slide1 n.
Download
Skip this Video
Loading SlideShow in 5 Seconds..
Ocean Circulation, dynamics and thermodynamics…. PowerPoint Presentation
Download Presentation
Ocean Circulation, dynamics and thermodynamics….

Loading in 2 Seconds...

play fullscreen
1 / 70

Ocean Circulation, dynamics and thermodynamics…. - PowerPoint PPT Presentation


  • 133 Views
  • Uploaded on

OCEAN 587. Ocean Circulation, dynamics and thermodynamics…. Equation of state for seawater General T / S properties of the upper ocean Upper ocean circulation Heat balance of the upper ocean Deep circulation.  =  ( S , T , p , ……). [Determined from laboratory experiments].

loader
I am the owner, or an agent authorized to act on behalf of the owner, of the copyrighted work described.
capcha
Download Presentation

PowerPoint Slideshow about 'Ocean Circulation, dynamics and thermodynamics….' - kanan


An Image/Link below is provided (as is) to download presentation

Download Policy: Content on the Website is provided to you AS IS for your information and personal use and may not be sold / licensed / shared on other websites without getting consent from its author.While downloading, if for some reason you are not able to download a presentation, the publisher may have deleted the file from their server.


- - - - - - - - - - - - - - - - - - - - - - - - - - E N D - - - - - - - - - - - - - - - - - - - - - - - - - -
Presentation Transcript
slide1

OCEAN 587

Ocean Circulation, dynamics and thermodynamics….

Equation of state for seawater

General T/S properties of the upper ocean

Upper ocean circulation

Heat balance of the upper ocean

Deep circulation

slide2

 =  (S, T, p, ……)

[Determined from laboratory experiments]

[ S = salinity; T = temperature; p = pressure ]

In order to use this relation to infer  it is necessary to clearly define the meanings of S, T, and p.

slide3

Salinity….

Why is the ocean salty? [ A great question ]

Hydrological cycle  fresh water 

 precipitation  continental runoff 

 evaporation  salinity

[seems unlikely]

Geochemical implications: Earth’s crust is rich in aluminum; rivers are rich in carbonates; oceans are rich in halides and sulfides.

Seawater is a complex chemical system that contains traces of nearly all naturally occurring elements. The dissolved part of seawater is about 78% NaCl by mass.

slide4

The chemical composition of seawater….

Amazingly, the proportions of the major constituents of seawater are nearly constant everywhere.

slide5

This constancy of proportions means that to a good approximation, only one component of seawater needs to be measured, and all of the other can be inferred from it. (what are the limitations?)

Instead of characterizing a seawater sample by knowing all of its chemical components, we can characterize the sample by a single parameter. We shall call this parameter the salinity.

Formal, traditional definition of salinity:

“The total amount of solid materials in grams contained in one kilogram of seawater when all the carbonate has been converted to oxide, the bromine and iodine replaced by chlorine, and all organic matter completely oxidized.”

slide6

34.60

34.70

34.72

How well do we need to measure S? The section of S along 43 S in the S. Pacific shows that there is considerable detail at signal levels  0.01 PSU.

slide7

Temperature (continued)….

Seawater is slightly compressible. What are the implications of this?

[warming]

[cooling]

Adiabatic changes in temperature: changes in temperature when there are no changes in heat.

slide8

Temperature (continued)….

Potential temperature (): the temperature that a seawater parcel would have if it were raised adiabatically to the sea surface. [Note:   T ]

The temperature change in adiabatic conditions is due to compressibility alone, since seawater is slightly compressible…..this can get confusing.

Given 2 parcels, which is heavier?

Care must be taken, since they are at different pressures.

Solution: refer both to a common pressure.

slide9

Temperature (continued)….

T

In situ temperature T increases as depth inceases.

Potential temperature  does not increase with depth.

slide10

Density….

 =  (S, T, p) is the symbolic equation of state.

Over the water column,  varies by a few per cent.

surf  1.021 g/cm3 ; bot  1.071 g/cm3 [typical values]

In the upper ocean,    (S, T) .

In the deep sea,    (p) .

Define the parameter  as a more useful representation of density:

 = (  1)  1000 (cgs) ;  =   1000 (mks)

slide11

There are several different types of density….

  •  = (S, T, p) (in situ sigma)
  • t = (S, T, 0) (“sigma-tee”)
  •  = (S, , 0) (“sigma-theta”)
  • note:
  • (4) po = (S, po, po) [po = (S, T, po) ]

[potential temperature]

[  = adiabatic lapse rate = T/p ]

slide12

Typical profiles in the ocean….

[ from Ocean Station P, 50N, 145W ]

t

T

S

slide14

Equation of state….

