Coastal Upwelling.
Coastal upwelling brings together many of the physical oceanographic processes that
we’ve learned this year. —and thus is perhaps an excellent ending to the class. These
include:
Wind stress
Ekman Transport
Ekman Depth
Rossby Radius
Continuity equation
Coastal upwelling occurs
near a coast when the wind
stress has a component in the along-shore direction with the shore to its left (in the
Northern Hemisphere). In the case of New Jersey this would be a southerly wind (wind
blowing to the north), while in California upwelling occurs when the wind blows from
the north to the south. Note that in both cases upwelling occurs when the wind blows in a
direction opposite to the direction of Kelvin wave propagation.
Upwelling occurs because the wind-driven
Ekman transport in the surface layer is away
from the coast—and this off-shore transport
of surface water is replenished by an onshore
transport of bottom waters. If the surface
water is warm and the lower layer cold—as
often is the case because cold water is
heavier—then the upwelled water will be
cold. Similarly, as often is the case, if the
lower layer is nutrient rich then the upwelling
will bring nutrient rich water up into the
euphotic zone and drive a phytoplankton
bloom. This is clearly seen in satellite
imagery of an upwelling event along the US
west coast that followed a period of persistent
northerly winds.
Chlorophyll- a
Temperature
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Similarly, along the New Jersey
Shore southerly winds generate
upwelling. In the summer months
this can drop beach temperature to
below 60 degrees in the middle of
July!.
In fact fact this past year the
upwelling was so persistent along
the US east coast that it attracted
the attention of the national media.
Interaction between bathymetry
and tends to focus the upwelling in
upwelling centers—and this give
rise to the along-shore variability
seen in the figure to the right. THe
increased plankton growth
associated with this upwelling—
and concentration of the associated
organic matter appears to produce
regions of recurrent low dissolved
oxygen along the New Jersey
Coast.
If the wind began to blow to the
north along the New Jersey coast a
10 m/s how long would it take for
the thermocline to upwell?
Grey indicates
recurrent regions
of low dissolved
oxygen
We know that the wind stress,  associated with this wind will be Cd W2, where Cd is a
drag coefficient (let’s assume it is 1x10-3) and W is the wind speed. This corresponds to
a surface wind stress of 0.1 Pa/m2.
We know that the Ekman Transport is
2

 fu
H
Where H is the Ekman depth.
In a two layer system—when the surface layer is shallow—we can often assume that the
Ekman depth is equal to the depth of the surface mixed layer. This occurs because the
strong stratification in the thermocline suppresses mixing and thus the wind stress is
confined to the surface layer.
So the off-shore transport in the upper layer is

, with f=10-4 1/s and kg/m3 the
f
off shore transport is 1 m2/s.
To calculate the vertical velocity we need to use the continuity equation—which if we
neglect along shore variability (i.e. dv/dy =0) – can be written as:
u w

0
x z
Integrating this vertically over the surface layer ( which we assume is the Ekman Layer)
we get
0

u
z 
x
H
0

w
z  0
z
H
0

uz  w 0  w H  0
x

H
Where wH is the vertical velocity at the base of the upper layer, and w0 is the vertical
velocity of the sea surface. If we integrate the first term over the surface layer—this is

simply the cross-shelf transport, which we know is
and if assume that the vertical
f
velocity at the surface is zero (this is not quite right—for the upwelling to occur requires
the surface layer to drop down a cm or two and this drives the on-shore flow at depth—
but most of the thinning of the upper layer is due to the upward vertical motion of the
thermocline—so the assumption is valid) we get:
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   
 w
x  f 
At the coast we know that the cross-shore transport is zero—and it increases to the

Ekman transport--- approximately on internal Rossby Radius, R, from the coast
f

than the cross shore divergence in the upper layer is simply
which equals the
Rf
upwelling velocity.
Rossby Radius
R is approximately 10 km and plugging rest of the values given above we f nind that the
upwelling velocity is 0.1 mm/s or approximately 8 meters per day. Given that the surface
mixed layer on the New Jersey shelf is typically 10 meter deep—upwelling will generally
occur if a 10 m/s wind were to blow persistently for two days.
Note that the cross-shore Ekman flow appears in the along-shelf momentum equation
because it includes the Coriolis term that involves the cross-shelf flow. What drives the
bottom layer on-shore is the tilting down of the surface layer. This cross-shelf slope is
expressed in the cross-shelf momentum equation.
v
 
g

 fu
t
x z
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In the interior geostropic flow would then be up the coast. In the cross shl
HOwever, other complications are:
Initial condition—for example if the cold water is far off-shore then it will take more
time for upwelling to occur for the cold water must first move onshore.
Also this 2-dimensional model does not fully describe the upwelling that we observe
because of complications arising due to bathymetry and spatial and temporal variability
in the wind field.
Example include the one I showed in class where off-shore transport in the surface layer
tends to agree with the Ekman Theory—but the on-shore transport in the lower layer does
not—suggesting that upwelling is a fully 3-dimensonal process that cannot be fully
described by this simple 2-dimenonal model.
Opposite of Upwelling is Downwelling.
Ask in class what sea-surface would do
What would the cross-shore circulation look like.
The above discussion is on the along-shore momentum balance, which is Ekman. But in
the cross-shore momentum balance (assuming winds are strongly along shore) the tilting
sea-surface, and tiltling isopycnals, develop a cross-shore pressure gradient which is
balanced by a strong along shore flows that are in geo-strophic balance.
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