Module “Atmosphere” - IMPRS-gBGC · IMPRS-BGC: Module “Atmosphere” – Day 1 – slide Kleidon 02/2016 Motivation for this Course Goal: provide the basic background of climate

Module “Atmosphere”Introduction and Basics

Dr. Axel KleidonMax-Planck-Institut für Biogeochemie

Day 1

29.02.20161

IMPRS-BGC: Module “Atmosphere” – Day 1 – slide Kleidon 02/2016

Motivation for this CourseGoal: provide the basic background of climate physics to

understand biogeochemical cycling within the Earth system

2

Field et al. (1998) Science

Patterns of net primary productivity reflect limitations imposed by climate

gC m -2 yr-1

2


Overview of the Module

Date Topic Lecturer

29.02. Introduction and Basics Kleidon

07.03. Radiation Feist, Marshall

14.03. Dynamics Gerbig, Rödenbeck

21.03. Land Surface Exchange Gerbig, Kolle

22.03. Feedbacks and Change Kleidon, Heimann

Structure:

33


About myself

4

Dr. Axel Kleidon

e-mail: [email protected]: A2.015 (BSYS)phone: 57-6217background: M.Sc. (physics), Ph.D. (meteorology)at MPI-BGC: group leader “Biospheric Theory and Modelling” (since 2006)scientific interests: atmosphere-biosphere interaction, vegetation modelling, thermodynamics of the Earth system, Gaia hypothesis

4


About yourself

• What is your name?

• What is your background in atmospheric/climate science?

• What are your expectations for this course?

55


Literature

66


Overview of Today

Time Topic09:00 Overview of the module09:00

1. Introduction to the climate systemShort break2. Physical principles3. Atmospheric structure

12:30 Lunch break14:00 4. Radiation14:00

5. Global energy balanceShort break6. Atmospheric dynamics7. Biogeochemical cycles

17:00 End

7

Ewith several simple exercises, marked by on the slides

7


1. Introduction

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

-30 -20 -10 0 10 20 30Annual Mean Temperature (°C)

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

0 500 1000 1500 2000Annual Precipitation (mm)

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

0 100 200 300 400Net Solar TOA (W/m2)

Forcing:solar radiation

8

Predictable mean variations in

physical variables

Temperature

Precipitation

Net solar radiation at the Top Of the Atmosphere (TOA)

(includes effects of clouds and ice)

Climatesystem

8


Dominant time periods associated with temperature variations, most are caused by external forcing, but amplified by internal dynamics

Introduction

Peixoto and Oort: Physics of Climate

9

Diurnal scale(day/night)

Annual scale(seasons)Orbital variations

(ice ages)

Synoptic scale(weather)

9


To understand climate, we need to understand the physical relation-ships between solar forcing, internal dynamics, and observable variables

Introduction

SolarForcing

Climate systeminternal dynamics:

transport and conversions

Characteristicpatterns

Observations(temperature,precipitation)

Physicallaws

10

Biogeochemicalcycles

10


The climate system consists of the atmosphere (gaseous), oceans (liquid), land (mixed), cryosphere (solid water) and lithosphere (solid)

The Climate System

11

Atmosphere

Ocean Land

Cryo-sphere

Lithosphere

Water

Surface area of 511 1012 m271% 29%

11


Water on Earth

Relevance to climate: • affects radiative transfer (reflectivity by clouds, snow; greenhouse effect)• transports large quantities of energy

12

Water reservoir % stateOceans 97,0000% liquidIce caps and glaciers 2,2000% solidGroundwater 0,7000% liquidLakes and streams 0,0130% liquidSoil moisture 0,0130% liquidAtmosphere 0,0009% gaseous

Hartmann, Table 1.2

Uneven distribution of water:

Solidfixed volume,fixed shape

Liquidfixed volume,

free shape

Vaporfree volume,free shape

2.5 MJ/kg

0.3 MJ/kg

2.8 MJ/kg

Phase transitions and energy:

requires heatreleases heat

Ocean 7%Snow 35-90%Clouds 20-70%

Hartmann, Tables 3.2 and 4.2

Reflectivity of water:

