) the distance between
land and the stream path for those cases in which the wind
was blowing from land. For each of these criteria, the
match-up was assigned either a 0, it did not meet the
minimum standard for that criterion, or a 1, it did meet
the standard. This assignment was made by visual inspection of the data. The ten cases, assigned a score of 3 or
more, were selected for analysis (Figure 1).
[16] Following selection, a parallelogram was defined
with two sides parallel to, and the same length as, the
straight line, least square fit to the Gulf Stream path segment
of interest. The orientation of the other two sides of the
parallelogram was determined by the angle that the mean
wind upwind of the stream made relative to the path of the
stream, the availability of data in the area, the proximity of
land and the location of other significant oceanographic
features. In all cases, the length of the sides not parallel to
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the stream segment is 800 km centered on the path of
the stream. This parallelogram defined the study area. The
parameters used to force the simulations introduced in the
next section and the values against which the simulation
results are compared were obtained from this study area. A
grid was then defined with one axis parallel to the least
square fit straight line and the other axis normal to this line.
The grid spacing is 0.35° of longitude in the along-path
direction and 25 km in the across-path direction. For each
selected scatterometer pass, the observed wind field was
interpolated to this grid such that all grid elements in the
parallelogram for which good scatterometer data existed
were filled. The u component of wind is defined to be
parallel to the front (the short side of the box), positive in
the downstream direction for the cases discussed herein.
The v component is in the cross-front direction, positive
shoreward. Winds greater than 15 m s1 and less than
3 m s1 were omitted from the analysis.
3. Numerical Model
[17] The vertical structure of the atmosphere is very
important for understanding the mechanisms by which it
adjusts to SST variations [e.g., Friehe et al., 1991; Song et
al., 2004; Xie, 2004]. Since scatterometer-derived winds do
not measure vertical structure, numerical modeling, validated
against observations, is used to gain insight into the response of the MABL and the processes giving rise to this
response. The numerical model used in this study is the fifthgeneration PSU-NCAR MM5. MM5 is a 3D, multi-nested,
non-hydrostatic or hydrostatic (only the non-hydrostatic
configuration was used in this study), terrain-following
s-coordinate model designed to simulate or predict mesoscale and regional scale atmospheric circulation [Dudhia,
1993; Grell et al., 1995], where s = (p pt)(ps pt)1,
p is the pressure, pt is the pressure at the top of the model
atmosphere (100 mb), and ps is the surface pressure.
[18] The performance of MM5 in the vicinity of an SST
front was evaluated through comparison of MM5 simulations with in situ observations made during the Frontal AirSea Interaction Experiment (FASINEX) [Friehe et al.,
1991; Song et al., 2004].
[19] To simulate MABL processes at high vertical resolution, five PBL schemes are available in MM5: the
Blackadar high-resolution scheme [Zhang and Anthes,
1982], the Medium-Range Forecast (MRF) scheme [Hong
and Pan, 1996], the Burk-Thompson scheme [Burk and
Thompson, 1989], the Eta scheme [Janjic, 1994] and
Gayno-Seaman scheme [Shafran et al., 2000]. These PBL
schemes are logically grouped into two general categories:
local schemes (Burk-Thompson, Eta and Gayno-Seaman)
and non-local schemes (Blackadar and MRF).
[20] The simulations using MM5 with the MRF boundary-layer model provided results at the mesoscale that are
statistically closer to the FASINEX observations than those
obtained with MM5 using any of the other four boundary
layer models [Song et al., 2004]. Therefore MRF was
selected for this study.
[21] To resolve the MABL response to changes in SST as
well as to compare this response to the scatterometerderived winds (resolution 25 km), MM5 was configured
with regional and local computational domains with hori-
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zontal grid spacings of 15 km and 5 km respectively. A total
of 40 s levels, with 25 concentrated in the PBL, the lower
2 km, were specified for both regional and local domains.
Physical options for all the model simulations include
Dudhia’s simple ice scheme and Grell’s cumulus parameterization. The cloud-radiation scheme has been selected to
account for longwave and shortwave interactions with
explicit cloud and clear-air, allowing diurnal variation
during the simulation. Non-hydrostatic dynamics with a
3D Coriolis force were used because of the high spatial
resolution.
[22] For all ten match-ups, the vertical profile used to
initialize the temperature and moisture fields in MM5 was
obtained from the NCEP reanalysis, available at 6-h intervals with a horizontal grid spacing of about 275 km. For
studies of the MABL response to horizontal variability in
the SST field using MM5, the vertical structure of the
atmosphere obtained from the NCEP reanalysis is an
acceptable proxy for the actual vertical structure in the Gulf
Stream region for the initialization of the model [Song et al.,
2004].
[23] The SST field used for each simulation was obtained
by averaging the two-day composite AVHRR SST fields in
the along-front direction across the sampling box and
interpolated into the MM5 outer domain at a grid spacing
of 15 km.
[24] In the past, MM5 has been configured for a bottom
boundary that is at rest, with the resulting wind stress of the
bottom on the atmosphere being in the opposite direction as
the wind. In general, this is not the case for the ocean,
especially in the vicinity of strong currents such as the Gulf
Stream; the direction of the current relative to the wind must
be taken into account in order to determine the wind stress
exerted by the ocean on the atmosphere. In light of this it
was necessary to modify the bottom boundary condition in
the model (see Song [2006] for details). For this study, the
bottom boundary was configured with a surface current
representing the Gulf Stream that was adapted from Rossby
and Zhang [2001]: a sharp velocity maximum (1.5 m s1)
flanked by rapidly decreasing velocities on either side. The
peak of the surface current speed is located 25 km to the
south of the digitized path.
[25] The atmospheric flow field above the MABL was set
to be horizontally uniform. Model boundary conditions,
which were held constant for the run, were adjusted such
that the upstream near surface wind at the end of the
simulation, the time at which the comparison between the
model simulations and the scatterometer observations was
made, matched the u and v components of the scatterometer
observations.
4. Results
[26] Each simulation involved a 15-h integration. Six
hours into the run, the fine-resolution local domain began
to run interactively with the outer domain. During the
simulation, the SST and sea surface current fields were held
constant. The MABL reached nearly steady state approximately 6 hours into the fine resolution simulation (approximately 12 hours into the run). The model results presented
in this paper are those at the end of the 15 hour integration.
The starting time of each simulation was chosen so that the
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Table 1. Numerical Experiments
Experiment Name
Sea Surface Current
SST Front
SIMCS
SIMNC
SIMNCS
Yes
No
No
Yes
Yes
No
ending time of the 15 hour integration is close to that of the
scatterometer pass.