Step 1, fresh water equation

Step 2, add salinity at zero pressure

Step 3, full equation with T, S, p

[a 29 term polynomial]

slide15

Equation of state….

The equation of state for seawater is nonlinear. This has major implications to ocean mixing.

slide16

Ocean circulation: surface currents….

[notice E/W asymmetry]

[ 0-500 dbar dynamic ht; maximum range ~ 2 m]

slide19

The Heat Balance of the Earth (and the ocean’s role)….

Solar heating is ultimately the source of all energy at the Earth’s surface, except for tides and geothermal heating.

The atmosphere and ocean exchange heat and energy in a complicated set of feedbacks.

Recall the 2nd law of thermodynamics….

Heat is conserved.

slide20

Simple heat transfer….

2 adjacent blocks

Assume T1 > T2

Let Tf be the final temperature

Heat lost from block 1 = m1c1(T1Tf) = H1

Heat gained by block 2 = m2c2(Tf T2) = H2

H1 = H2

Tf = (m1c1T1 + m2c2T2)/(m1c1 + m2c2)

slide21

Atmosphere-ocean heat transfer….

Suppose we change T in the lower 10 km of the atmosphere by 10 C (a massive change!).

Suppose we put all of this heat into the upper 100 m of the ocean.

How much would T in the upper 100 m of the ocean change?

Atmosphere

Ocean

(.0013) (10) (0.25) (10)

(1) (0.1) (1) (?)

[small! the ocean acts as a buffer to the atmosphere]

result:

slide22

The heat content in the upper 300 m of the ocean has varied over the past 50 years, generally increasing…where has this heat come from?

[heating rate  0.6 C/century]

[from Levitus]

slide23

The concept of a flux….

area A

amount ofstuff S per unit time

The flux of S is defined as the amount of S crossing the area A per unit time; a flux is just (amount)/(area  time).

slide24

The global heat balance (schematic)….

atmosphere

ocean/land/cryosphere/biosphere

slide25

The heat balance of the ocean….

QT = QS + QB + QH + QE + QV + QG

where

QT = net heat input to the Earth ( 0 watts/m2)

QS = direct solar input ( +150 watts/m2)

QB = black body radiation (50 watts/m2)

QH = sensible heat loss (10 watts/m2)

QE = evaporative heat loss (90 watts/m2)

QV = advective heat transport (0 watts/m2)

QG = geothermal heating ( +0.01 watts/m2)

[note: there are errors and uncertainties associated with each of these estimates]

slide27

QB, black body radiation….

[Sun 0.7 m]

[Earth 15 m]

Wien’s Law: maxT = A (a constant)

Sun: high T (short ); Earth: low T (long )

slide28

QB, black body radiation….

The amount of heat radiated by a black-body of temperature T can be estimated from the Stefan-Boltzmann relation as

QB = T4 ,

where  is the Stefan-Boltzmann constant.

For the case where there is no atmosphere and the heat balance is in equilibrium, on the Earth we would have

QB = So (1-)/4 = T4T = (222/)1/4  278 K .

This is too cold for the globally averaged T (actually about 300 K).

[  = 5.67  10-8 W/(m2 K4) ]

slide30

The concept of flux-gradient diffusion….

[note: this is generally true in the molecular case only]

The flux is proportional to the gradient of the concentration, with the constant of proportionality being the diffusivity .

slide31

QH, conductive (or sensible) heat flux….

Since the ocean is generally at a different temperature than the atmosphere, there will be a conduction of heat between them (recall the two blocks).

Fourier’s Law of heat conduction:

F  T or

F =  K T

[K = diffusivity]

So, QH = c T/z,

c = CpK

[this is an empirical result]

slide32

QH, continued….

Note 1: K = K (T, wind, sea state, humidity…..)

 10 – 103 cm2/sec at the sea surface

Note 2: Strictly, the concept of a diffusivity K only applies to laminar flows. Where turbulence occurs (everywhere !), this formulation serves as a parameterization only (“eddy diffusivity”).

slide33

QH for the Pacific….

[note the east-west asymmetry]

slide34

Evaporative heat flux, QE….

The latent heat of evaporation L for fresh water is

L = 2494 joules/kg @ 2 C

This much energy must be supplied in order for the state transition from liquid to vapor to occur. In the oceanic case, the energy will be extracted from surface seawater in the form of heat, thus having a net cooling effect.

This evaporated seawater will appear as water vapor in the atmosphere, and the latent heat L will be released into the atmosphere upon condensation of the water vapor.

slide35

QE, continued….

Note: QE is generally the largest term in the heat balance equation, after direct solar input QS.

How can QE be estimated? In general the evaporation rate and associated heat flux are a complicated function of a number of variables and can only be estimated empirically,

QE = QE(Toc, Tatm, S, wind, sea state, humidity,….)