E

12


Internal dynamics are ultimately driven by differences in radiative exchange with space and coupled by energy and mass fluxes

The Climate System

Atmosphere

Ocean Land

Cryo-sphere

13

Solar radiation

Lithosphere

Terrestrial radiation

Driver: differences in solar radiationRelevance: distribution and conversions of heat and mass

Driver: wind, differences in densityRelevance: heat storage, distribution of heat and mass, evaporation

Driver: radiative heating, windRelevance: heat storage, vertical exchange of heat and mass

342 W m-2 102 W m-2 240 W m-2

13


Atmosphere: Composition

Constituent m-w conc. Total mass (g)Nitrogen N2 28 78% 3.9 1021

Oxygen O2 32 21% 1.2 1021

Argon Ar 40 1% 6.6 1019

Water vapor H2O 18 variable 1.7 1019

Carbon dioxide CO2 44 353 ppmv 2.8 1018

Neon Ne 20 18.2 ppmv 6.5 1016

Krypton Kr 84 1.14 ppmv 1.7 1016

Helium He 4 5.24 ppmv 3.7 1015

Methane CH4 16 1.72 ppmv 4.9 1015

Xenon Xe 131 87 ppbv 2.0 1015

Ozone O3 48 variable 3.3 1015

Nitrous oxide N2O 44 310 ppbv 2.3 1015

Carbon monoxide CO 28 120 ppbv 5.9 1014

Hydrogen H2 2 500 ppbv 1.8 1014

Ammonia NH3 17 100 ppbv 3.0 1013

Nitrogen dioxide NO2 46 1 ppbv 8.1 1012

CFC-11 CCl2F2 121 480 pptv 1.0 1013

…

based on Hartmann, Table 1.1

14

• > 99% of the atmosphere consists of inert gases (they do not react without addition of energy)• Gases of particular importance:

• Greenhouse gases (absorb terrestrial radiation), e.g. H2O, CO2, CH4, N2O, CO

• UV absorption: oxygen (O2 and O3, largely of biotic origin, reactive)

• Atmospheric chemistry (e.g., CFC-11; CFC-12 are man-made compounds that affect ozone chemistry)

• Water vapor: highly variable • Also: droplets, aerosols

m-w: molar weight, conc.: volume concentration

14


Atmosphere: Temperature

Hartmann, Figure 1.2

• The atmosphere shows a typical vertical profile of temperature• The lapse rate describes the rate of decrease in temperature with height• Distinct layers defined by a sign change of lapse rate• Inversions describe areas in which temperature increase with height• Radiative and convective processes shape this profile• Most climate-relevant dynamics take place in the lowest layer (troposphere)

1515



Atmosphere: Pressure

• Pressure defined as force per unit area in units of hPa

• Air pressure typically decreases nearly exponentially with height• Total pressure is composed of sum of partial pressures of different compounds• Partial pressure of water vapor decreases much faster with height

16

p =F

A=

mg

A

p! pressurem! massF! force

A! areag! = 9.81 m s-2

E

16


Ability of air to hold moisture is restricted up to its saturation vapor pressure, esat, which depends strongly on temperature T.

Atmosphere: Water vapor

17

0

2000

4000

−20 −10 0 10 20 30

Air Temperature (°C)

Satu

ratio

n pr

essu

re (P

a)

T: Temperature in Kelvinesat: partial pressure in Pa

Saturation increases by about 6.5%/K at global mean temperature of

15°C

esat = 611 · e19.83�5417/T

17


Oceans: Composition

Component g/kg

Chloride 19,353Sodium 10,760

Sulfate 2,712Magnesium 1,294Calcium 0,413

Potassium 0,387Bicarbonate 0,142Bromide 0,067

Strontium 0,008Boron 0,004

Fluoride 0,001

Hartmann Table 1.3

18

Mean Compositionof Sea Water

Spatial deviations of mean salinity (35 g/kg)