[27] Scatterometers respond to wind stress at the surface
as opposed to wind speed, hence the primary product
derived from scatterometer data, and used in this study, is
the neutral equivalent wind 10 m above the sea surface, UN10.
To facilitate the comparison of the MM5 results with the
scatterometer winds, UN10 was calculated for the MM5 runs
as follows:
N
U10
u?
10
¼
ln
k
z0
ð1Þ
where u? is the friction velocity, k is von Karman’s constant,
and z0 is the roughness parameter. Friction velocity u? and
roughness parameter z0 are standard output of the numerical
model.
[28] In order to examine the sensitivity of the MABL
structure to an SST front and the current generally associated with such a front, three model runs were performed. All
of the model runs use identical physical-process parameterizations, the only difference being the SST and sea surface
current forcings (see Table 1). A base run, SIMNCS, was
performed for which there was no SST front and no current.
The SST field for this run was chosen as the mean SST field
on the upwind side of the front and the wind in the free
atmosphere was chosen such that simulated surface winds at
the end of the 15-hour simulation match the upwind mean
of the scatterometer-derived winds. The vertical profile of
the atmosphere was initialized as described in the previous
section.
[29] For each of the ten cases the scatterometer winds
were averaged in the along-front direction across the rectangular box at each of the 25 km grid points to obtain a
mean cross-frontal surface wind section normal to the path
of the stream. The simulation can be regarded as a 2D
simulation; there is little to no variation along the straight
oceanic front. Figure 2 shows scatterometer observations
versus numerical simulations (SIMCS and SIMNC) for
eight match-ups (the other two cases are discussed in detail
below). SST, u and v components of wind, and wind curl are
plotted as a function of distance away from the digitized
path of the stream. The position of the stream’s path is at
zero on the vertical axes in these plots. Positive (negative)
distance represents the shoreward or left (seaward or right)
side of the path looking downstream. The generally good
agreement between full simulations (SIMCS – red lines) and
observations (thick blue lines) for all eight cases indicates
that MM5 captures the proper physics in the MABL forced
by oceanic fronts with the proper initialization and boundary conditions. For the eight cases, the surface wind
increases as it blows from cold to warm water with a
minimum in the curl of the wind close to the front. The
difference between SIMNC (cyan line) and SIMCS indicate
the effect of an ocean current on UN10. The effect of the
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current on the cross-frontal wind stress is discussed in more
detail in section 4.3.
[30] The occasional large differences between the model
runs and the observations near the up- or the downwind
edge of the inner model domain, 400 km from the SST
front, are a result of changes in the large scale meteorology
in the area. They do not result from a response of the wind
to the front.
4.1. The 24 March 2000 Cold-to-Warm Case
[31] Consistent with the overall selection criteria, the
scatterometer winds are uniform upwind of the path of the
stream in the region selected for analysis from the 24 March
2000 match-up (the blue box in Figure 3a). The position of
the stream path was digitized from the SST image (the blue
circles in Figure 3a). The stream is relatively straight across
the study region and no atmospheric fronts are evident near
the study area. The study area was oriented such that the
wind upwind of the stream is approximately parallel on
average to the long side of the box. The wind was blowing
from cold to warm water, very nearly perpendicular to the
stream (Figure 3b), and the SST step across the stream is
large, approximately 9°C over 80 km. The Gulf Stream
(Figure 3a) is about 100 km wide separating the slope water
(green) from subtropical water (orange). Figures 3d – 3g
present the model results (SIMCS - red line and SIMNC cyan line) with the scatterometer observations (thick blue
line). The u and v wind components in Figures 3d and 3e
correspond to the equivalent neutral wind, UN10.
[32] As the vertical structure of the MABL is critical to
understanding the physical processes that modify the wind
as it crosses the stream, a cross section of potential temperature, q, normal to the SST front from the surface to 750 mb
is shown (contours) in Figure 4. Considerable horizontal
variability of the potential temperature field is evident in the
MABL. Above the mixed layer, pressure <925 mb, the
horizontal change in potential temperature is small. The
largest horizontal gradient in q is found across the front at
the surface. The sharp increase of q between 975 mb and
925 mb indicates that the top of the mixed layer is slightly
higher on the downwind warm side of the front. The
upwind, cold side of the front is stably stratified as indicated
by the continuous increase of q. The potential temperature
on the downwind, warm side of the front is well mixed. The
near surface potential temperature over the cold water
(290.5°K) is higher than the underlying SST (288°K).
As the air approaches the front, this SST difference
decreases and then reverses sign as the front is crossed,
292.25°K over 294°K. The air-sea temperature difference
results in suppressed surface heat fluxes over the cold water
and enhanced heat fluxes over the warm water. Over warm
water, increased surface heat fluxes destabilize and deepen
the MABL through strong vertical turbulent mixing.
[33] Surface wind speed (color in Figure 4) increases
from 6.5 m s1 to 9.0 m s1, 27%, as the wind crosses the
front from cold to warm water. Contrary to the increase in
surface wind speed, the wind near the top of the MABL,
centered at 975 mb, shows a gradual decrease across the
front. This characteristic of the wind profile within the
MABL was also reported by Friehe et al. [1991] in
FASINEX and by Wai and Stage [1989] and Small et al.
[2003] using numerical simulations.
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Figure 2. Two numerical simulations (SIMCS – red and SIMNC – thin cyan lines) versus observations
(thick blue lines) for each of the 8 match-ups. SST in °C, u and v in m s1, curl in 105 s1. The arrow in
the SST panel represents the surface wind direction relative to the front. The black line is the 200 m
isobath.
[34] The divergent region in the MABL is fed from the
convergent region above the MABL. This is best seen in the
context of the vertical velocity field shown in Figure 5 with
the horizontal divergence. The horizontal divergence reaches its maximum value, 2 105 s1, immediately above
the SST front. Above this divergent region, there exists a
convergent (negative) region of relatively small magnitude,
4 106 s1. Associated with the horizontally divergent
region, there is a system of up and downdrafts. The core of
the upward motion, with a magnitude of 0.6 cm s1, is
horizontally narrow and extends to approximately the top of
the mixed layer, 925 mb. By contrast, the downward
motion associated with the surface divergent region extends
from the free atmosphere, well above the MABL, down to
approximately the 1000 mb level, just above the large
divergent core immediately over the front. The downdraft
is relatively large in the free atmosphere, 0.8 cm s1,
compared with its magnitude through most of the MABL,
0.4 cm s1.
