Download To change the ocean water density we can: provide heating/cooling

OCEAN RESPONSE TO AIR-SEA FLUXES
The main mechanism: mechanical and
convective mixing in the surface layer.
Both work to provide a deepening of the
upper ocean mixed layer.
Temporal seasonal evolution
of the surface mixed layer
Oceanic and atmospheric mixed
layers in contact
2 non-dimensional
parameters:
Simple mixed layer model:
T
Ts
h
TH
z
H
Integration within surface mixed layer gives:
Vertical heat exchange equation:
Integration within surface
mixed layer gives:
Ts (t )  T ( z, t )

,
Ts (t )  TH
z  h(t )

H  h(t )
T
Q

t
z
Ts
h
 Q
t
Surface
net flux
Ts
h
 Q
t
Integration from h to H:

h
T
(
z
,
t
)
dz

T

Q
s

t h
t
H
This integral can be written as
H
 T ( z, t )dz  (T
H
 Ts )  Ts ( H  h)
h
Ts  T ( z, t )
 
dz   ( H  h),
Ts  TH
h
H
where
:
1
   d
0
Ocean circulation characteristics from
sea-air interaction parameters
Meridional Transport of Heat (MHT): OCEAN VIEW:
The total MHT can be represented as the sum of the advective and diffusive
transports of heat:
MHT   C p  dl  vT dz   C p  dl 
L
z
L
z
T
k
dz
y
Advective MHT can be decomposed into an overturning component which
corresponds to the transport of the zonally averaged circulation, and a gyre
component which corresponds to the transport by horizontal gyres:
MHT   C p  dl  vT dz   C p  dl  v T dz   C p  dl  v' T ' dz
L
z
L
z
L
z
where  is the zonal averaging operator and prime stands for the variations of
temperature T and meridional component of velocity v around the zonal
mean.
If we know surface net heat flux, can
we say something about the MHT?
M1=M0+Q1
M2=M1+Q2=
M0+Q1+Q2
Q2
M2
S
Q1
M0
M1
N
Estimates of MHT can be obtained from the integration of
Qnet = SW-LW-Qh-Qe
over the ocean surface north of the latitude considered:

MHT   d  Hd  M 0
0
where Mo is a boundary condition at the north which should be taken from alternative
source (observations) or set to zero.
Being computed from a
balanced surface flux,
global MHT should
converge to zero at the
north and at the south
Meridional heat transport shows basic features of the ocean general
circulation, in particular it marks a specific role of the Atlantic Ocean,
which indicates the northward MHT at all latitudes in both Hemispheres.
Ocean heat fluxes are influenced by the uncertainties (observational, sampling,
etc.), which impact on the zonal heat balance and propagate quickly to MHT
through the integration: MHT, being physically a very important parameter,
remains not very effective measure of the reliability of surface flux fields from a
methodological viewpoint. Uncertainty of 10 W/m2 for the Atlantic Ocean results
in 0.5 PW error in MHT at 20S
MHT, 10**14 W
25
20
30
15
20
10
10
Closure of surface heat
balance is not achieved in
most climatologies of seaair fluxes:
SOC (Josey at al. 1999):
-30 W/m2, negative MHT at
12N in the Atlantic
5
0
10
scheme-to-scheme variation
variation due to variable
-5
corrections
VOS - ensemble average
20
-10
UWM (Da Silva et al. 1994) :
- 30 W/m2
NCEP/NCAR
NCEP/NCAR undersampled
30
SOC climatology
UWM climatology
IFM climatology
-15
ocean
inverse
model
Ganachaud (1999)
-20
-20
-10
0
10
20
30
latitude
40
50
60
70
Water mass formation characteristics from surface fluxes
Ocean waters are characterized by temperature and salinity. Equation
of state links density with these characteristics.
Highest
density
Smallest
density
Surface density in the North Atlantic
Typical density of ocean waters are from 1023 to 1028 kg/m3. For practical
reasons ocean densities are estimated in relative units:
=-1000
Surface water
dzdS
Bottom
water
T,S – diagram is the
ocean analog of
P,T - diagram
Volumetric
T,S - diagram
To change the ocean water density we can:
 provide heating/cooling by surface net heat flux (W/m2)
 provide precipitation/evaporation (m3/sec)
But how to know how many kilograms we added (or
extracted) to (from) the ocean by the joint application of
these two processes?
The density flux (in fact a virtual mass flux since it has the unit of
kgm2s-1) at the sea surface has been derived as:

( E  P) S
F
Qnet   0 
CP
1 s
This flux is applied as a
surface boundary condition to
ocean models
where CP is specific heat of sea water at constant pressure, 0 is a reference
density of sea water, s is salinity in portions of unity,  and  are the thermal
expansion and haline contraction coefficients:
 = /T,
 = /S
North Atlantic surface density flux in kg/m2s, computed
from the net flux and evaporation minus precipitation
Water mass transformation by surface density flux
Integration in space and time of the density flux gives the
transformation rate of waters at given density, F(ρ), which
represents the time average over a period T (generally one
year) of the density change of a unit water volume of density ρ
which results from the action of the atmospheric forcing:
1
F     dt  f      d
TT 
where the  function samples the density flux f over the area
where waters of density  are outcroping within the integration
area . This quantity has a unit of kgs-1 and can be scaled
with the unit density to be expressed in m3s-1 or Sv. In this form
the transformation rate shows the volume of water of density
0 which is transformed, during a given period, into higher
densities (F(0) > 0) or lower densities (F(0) < 0).

1

 dt  f      d


1E+007
haline
transformation
0E+000
thermal
transformation
-1E+007
-2E+007
22
24
26
sigma-0, kg/m**3
Surface tropical
waters
STMW
LSW
1E+007
0E+000
transformation
captured by
the outcrop
MLD>150m
-1E+007
transformation
diagnozed by
model
-2E+007
20
28
LSW
transformation
diagnozed
by original
sea-air flux
2E+007
transformation, m**3/sec
F   
total transformation
2E+007
transformation, m**3/sec
Трансфрмация вод
на поверхности
STMW
20
22
24
26
sigma-0, kg/m**3
28
Thermal versus haline
contributions
3E+007
CLIPPER-1/6, model
ATL6
ATLN1
ATLN6
transformation, m**3/sec
2E+007
haline
1E+007
0E+000
ATL1
-1E+007
thermal
-2E+007
-3E+007
20
22
24
26
sigma-0
28
30
Water mass formation:
The formation rate was defined by the differentiation of the transformation
rate, as the gradient of F(T,S) orthogonal to isopycnals:
M(T,S) = -[F(T,S)  (T,S)] / ,  = (S0S, T0T).
Strong formation of STMW.
Strong formation of LSW.
Overturning induced
by the water mass
transformation
 ( ,  ) 

ATL1
0
 d  f ( ,  ) H ( )d ,

1,  surf  
H 
0,  surf  
ATL6
The role of surface fluxes in forming ocean circulation
-100
70
-90
-80
-70
-60
-50
-40
-30
-20
-10
0
10
20
60
50
30
70
60
Incorrect position
50
40
40
30
30
20
20
10
10
0
0
Correct
position
-10
-100
-90
-80
-70
-60
-50
-10
-40
-30
-20
-10
0
10
20
30
Importance of the correct location of the Gulf stream in relation to
air-sea fluxes