• Relevance: salinity affects density of sea water (and thereby buoyancy, thermohaline circulation)• Driver: mostly the hydrologic cycle

http://www.nodc.noaa.gov/OC5/

18


Oceans: Structure


Mixed layer

19

Thermocline

Deep ocean

The vertical temperature profile can typically be partitioned into three regions:• Mixed layer: top 20-200m, well mixed, exchanges with atmosphere• Thermocline: rapid change in temperature• Deep ocean: below thermocline

19


Cryosphere

Reservoir Volume (km3) %Antarctic Ice Sheet 30.1 106 89,3 %Greenland Ice Sheet 2.6 106 8,6 %Mountain Glaciers 0.3 106 0,8 %Permafrost 0.2-0.5 106 1,0 %Seasonal snow 2-3 103 Sea Ice 4-10 104

• Mostly in polar regions• > 10% of Earth’s surface (3% ice sheets, 7% sea ice)• Highly reflective surface (reflection of 60% and more)• Stores freshwater• Affects ocean salinity, sea level• Forms topography

20

NASA/Goddard Space Flight Center E

20


Land

• 29% of Earth’s surface• Surface properties are strongly affected by vegetation

• reflectance (e.g., Sahara: 35%, rainforest 14%)• ability to evaporate (stomatal conductance, rooting depth)

• Topography affects precipitation, heating, and flow• relevant to humans:

• 10-13% farming and settlements• 20-25% grazing lands

180˚ 270˚ 0˚ 90˚ 180˚-90˚ -90˚

-60˚ -60˚

-30˚ -30˚

0˚ 0˚

30˚ 30˚

60˚ 60˚

90˚ 90˚

-10000 -7500 -5000 -2500 0 2500 5000 7500 10000

height (m)

180˚ 270˚ 0˚ 90˚ 180˚-90˚ -90˚

-60˚ -60˚

-30˚ -30˚

0˚ 0˚

30˚ 30˚

60˚ 60˚

90˚ 90˚

-10000 -7500 -5000 -2500 0 2500 5000 7500 10000

height (m)

2121


2. Physical Principles

1. Conservation laws: Some basic physical quantities are conserved. This results in balance equations which account for the changes of content with the difference of in- and outfluxes

• energy

• mass

• momentum

• [angular momentum, charge, ...]2. Thermodynamics: Conversions of energy, mass and momentum are associated with energy conversions. Thermodynamics sets the rules:

• energy is conserved during conversion process

• energy is dispersed (increase of entropy) => sets direction

• sets limits to the magnitude of energy conversions22

22


Balance Equations

• Energy balances describe changes in heat storage in relation to heat fluxes and predict changes in temperature

• Example: surface energy balance on land

• Specific heat capacity cp and density ρ determine how fast temperature changes as a result of heat fluxes.

23

E

cp⇢dTs

dt= Rs �Rl �H � �E �G

change in heat storage

radiativefluxes

sensible andlatent heat fluxes

groundheat flux

23


Balance Equations

• Simplest representation of the mass balance is the linear reservoir

• Steady-state represents the case in which contents do not change in time (dW/dt = 0). In this case, contents are given by:

• Other example: diffusion equation

24

ContentsW

InfluxJin

OutfluxJout

dW

dt= Jin � W

�

• outflux assumed to be proportional to contents W with a residence time τ

Wss = Jin �

change incontents

sum offluxes

24


Thermodynamics

There are a total of four laws of thermodynamics. The first and second law are the most important for calculations.

• First law states energy conservation, with heat exchanged dQ being balancedby internal changes dU and work performed in time:

• For heat fluxes through time, it canbe expressed as:

• can be generalized to all forms of energy conversions

25

heatcontent

U

heat fluxJin

heat fluxJout

powerP

dU = dQ� dW

cp

⇢dT

dt=

dU

dt= (J

in

� Jout

)� P

25


heatcontent

U

heat fluxJin

heat fluxJout

Thermodynamics

• The second law describes the direction of increased energy dispersal among energy conversion processes. It requires the entropy S during a conversion process to increase (or stay unchanged)

• Typically, this is evaluated in the contextof an entropy balance of the system:

26

powerP

dS � 0Tin

Tout

dS

dt= � +

Jin

Tin

� Jout

Tout

entropy producedby processes

within the system(σ ≥ 0)

entropy exchangeof the system

change ofentropy ofthe system

26


Thermodynamics

27

Externalheating

Externalcooling

System A(isolated)