[35] To examine the extent to which the SST front
modifies the atmosphere, mass fluxes were calculated for
two rectangular boxes, a boundary layer box and a free
atmosphere box. The interface between the boxes is at the
top of the mixed layer (925 mb). The horizontal dimensions
of the boxes are 125 km in the along-front direction and
200 km, centered on the path of the stream, in the acrossfront direction. The fluxes shown in Figure 6 are the
differences between the SIMCS and the SIMNCS runs for
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Figure 3. Wind vectors from a QuikSCAT pass on 24 March 2000 plotted on an AVHRR-derived SST
field. Comparison of MM5 simulations and observations on 24 March 2000: (a) AVHRR-derived SST
field with scatterometer-derived wind vectors, black line – 200 m bathymetry, blue box – study area for
this SST-wind pair, (b) grid representation of the scatterometer wind, (c) mean SST normal to the selected
stream segment, (d) u component of equivalent neutral wind, UN10, (SIMCS - red line, SIMNC - cyan line
and the scatterometer observations - thick blue line), (e) v component of UN10, (f) UN10 and (g) wind curl.
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Figure 4. (top) Potential temperature and wind speed as a function of cross-frontal distance and
pressure for the cold-to-warm case at the end of the 15 hour simulation. Contour lines: potential
temperature on a contour interval of 0.25°K. Colors represent wind speed with color steps in 0.25 m s1.
(bottom) SST and wind direction at the surface.
24 March 2000. The total flux (not shown) across the
upwind interface of the free atmosphere box for the no
current, no SST front case (SIMNCS) is 56.81 108 kg s1 while that across the upwind interface of the
boundary layer box is 13.01 108 kg s1. The increase in
flux across the upwind interface of the boundary layer box
resulting from the introduction of the SST front, 0.4 108 kg s1, is the smallest change of those shown, with the
Figure 5. (top) Horizontal divergence and vertical velocity as a function of cross-frontal distance and
pressure for the cold-to-warm case at the end of the 15 hour simulation. Contour lines: vertical velocity in
cm s1. Colors represent horizontal wind divergence in 106 s1. (bottom) SST and wind direction at the
surface.
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Figure 6. Schematic showing changes in the vertical and horizontal (in the cross frontal directional
only) mass fluxes from the no current, no SST front model run (SIMNCS) to the complete simulation
(SIMCS) for the atmosphere divided into two regions: a free atmosphere region and a MABL region. All
fluxes are in units of kg s1.
exception of that through the top of the free atmosphere
box. By contrast, as a result of the acceleration of the air
across the front, the flux across the downwind interface
increases by 1.4 108 kg s1; i.e., the boundary layer box
is horizontally divergent. The extra mass leaving the boundary layer box in the horizontal is supplied from above, from
the free atmosphere box, as shown in the figure. The bulk of
this flux descends in the downwelling cell immediately
downwind of the front. This cell is approximately 100 km
wide with a mean downwelling speed of approximately
0.6 cm/s (Figure 5) resulting in a downward mass flux of
(105 m 1.25 105 m 6.0 103 m s1 1 kg m3)
0.75 108 kg s1. Of particular interest here is that the
perturbation of the flow in the free atmosphere box extends
well past the up- and downwind edges of the region shown
in Figure 6; the disturbance is felt more than 150 km
upstream of the front (not shown) and approximately
100 km downstream. Thus, the entire atmosphere in the
vicinity of the front is adjusting to the forcing associated
with the front, with the largest changes in wind speed near
the bottom of the MABL and the largest changes in mass
flux in the free atmosphere.
[36] It is interesting to compare our results with the 2D
simulation of Wai and Stage [1989], briefly described in
section 1. They examined a 1000 km domain with a
relatively weak SST front, 13°C in 350 km, in the middle.
Their simulation resulted in a secondary circulation pattern
with air rising over the warm water in the downwind portion
of the domain and descending over the cold water in the
upwind portion of the domain. The secondary circulation
direction in their study is opposite to that found in the MM5
simulation performed here. They concluded that the circulation is directly induced by underlying baroclinic thermal
forcing. However, our results suggest that the circulation
immediately above the front is more closely related to the
horizontal divergence in the wind field. Indeed, the diver-
gence is ultimately induced by the underlying oceanic
forcing. The difference between the results of these two
studies is conceivably due to either different physical
parameterizations within the MABL processes in the two
numerical models or different physical processes associated
with the different spatial scales.
[37] To gain further insight into the physical processes
that account for the relationship between the oceanic front
and the surface wind anomaly, residual fields were obtained
by subtracting the control run, SIMNCS, from the full
simulation, SIMCS. Cross sections of the residual fields
are plotted in Figure 7 showing perturbation pressure, wind
components u and v and wind speed. The resemblance of
the dipolar pattern of the perturbation pressure within the
MABL to that of the wind speed suggests the importance of
the perturbation pressure gradient in driving the wind. In the
MABL, the largest horizontal gradients in the perturbation
pressure are found over the front while the largest vertical
gradients are found over the warm water. The underlying
warm water induces a local negative perturbation pressure
near the sea surface. A positive perturbation pressure, albeit
of relatively smaller magnitude (0.06 mb compared with
0.12 mb near surface) is found above the negative
perturbation pressure. Corresponding to this low perturbation pressure is a positive near surface wind speed anomaly.
Since the mean surface wind blows from north to south, the
negative u and v anomalies near the surface (Figures 7c and
7d) mean that the surface wind is accelerated and rotates to
the right of the mean wind. Again, the upper portion of the
MABL clearly shows that the wind speed decreases (and
rotates to the left) over the warm water, a point that was also
discussed in the context of Figure 4.
4.1.1. Maintenance of the Pressure Gradient
[38] The fact that the PBL is in steady state, raises the
following two questions: (1) what gives rise to the dipolar
perturbation pressure pattern in the MABL over the warm
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Figure 7. SIMCS-SIMNCS: (a) perturbation pressure in Pa (0.01 mb), (b) wind speed anomalies in
m s1, (c) u component in m s1 and (d) v component in m s1.
water south of the Gulf Stream, and (2) how is the persistent
pressure gradient over the front maintained at different
levels. To address these questions, two vertical profiles
(Figure 8), one representing the atmosphere on the cold
side of the front, at +200 km, and the other representing the
warm side, at 200 km, were extracted from the full
simulation, SIMCS. The profiles of potential temperature,
moisture mixing ratio and wind speed all show enhanced
vertical mixing over the warm side of the front compared
with the cold side. Over the cold side, the potential
temperature profile (Figure 8a) indicates stable stratification. The small vertical variation of potential temperature
over the warm water results from the strong vertical mixing
from the sea surface up to about 950 mb. As air crosses the
front a mixed layer forms and deepens in response to the
heat fluxes and mixing by turbulence.