Many estimates of the evaporative heat flux exist, using a variety of parameterizations. Estimated errors are not small, in general.

slide36

QE for the Pacific….

[note the east-west asymmetry]

slide37

The heat balance….

QT = QS + QB + QH + QE + QV + QG = 0

= Qsurf + QV + QG  Qsurf + QV = 0

where Qsurf = QS + QB + QH + QE

What is the total surface heat flux from the ocean, Qsurf ?

slide38

Qsurf for the Pacific….

[note east-west asymmetry]

slide40

Qsurffor the Atlantic….

[note east-west asymmetry]

(watts/m2)

-150

-250

0

-75

-50

-100

slide41

QV , advective heat flux….

Qsurf + QV= 0

[Globally, QV = 0, but this might not be true locally.]

If there is a gain or loss of heat from the sea surface, then the ocean must export or import heat in order to keep the system in thermodynamical equilibrium.

Result: Ocean Circulation.

[The integration proceeds over the surface area of a sector of ocean, including the surface, sidewalls, and the bottom.]

Thus,

slide42

QV, continued….

an ocean losing/gaining heat

at the sea surface

an ocean losing/gaining heat

via advection through the sides

slide43

QV, continued….

+  flux in,

  flux out

(1)

(2)

If (1) > 0 (ocean gains heat from the atmosphere), then (2) < 0 (ocean circulation must transport heat away)

If (1) < 0 (ocean loses heat to the atmosphere), then (2) > 0 (ocean circulation must transport heat in)

Globally, the surface heat budget for the ocean has been estimated from bulk formulae, marine observations, satellites, etc. If we have an estimate of Qsurf, we should be able to estimate Qv.

slide44

QV continued….

units: 1013 watts

Global ocean heat transport, basin integrated

slide45

QV, averaged along latitude by ocean….

[note: Atlantic is different]

slide46

QV(continued)….

Parameterization of Qv :

Let V equal a velocity normal to some surface,

and let T be the temperature of the water flowing with V. The heat transport associated with V and T is then

QV = CVT

where C is the heat capacity and  the density.

Check the units:

[heat/area/time, a flux by definition]

slide47

QV , continued….

[note: T is measured relative to some reference temperature]

QV = CVT

Note the meaning of this parameterization:

Increasing either the temperature, or the flow, will increase the heat transport.

Also note that

warm water flowing north = cold water flowing south

(T>0, V>0) (T<0, V<0)

[consider the Atlantic]

slide48

QV , continued….

Ocean

Atmosphere

Total

Northward heat transport by the ocean and atmosphere

slide49

QV, continued….

Hydrographic lines in the World Ocean Circulation Experiment (WOCE)

slide50

Ocean circulation: surface currents….

[notice E/W asymmetry]

[ 0-500 dbar dynamic ht; maximum range ~ 2 m]

slide51

Annual mean wind stress

The upper ocean circulation is largely wind-driven.

slide52

wind

ocean circ.

Winds: symmetric; Ocean circulation: asymmetric….why?

slide53

(3)

(2)

  • 3 effects:
  • wind
  • friction
  • planetary

(1)

(3)

(2)

Symmetric gyre: vorticity (“spin”) cannot be balanced.

slide54

(3)

(2)

  • 3 effects:
  • wind
  • friction
  • planetary

(1)

(2)

(3)

Asymmetric gyre: vorticity (“spin”) can be balanced.

slide55

Large-scale asymmetric gyre….

possible

[See Stommel (1948) for the details]

Large-scale symmetric gyre….

not possible

slide56

The surface mixed layer in the ocean….

winter

mixed layer

summer

[ thermostad remains

after winter ]

[western N. Atlantic]

slide57

Typical winter

mixed layer in the western N. Atlantic, 3/98

http://flux.ocean.washington.edu

slide58

N

S

winter

Model results,

western N. Atlantic

mode water

summer

slide59

N

S

mode water

Mode water, resulting from winter mixed layer,

exists over much of the western N. Atlantic

slide61

The North Atlantic Oscillation (NAO)….

Low NAO ~ Colder SST in the subtropics

High NAO

Low NAO

slide65

Deep circulation….

Stommel’s T-O2 diagram, suggesting locations of deep water sources

(1958)

slide66

 Stommel’s deep water sources

slide67

Source function for tritium (HTO) input to the ocean

(tritium half-life  12.5 years)

slide68

Atlantic

Tritium in the thermocline of the world ocean, 1970s

Pacific

Indian

slide70

Deep convection and the conveyor belt….

Importance of the global deep circulation

 long term memory (~ 1 ky)

 storage of heat (cold)

 sequestration of carbon

[ apparent time scale: ~1000 years, from 14C ]