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and forecasting; the speed of storm advance is sensitive to
degree of surface friction at the ocean-atmosphere boundary
(Section 6.2.4).
An important roughness feature on the sea surface is
water waves. These occur on a variety of scales. The waves
take the form of smooth sinusoids through to breaking
waves (whitecaps). Wave formation from a flat sea surface
is due to velocity and pressure fluctuations in the wind set
up during wind-wave coupling over initially small waves
that then grow to equilibrium. They are broadly analogous
to the Kelvin-Helmholz waves described in Section 4.9.
The growing waves then strongly influence momentum,
energy, and mass exchanges at the interface. Any overall
w
ind velocity,
u
, is the sum of a long-term mean velocity,
, a turbulent fluctuation from the mean,
u
u
, and a peri-
odic velocity associated with the waves, so that
. The same applies to Bernoulli pressure
fluctuations induced by the wind: when the wave pressure
is less on the lee side than the stoss side of the wave then
energy is transferred from air to wave and the wave grows.
We might ask ourselves how far upward the log layer
extends in natural ABLs which may be in motion for many
hundreds of meters to kilometers upward above any meas-
uring platform. The answer comes from data collected in
connection with hurricane studies in which radio sondes,
remotely-tracked by Global Positioning System (GPS)
receivers, were released into the ABL from aircraft
(Fig. 6.15). The frictional influence of the flow boundary
(ocean water in this case) extends upward as a log layer
c
.200 m. Maximum flow velocities were reached at about
500 m, gradually weakening upward to a height of 3 km.
Surprisingly, the most energetic storm winds lead to a reduc-
tion of surface roughness friction, despite the production of
larger surface waves. This is thought to be the result of the
production of abundant surface foam, which somehow acts as
a
drag-reduction agent
. The reduced friction enables tropical
cyclonic storms to move faster than predicted from conven-
tional considerations of boundary drag. Such momentum-
exchange processes obviously lead to gaseous exchange
from the ABL into the ocean, but also vice versa, in surface
ocean layer gases from planktonic photosynthesis. Also, both
sensible and latent heat are exchanged via radiation gains
and losses, evaporation of surface waters due to forced
convection, and from the condensation of water vapor.
˜
u
u
u
u
˜
Fig. 6.13
The stable FLIP platform for measurement of wind velocity/
boundary layer data in the offshore environment 50 km west of
Monterey, California.
t
=
C
d
ru
*
2
du
Individual time-mean
velocity measured
by flow meter
dz
u = k log z
The shear velocity,
u
*
, is
proportional to the velocity
gradient,
δ
u
/
δ
z
.
Intercept on ordinate
defines roughness length
Mean flow velocity,
u
6.2.2 Dynamic ocean topography: Atmospheric
wind forcing of surface ocean currents and circulation
Fig. 6.14
Schematic of typical ABL velocity distribution with height
to show the logarithmic relationship and the computation of shear
velocity and hence shear stress,
The low resistance of ocean water to surface shear by
the blowing wind leads to a net mean motion of the
, as a measure of flow turbulent
momentum exchange.
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