[39] The column integral of moisture mixing ratio over
the warm side of the front is larger than that over the cold
side (Figure 8b), indicating the injection of water vapor
from the ocean surface into the MABL. The moisture
mixing ratio increases from the cold side to the warm side
between 980 and 925 mb, from the middle to slightly above
the top of the mixed layer, and it decreases from 980 mb to
the ocean surface. Little vertical change of moisture mixing
ratio is found between 1000 mb and 960 mb. These features
indicate that strong vertical mixing throughout the entire
MABL efficiently redistributes the water vapor in it.
[40] Significant variation of wind speed profiles within
the MABL is also found between the cold and warm
sides of the front (Figure 8d). These profiles show an
increase in wind speed from cold to warm water near the
sea surface and a decrease in the middle of the MABL,
centered at 960 mb. A strong wind shear is also evident
over cold water where the bottom of the MABL is well
stratified.
[41] The above analysis reveals a reduced vertical variation in the moisture mixing ratio and less wind shear
over the warm water side of the front compared to the cold
side. Thus the effect of vertical turbulent mixing induced
by a positive SST anomaly is to homogenize the air mass
within the MABL. This corresponds to an upward mass
flux that tends to increase the density of the air in the
upper part of the MABL while reducing it near the surface.
At the same time, the mean density of the air in the
MABL – downstream of the SST front – decreases as it is
warmed.
[42] The combination of these two effects is clearly seen
in a cross section of potential air temperature (Figure 4).
The horizontal air temperature gradient near the surface
resulting from the differential SST heating is readily apparent in Figure 4 as is the decrease in air temperature above
980 mb downstream of the front. The near surface air
temperature gradient is not exactly collocated with the SST
gradient but centered about 50 km downwind of the path of
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Figure 8. (a) Potential temperature in °K, (b) moisture mixing ratio in g kg1, (c) perturbation pressure
anomaly in Pa (0.01 mb) and (d) wind speed in m s1.
the stream. This is due to the horizontal advection of the
warmed air by the mean surface wind. So, as the air moves
over the warmer SST downwind of the front it is warmed,
its density (inversely related to the virtual temperature)
decreases reducing the vertical stratification, and turbulent
mixing results. The turbulent mixing moves denser surface
air upward as it moves downwind while at the same time the
mean density of the air in the MABL decreases. These
changes in density result in an increase in the perturbation
pressure in the middle of the MABL and a decrease toward
the bottom of the MABL – the dipolar perturbation pressure
pattern in the MABL seen over the warm water south of the
Gulf Stream (Figure 7a).
4.1.2. Momentum Budget of the Mean Field
[43] In this and the following section, we examine the
momentum budget in the MABL in an effort to better
understand the physical processes contributing the modification of the MABL by the SST front and associated
current. In this section, we focus on the momentum field
for the run with the SST front and current, SIMCS. We refer
to this as the mean momentum field in the following. In the
next section the focus is on the residual momentum field.
The residual momentum at a given location is defined as the
momentum at that location for the SIMCS run minus the
momentum at the same location for the control run,
SIMNCS - no SST front and no ocean current.
[44] The horizontal momentum for the mixed layer and
free atmosphere can be written as follows:
Du
1 @p
@u0 w0
þ Du
¼
þ fv @z
Dt
r @x
ð2Þ
Dv
1 @p
@v0 w0
þ Dv
¼
fu @z
Dt
r @y
ð3Þ
where u and v are the horizontal wind components, p is
pressure, r is air density, f is the Coriolis parameter, u0 and
v0 are fluctuations of the horizontal wind components u and
v respectively, and w0 is the fluctuation associated with the
vertical component of the wind. Following the traditional
definition, the terms in the above equations represent, from
left to right, the acceleration of an air parcel, the pressure
gradient force, the Coriolis force, the momentum flux
divergence (referred to as the friction terms hereafter) and
horizontal mixing. The horizontal mixing terms, Du and Dv,
are very small at all levels compared with the other terms.
Therefore only the pressure gradient term, the Coriolis term
and the momentum flux divergence are discussed in the
following.
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Figure 9. The momentum associated with the three dominant terms at 30 m in the (a) cross-front, y, and
(b) along-front, x, directions. (c) SST. (d) The vector representation of the three dominant terms (left hand
column) and the vector sum of these terms, the local acceleration (right hand column).
[45] The PBL model, MRF, uses a non-local turbulence
scheme that incorporates the contribution of large-scale
eddies into the total flux. Momentum fluxes above the
surface layer are not routinely calculated by MM5, so the
turbulence statistics are calculated in this study from variables that are saved based on the formulation described by
Hong and Pan [1996].
[46] The pressure gradient, Coriolis force and momentum
flux divergence for the mean field are plotted for the lowest
model level, 30 m, (Figure 9) as a function of distance from
the path of the stream. The y component of the momentum
at 30 m (Figure 9a) is dominated by the pressure gradient
and momentum flux divergence, with only a slight contribution from the Coriolis term over warm water. The y
component of the pressure gradient is toward the south
and that of the momentum flux divergence is toward the
north. For the x component of the momentum (Figure 9b),
all three terms are important. Again friction is opposite to
the pressure gradient force. The cross-frontal acceleration of
the surface wind (thin black line in Figure 9a) is seen to
result from a combination of the pressure gradient and the
Coriolis terms, the pressure gradient term decreasing and
the Coriolis term increasing slightly compensating for part
of the decreased (more negative) pressure gradient. There is
little change in the friction term in the vicinity of the front.
[47] The vector form for the momentum budget of the
mean field near the surface (Figure 9d) shows the Coriolis
force and the friction term balancing the pressure gradient
force away from the front. In the free atmosphere (not
shown), the mean field is very nearly geostrophic; the
Coriolis term balances the pressure gradient term.
4.1.3. Momentum Budget of the Residual Field
[48] Figure 10 shows the horizontal momentum budget of
the residual field at 30 m. The residual momentum in the
cross-front, y, direction (Figure 10a) is dominated by the
perturbation pressure gradient and the Coriolis force. Over
the front, the perturbation pressure gradient is much larger
than the Coriolis force, resulting in an acceleration of the
wind in the negative y-direction (southward). Because
homogeneity is assumed in the x, along-front direction,
there is no frontally induced perturbation pressure gradient
in the x momentum equation. The x component of the
residual momentum (Figure 10b) is dominated by the
momentum flux divergence and the Coriolis force, with
the acceleration of the wind in the negative x-direction
(westward). In Figure 10d the dominant terms of the
residual momentum budget at 30 m are shown in vector
form. The vector sum of the residual momentum terms
(right column of Figure 10d) clearly shows the acceleration
of the residual horizontal wind above and immediately
downwind of the SST front. Away from the front, the net
acceleration is negligible. This indicates that the actual
surface wind is accelerated toward the south near the frontal
region primarily due to the perturbation pressure gradient.
Although the Coriolis force itself does not change the wind
speed, it does rotate the residual wind to the west as a result
of the westward acceleration applied to it.
[49] For the residual momentum in the free atmosphere
(Figure 11, 1350 m), there is no apparent acceleration of the
wind in the along-front direction. A southward (negative)
acceleration exists above the front as seen in the y momentum
budget, but this acceleration is small (order of 105 m s2),
as are all of the residual momentum terms, compared with that
near the surface (104 m s2). This acceleration leads to a
slight cross-frontal increase in the free atmosphere wind
speed, approximately 0.2 m s1, compared with a crossfrontal increase in the surface wind of order 2.0 m s1.
4.1.4. Momentum Budget Summary for the 24 March
2000 Case
[50] In the free stream the momentum balance is geostrophic, as expected, and the pressure field changes due to
the front are very small. Near the surface the mixing terms
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Figure 10. The residual momentum associated with the three dominant terms at 30 m in the (a) crossfront, y, and (b) along-front, x, directions. (c) SST. (d) The vector representation of the three dominant
terms (left hand column) and the vector sum of these terms, the local acceleration (right hand column).
become important and the pressure gradient is very similar
to that in the free stream, resulting in modified near surface
winds. This picture seems to be consistent with the 1D view,
i.e., modified mixing introduces modified surface wind
vectors. However, examination of the residual fields, shows
that, even a few hundred kilometers downstream of the front
where one might have expected the wind field to have
adjusted to a new equilibrium state, the pressure gradient
field near the surface is not the same as that in the free
stream, i.e., the system has not fully adjusted. Specifically,
the residual momentum balance is a three way balance with
the pressure gradient, the friction term and the Coriolis term
each contributing significantly — pressure modulation is as
important as friction modulation. Furthermore, in the vicinity (less than a few hundred kilometers downwind) of the
front the residual field is dominated by the residual pressure
gradient and the actual acceleration of the fluid, i.e., these
fields are far from being in a 1D balance.
[51] Park and Cornillon [2002] used a 1D model in their
analysis of the wind modulation due to an SST front. Their
assumption was that the MABL reaches a new equilibrium
shortly after crossing the front. It is clear from the results of
Figure 11. Same as Figure 10, except at 1350 m altitude (850 mb).
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the simulation presented here that this is not the case, that
the MABL is still in transition several hundred kilometers
from the front. This is clearly seen in Figure 9d. Upstream
of the front the Coriolis term and the friction term are
similar in magnitude and balance the pressure gradient term.
Downstream of the front the Coriolis term is substantially
larger than the friction term and the friction term continues
to decrease in magnitude to the southern most extreme of
the inner domain. This adjustment is also evident in the
perturbation terms shown in Figure 10.
4.2. The 5 September 1999 Warm-to-Cold Case
[52] The second case selected for detailed analysis,
5 September 1999 (Figure 12), differs from the first in three
fundamental ways: (1) The wind blew from warm to cold
water; (2) The wind blew obliquely to the stream, intersecting the front at approximately 130° (angle measured clockwise from the vector that is perpendicular to the stream
pointing from warm to cold to the direction from which
the wind was blowing); (3) The SST step across the edge of
the stream was much weaker, O(2°C), from 28°C on the
upwind side to 26°C on the downwind side.
[53] Figures 12d– 12g present the warm-to-cold model
results (SIMCS and SIMNC) together with scatterometerderived winds for the u and v components of the neutral
equivalent wind UN10, the magnitude of UN10 and wind curl,
all as a function of cross-frontal distance. The results of the
full MM5 simulation, SIMCS, capture the basic sense of the
change in QuikSCAT winds. Wind speeds from both
SIMCS and the scatterometer decrease from about
7.0 m s 1 over the upwind warm water to around
5.5 m s1 over the downwind cold water. Both SIMCS and
QuikSCAT show a maximum in wind speed near the front, a
feature that was not properly simulated by SIMNC (no sea
surface current forcing, cyan line in Figures 12d – 12g).
[54] Cross sections of potential temperature, q, and wind
speed produced from the full simulation, SIMCS, are
plotted as a function of horizontal distance and pressure
from the surface to 750 mb (Figure 13). While there is no
along-front variation in potential temperature and wind
speed at any given height in the simulation, considerable
variability in q and wind speed in the across-front direction
is evident throughout the MABL. Over the upwind warm
water the MABL is well mixed due to turbulent mixing.
This mixing results from the slightly warmer SST (301°K)
than the overlying air temperature (299.5°K) in the
upwind portion of the study area, i.e., the upwind marine
boundary layer is not in equilibrium with the underlying
water. This comes about because the SST far upwind of the
stream, to the southeast of the upwind portion of the study
area, appears to be a bit warmer than the SST in the study
area upwind of the stream (compare the SST outside and
southeast of the blue box in Figure 12a with the SST in the
blue box south of the stream). The largest horizontal
gradient in q is found across the front at the surface. The
top of the MABL, indicated by the sharp increase of q, is
slightly lower on the downwind, cold side of the front.
Within the MABL the contrast of the well mixed section
over the warm water to a more stratified section downstream
indicates that a stable internal boundary layer (IBL) has
formed near the surface over the cold water. Corresponding
to the formation of the stable IBL, there is a 20% reduction
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in the wind speed near the surface over the cold water. This
decrease in wind speed apprears to be caused by a constraint
on the transfer of momentum from the higher momentum
winds aloft due to the thermal stability of air near the
surface [Koracin and Rogers, 1990; Song et al., 2004].
This will be addressed in detail in the sections below in
which the momentum budget is analyzed.
[55] Cross sections of horizontal divergence and vertical
velocity are shown in Figure 14. The vertical velocity field
is characterized by features falling into two general groups,
one that is domain-wide and the other that appears to be
related to the SST front. The dominant domain-wide feature
consists of subsidence from the top of the region modeled,
750 mb, down to approximately 960 mb. The domain-wide
subsidence above the MABL results from an imbalance in
the upwind boundary condition and the air-sea temperature
difference upwind of the front. The upwind boundary
condition was obtained from the NCEP reanalysis which
assumes that the air is in equilibrium with the SST. This is
not the case here, so the MABL is continually undergoing
adjustment in the upwind region. This is what gives rise to
the strong turbulent mixing in the MABL in this region. The
large-scale subsidence, with maximum downward velocities
of 1.4 cm s1, evident above the mixed layer (900 mb) is
part of the adjustment process. It maintains an inversion cap
(evident in in situ vertical temperature profiles not shown
here) at the top of the MABL. Of more interest to this study
is the effect of the front on the vertical velocities: a
significant enhancement of the upwelling in the lower
MABL downstream of the front and a maximum in the
subsidence above the MABL over the front. A negative
(convergent) region in horizontal divergence forms over the
front. Associated with this convergence is an updraft with a
maximum vertical speed of 0.8 cm s1.
[56] Relative to the cold to warm case of 24 March 2000,
the range in horizontal divergence in this case is relatively
smaller over the study area, 9 106 s1 for this case
compared with 25 106 s1 for the 24 March 2000
case. This results in a smaller effect on the vertical velocity
in the upper MABL and above than was seen in the cold to
warm case.
[57] Residual fields for pressure, wind speed, and the u
and v components of the wind (Figure 15) were obtained, as
in the cold-to-warm case, by subtracting the corresponding
fields obtained in the control experiment, SIMNCS, from
the full simulation, SIMCS. The SST in the control run,
SIMNCS, consisted of a uniform field at the mean value of
the SST upwind of the SST front. Thus, the residual fields
represent the actual response of the entire atmosphere,
especially the MABL to the thermodynamic forcing of the
oceanic front and the cold water downwind of the front. As
discussed above, different mixing processes contribute to
the adjustment of the MABL to the SST front for the coldto-warm case and warm-to-cold case, but the spatial resemblance between the perturbation pressure (Figures 7a and
15a) and wind (Figures 7b – 7d and Figures 15b – 15d)
patterns within the MABL emphasizes the importance of
the perturbation pressure gradient in driving the air flow
across a front regardless of the wind direction.
[58] Associated with the IBL over cold water is a positive
perturbation pressure region (Figure 15a). A corresponding
negative perturbation pressure sits on top of the positive
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Figure 12. Wind vectors from a QuikSCAT pass on 5 September 1999 plotted on an AVHRR-derived
SST field. Comparison of MM5 simulations and observations on 5 September 1999: (a) AVHRR-derived
SST field with scatterometer-derived wind vectors, black line – 200 m bathymetry, blue box – study area
for this SST-wind pair, (b) grid representation of the scatterometer wind, (c) mean SST normal to the
selected stream segment, (d) u component of equivalent neutral wind, UN10, (SIMCS - red line, SIMNC cyan line and the scatterometer observations - thick blue line), (e) v component of UN10, (f) UN10, and
(g) wind curl.
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Figure 13. (top) Potential temperature and wind speed as a function of cross-frontal distance and
pressure for the warm-to-cold case at the end of the 15 hour simulation. Contour lines: potential
temperature on a contour interval of 0.25°K. Colors represent wind speed with color steps in 0.25 m s1.
(bottom) SST and wind direction at the surface.
Figure 14. (top) Horizontal divergence and vertical velocity as a function of cross-frontal distance and
pressure for the warm-to-cold case at the end of the 15 hour simulation. Contour lines: vertical velocity in
cm s1. Colors represent horizontal wind divergence in 106 s1. (bottom) SST and wind direction at the
surface.
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Figure 15. SIMCS-SIMNCS: (a) perturbation pressure in Pa (0.01 mb), (b) wind speed anomalies in
m s1, (c) u component in m s1 and (d) v component in m s1.
region at the upper part of MABL. This dipolar pattern is
also a feature of wind in the residual field (Figures 7b– 7d).
Positive perturbation pressure corresponds to a reduction in
wind speed and vice versa. The perturbation pressure varies
both across the front in the MABL and with height over the
downwind cold water. For the warm-to-cold case, the
magnitude of the perturbation pressure (0.03 mb) is
much smaller than that of the cold-to-warm case (0.1 mb),
with significantly less variability near the front (Figures 7a
and 15a).
[59] To explain the horizontal and vertical variation of
perturbation pressure, two vertical profiles of potential
temperature q, vapor mixing ratio, perturbation pressure,
and wind speed were plotted at two locations ±200 km away
from the path of the stream (Figure 16). These profiles are
very different from the cold-to-warm case (Figure 8). The
potential temperature profiles show little to no vertical
variations between 1000 mb and 940 mb, indicating that
for both the warm and cold water the MABL is well mixed.
The only significant difference is that q is cooled by
approximately 1°K over the cold water downwind of the
front. Potential temperature, q, increases with height very
close to the surface over the cold water, indicating a stable
IBL. Close to the surface, over the cold water, the air
contains more water vapor than the warm side as shown
in Figure 16b. The IBL inhibits the vertical transport of
water vapor to higher levels. It is interesting to note that
between 950 mb and 900 mb the water vapor content is less
over the cold side than over the warm side. This results from
the downward transport of relatively drier air (note the slope
of the vertical profile of the vapor mixing ratio in
Figure 16b) corresponding to the subsidence described
earlier (Figure 14). The downwelled drier air is then
advected by the mean wind from the warm to the cold side
of the front. The vertical barrier provided by the IBL
prevents the transport of moist air from the sea surface to
the free atmosphere, the only way that the moisture content
of the upper atmosphere could be increased.
[60] The above analysis enables us to propose a mechanism that explains the dipolar pattern of the perturbation
pressure over the cold water downwind of the front. The
relatively stable IBL, resulting from cooling over the cold
water, reduces the vertical exchange of air compared with
the control run, so cooler, denser air remains near the
surface downwind of the front resulting in a positive
perturbation pressure. By contrast, the inversion-penetrating
effect at the top of the MABL heats the air of the upper part
of the MABL helping to form the mid-MABL low perturbation pressure region over the cold water.
[61] The small crossfrontal variation of the potential
temperature in the atmosphere that accounts for most of
perturbation pressure gradient across the front close to the
surface is evident in Figure 13.
4.2.1. Momentum Budget of Mean Field
[62] The large-scale pressure gradient, Coriolis force, and
momentum flux divergence 30 m above the sea surface are
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Figure 16. (a) Potential temperature in °K, (b) vapor mixing ratio in g kg1, (c) perturbation pressure in
Pa (0.01 mb) and (d) wind speed in m s1.
plotted in Figure 17 for the mean field as a function of the
horizontal distance from the path of the stream. The y
component of momentum is dominated primarily by the
pressure gradient and the Coriolis force while all three terms
are important in the x momentum equation. A clear picture
of the horizontal variability of the three forces is shown
vectorially in Figure 17d. In response to the large-scale
pressure gradient that points toward the southwest
(Figures 17a and 17b), the large scale surface wind blows
from the southeast (Figure 12a). By contrast, the Coriolis
force and momentum flux divergence are directed roughly
opposite to the pressure gradient. The momentum flux
divergence is equivalent to a friction term and varies
substantially in the mean field. Over the cold water, the
momentum flux divergence is larger due to the IBL thus it
applies more friction on the air parcels. However, the flux
divergence does not change rapidly near the front because
there is upward vertical motion near the surface associated
with the convergent field discussed above. This convergent
field tends to reduce the momentum flux divergence over
the front. The momentum budget of the mean field in the
free atmosphere (figure not shown) indicates that at largescale the field is in geostrophic balance as was found in the
cold to warm case (24 March 2000).
4.2.2. Momentum Budget of Residual Field
[63] The pressure gradient, Coriolis force, and momentum flux divergence 30 m above the sea surface are plotted
in Figure 18 for the perturbation field as a function of the
horizontal distance from the path of the stream. From
Figure 18a it is evident that all three terms are significant
in the across-front, y, direction downstream of the front. In
response to the underlying SST cooling near the front and
over the cold water, a perturbation pressure gradient is set
up from cold water (north) toward warm water (south). The
direction of the perturbation pressure gradient is the same as
that in the cold-to-warm case, indicating that the perturbation pressure is thermally induced by the underlying SST; it
always points from cold to warm water regardless of the
direction of the mean surface wind. In the x direction, the
balance is between the residual momentum flux divergence
and the residual Coriolis force (Figure 18b). The southward
acceleration of air parcels begins near the front decelerating
the mean surface wind as it blows from southeast to
northwest. The deceleration acting on the southeast air flow
near the surface also rotates the wind to the left.
4.3. Effects of Sea Surface Current
[64] Studies by Cornillon and Park [2001], Kelly et al.
[2001] and Park et al. [2006] show that surface currents can
have an effect on the wind stress measured from scatterometers, and hence on UN10. In particular, the studies of
Cornillon and Park [2001] and Park et al. [2006], in
analyses of scatterometer winds over Gulf Stream rings,
have shown that on average, the component of the scatterometer derived wind in the direction of the current is
decreased by the speed of the current and the component
in the direction perpendicular to the current is unaffected.
The results of this study are consistent with these results. In
all ten match-ups (Figures 2, 3, and 12), simulations made
with sea surface current forcing (SIMCS – red) are in better
agreement with scatterometer observed winds than those
made without surface current forcing (SIMNC – cyan).
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Figure 17. The momentum budget of the three dominant terms at 30 m in the (a) cross-front, y, and
(b) along-front, x, directions. (c) SST. (d) The vector representation of the three dominant terms (left hand
column) and the vector sum of these terms, the local acceleration (right hand column).
[65] Furthermore, although difficult to see in these
figures, the y component of UN10 shows little to no variation
due to the current while the profile of the x component
across the front, obtained by subtracting SIMCS
from SIMNC, is very similar to the sea surface current
incorporated into MM5. This suggests that the thermal
effect of an SST front and the dynamical effect of a surface
current on the wind are nearly additive.
[66] Figure 19 is obtained by subtracting SIMNC from
SIMCS to show the effect of the sea surface current on
the atmosphere for the 24 March 2000 case. Cross
sections of the u component of the wind, vapor mixing
ratio, temperature, and vorticity in the residual field are
plotted. The u component of the wind plotted here
(Figure 19a) is the actual wind calculated in MM5, not
the wind calculated from MM5 results for comparison
Figure 18. The residual momentum budget of the three dominant terms at 30 m in the (a) cross-front, y,
and (b) along-front, x, directions. (c) SST. (d) The vector representation of the three dominant terms (left
hand column) and the vector sum of these terms, the local acceleration (right hand column).
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Figure 19. SIMCS - SIMNC, 24 March 2000 case: (a) u component of wind, m s1, (b) relative
vorticity, 105 s1, (c) vapor mixing ratio, g kg1 and (d) temperature, °C.
with the scatterometer winds (Figures 2, 3, and 12). The
positive u component anomaly observed above the front
indicates that the sea surface current applies friction to
the air, tending to ‘‘drag’’ it in the direction of the
current. This increase in the u component of the wind
is mixed upward and advected downwind, hence the
upward sloping wind anomaly downwind of the front.
The sea surface current applies a positive torque on the
near surface air to the left of the current peak leading
to the weak positive anomaly in the relative vorticity
upwind of the front (Figure 19b). Similarly, the current
induced shear results in negative vorticity to the right of
the current peak. However the effect of the surface
current on mixing ratio and temperature is insignificant.
The largest surface current induced variation of temperature occurs near the surface and is on the order 3 102 °C (Figure 19d) compared with the thermodynamic
induced changes in the same general area of order 2°C
(Figure 8a). The largest surface current induced variations
of the mixing ratio occur downwind of the current at
925 mb and are order 102 g kg1 (Figure 19c)
compared with thermodynamically induced changes in
the same general area of order 0.5 g kg1 (Figure 8b).
[67] The residual SIMCS minus SIMNC fields for the
5 September 1999 case (Figures 20a – 20d) show a pattern
for the u component of wind that is similar to that of the
24 March 2000 case, but with a rapid attenuation with
height. This is probably due to the formation of the IBL
over the cold water that restricts the vertical propagation of
wind anomalies.
[68] The above analysis shows that the current, in addition to affecting the wind stress as measured by scatterometers, can alter the dynamical structure of the MABL.
However, the current has little effect on the thermal structure of the MABL such as the mixing ratio and air
temperature. The thermal structure is primarily affected by
the discontinuity in SST that results in differential heating of
the MABL. The fact that the current does not lead to
thermodynamic changes in the atmosphere is what led to
the observation made above that the dynamical effect of a
sea surface current on the MABL and the thermal effect of
an SST gradient on the MABL are very nearly additive.
5. Discussion and Conclusions
[69] This paper presents an analysis of the surface wind
response to an SST front based on a combination of
scatterometer (NSCAT and QuikSCAT) derived wind fields
and AVHRR derived SST fields in conjunction with PSUNCAR MM5 simulations of the observed conditions. Our
results show that the perturbation pressure gradient is the
main force modifying the air flow across the front. When
the wind blows from cold to warm water across an SST
front, strong mixing homogenizes the entire MABL downwind of the front. This, together with an increase in the
mean temperature of the MABL, leads to low and high
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Figure 20. SIMCS - SIMNC, 5 September 1999 case: (a) u component of wind, m s1. (b) relative
vorticity, 105 s1, (c) vapor mixing ratio, g kg1 and (d) temperature, °C.
perturbation pressure anomalies from the ocean surface to
the upper part of the MABL. Specifically, there is a negative
perturbation pressure anomaly near the surface over, and
immediately downwind of, the front and a positive perturbation pressure anomaly aloft within the MABL. The
negative perturbation pressure anomaly accelerates the
surface wind. The positive perturbation pressure anomaly
decelerates the wind aloft. The acceleration of the wind near
the surface results in a divergence over the front. This
divergence draws air into the MABL from the free atmosphere giving rise to a modification of the free atmosphere
that extends well upwind and downwind of the front. Large
scale subsidence (negative vertical velocity) in the free
atmosphere maintains the sharp property gradient at the
top of the MABL.
[70] When the wind blows from warm to cold, the
underlying cold water tends to stabilize the MABL. In the
case examined here, a stable internal boundary layer (IBL)
was formed downwind of the front that reduced the vertical
exchange of air in the MABL. This, together with a decrease
in the mean temperature of the MABL downwind of the
front, resulted in a high and a low perturbation pressure
anomaly between the ocean surface and the upper part of the
MABL, a reverse dipolar pattern of perturbation pressure
downwind of the front in contrast with what was found in
the cold to warm case. As a result of the increase in
pressure, the wind near the surface was decelerated on the
leading edge of the boundary layer resulting in a region of
horizontal convergence.
[71] The mechanism of wind adjustment discussed above
indicates that the formation of a perturbation pressure
gradient is directly induced by the underlying SST field.
The SST front sets up a perturbation pressure gradient that
points from cold to warm water regardless of the direction
of mean wind. It is this perturbation pressure gradient,
together with the Coriolis force, that modifies the air flow
in the presence of an SST front.
[72] The analysis presented above contrasts with the three
mechanisms [Wallace et al., 1989; Wai and Stage, 1989;
Lindzen and Nigam, 1987] discussed in the Introduction
that have been proposed to explain the positive correlation
between the underlying SST and wind speed. The first two
mechanisms (Lindzen and Nigam [1987] and Wallace et al.
[1989]) were proposed to explain variations of the crossequatorial flow in response to tropical instability waves in
the equatorial Pacific. There are two major assumptions
made by Lindzen and Nigam that can be compared with this
study. First, they assumed that the MABL is well mixed on
either side of the SST front and thus there is no vertical
variation in potential temperature. This assumption precludes a vertical variation of the pressure gradient over
the front. Our study shows a clear variation in the pressure
gradient with height over the front. Second, they impose a
rigid lid at the top of the mixed layer and there are no
interactions between the MABL and the free atmosphere.
Our study suggests that the dipolar perturbation pressure
pattern results from the enhanced vertical turbulent mixing
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due to surface heating as well as the interaction between the
MABL and the free atmosphere.
[73] Wallace et al. [1989] argued that, if the tropical lower
troposphere is relatively well mixed in the vertical, one
would expect horizontal gradients of temperature (or virtual
temperature) beneath the trade inversion (usually between
the 850 and 700 mb levels) to be similar to their surface
values [Lindzen and Nigam, 1987]. However, they found
the surface winds were strongest over warm water to the
north of the strongest SST gradient. From this they concluded that modification of boundary layer shear rather than
changes in the atmospheric sea level pressure was the
dominant process affecting surface winds. By contrast, our
simulations show that advection by the mean wind is what
determines the phasing between wind speed and SST
gradient in the vicinity of the front and that sea level
pressure is the dominant term resulting in the modification
of the surface winds by an SST front.
[74] Our simulation also contrasts with the results of Wai
and Stage [1989] described in section 4. Wai and Stage
[1989] argued that the circulation pattern that they observed
in their simulations was due to an underlying baroclinic
thermal forcing. As noted previously, the sense of the
circulation pattern that they obtained was opposite to the
one that we obtained for a similar configuration, wind
blowing from cold to warm water.
[75] The reason that previous theories tended to discount
the importance of cross-frontal pressure gradients in modifying the wind field is that most observational studies failed
to detect a change in the pressure field in the vicinity of a
front [e.g., Hashizume et al., 2002; Tokinaga et al., 2005].
However, the failure of most observational studies to detect
a cross-frontal pressure gradient is not inconsistent with our
results. The perturbation pressure anomalies observed in our
simulations were generally smaller than 10 Pa (0.1 mb),
which is near the detection limit of current mercurial
barometers, but as, demonstrated in the previous section,
perturbation pressure anomalies on this order can lead to
substantial changes in the wind speed. In fact, Cronin et al.
[2003] reported an SST induced variation in pressure in
tropical instability waves which is consistent with our
model results.
[76] In addition to explaining the cause in the change of
wind speed across an SST front, our results suggest that a
1D boundary layer model, applied locally as in Park and
Cornillon [2002], cannot explain the wind modulation due
to an SST front, even a few hundred kilometers downstream
from the front, because the pressure field is still in transition
to the new equilibrium state and it is therefore different from
that in the free stream. In other words, the residual pressure
gradient inside the boundary layer plays an important role
over a large spatial scale, almost as important as the
modulated mixing effect.
[77] Finally, we note that our results help explain (manuscript in preparation) the large gradients in long term
averages of wind stress curl observed in the vicinity of
strong boundary currents Chelton et al. [2004].
[78] Acknowledgments. This study was performed with support from
the National Aeronautics and Space Administration (grant NAS5-32965)
via Oregon State University as part of the SeaWinds program. Salary
support for P. Cornillon was provided by the State of Rhode Island and
Providence Plantations. The scatterometer wind data were obtained from
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the Jet Propulsion Laboratory (JPL) via OPeNDAP. The Mesoscale Model
(MM5) was obtained from the National Center for Atmospheric Research.
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