Steady state

Redistributionof heat

Initial stateTime

System B(non-isolated)

Thermodynamic equilibrium

Thermodynamic disequilibrium



Hot Cold

Hot Cold Externalheating

Externalcooling

Warm Cool

Warm Warm

Example

Cup of cold coffee

Pot ofboiling water

Examples of how the Second Law sets the direction(in general: fluxes deplete their driving gradient)

27


Thermodynamics

Atmospheric dynamics: generation of kinetic energy and frictional dissipation affect wind speeds

• generation of kinetic energy out of heating differences (surface-atmosphere, tropical-polar)

• dissipation of kinetic energy due to friction

Biotic activity: photosynthesis and respiration alter biomass

• “endotherm”: ! CO2 + H2O + light → CH2O + O2

• exotherm:! ! CH2O → CO2 + H2O + heat

solar radiation heatchemical energy

28

heating difference heatkinetic energy

28


heatcontent

U

heat fluxJin

heat fluxJout

Thermodynamics

• Both laws combined set the Carnot limit on energy conversions in steady state. It represents the idealized upper limit on energy conversion that fulfills the second law when

• Then, the two laws take this form:

29

powerP

Tin

Tout

� = 0

Jin

Tin

� Jout

Tout

= 0

Jin

� Jout

� P = 0first law:

second law:

Carnot limit: P = Jin

Tin

� Tout

Tin

⌘ =Tin

� Tout

Tin

Carnot efficiency:

29


Thermodynamics

30

Heat fluxWarm temperature, low entropy

Power = work/timeNo entropy

Waste heat fluxCold temperature, high entropy

Trade-off between power and waste heat flux affects

entropy exchange of the system

a. Steam engine b. Atmospheric heat engine

Surface

Absorption of solar radiation Warm temperature, low entropy

Emission of terrestrial radiation Cold temperature, high entropy

Heatengine

Power

MotionHeat flux

Atmosphere

Radiativeexchange

Thermodynamics sets limits to work and energy conversions

E

30


Thermodynamics

31

Solar radiation

Not reflected70%

Solar radiative heating

Differential heating

Atmospheric circulation

Oceanwaves

Oceanic circulation

≈ 175000 TW

Geometric differences40%

Maximum power efficiency2%

≈ 123000 TW

≈ 49000 TW

≈ 1000 TW

Conversion6%

≈ 60 TW

≈ 5 TW

Conversion8%

Work done by Earth system processes is small compared to radiative fluxes, but

essential to maintain dynamics

Comparison: 17 TW of human primary energy consumption

(0.4 TW in Germany)Kleidon (in press)

1 TW = 1012 W

31


Thermodynamics

• The ideal gas law describes relationship between pressure p and temperature T of air. It is typically expressed for air in terms of air density ρ and the gas constant Ra as

• very often used to determine air density!

32

p = ⇢RaT

E

T temperature K

Ra gas constant for air 287 J kg -1 K-1

ρ density kg m-3

p pressure Pa

T (in °C) ρ (in kg m-3)0 1,2935 1,270

10 1,24815 1,22620 1,20525 1,18530 1,165 at p = 1013.25 hPa

32


3. Atmospheric Structure

33

The basic structure of the atmosphere in terms of vertical temperature and pressure variations can be derived from simple physics.

First law: conserve total energy (heat + potential) when air is lifted by dz

cpdT + gdz = 0

Hydrostatic balance: balance of gravity with pressure drop across a layer of thickness dz:

dp = ��gdz

Ecommonly used to convert pressure into heighteffect of moisture?

33


Atmospheric Structure

When air is lifted adiabatically (no heat exchange), pressure and temperature changes. To make air parcels comparable, a potential temperature is introduced.The potential temperature θ isdefined as the temperature ofair with (T, p) when brought toa reference pressure p0:

- is derived from heat engine- is used to determine vertical motion

34

� = T

✓p0p

◆R/cp

R/cp: !2/7p0: ! Reference pressure

(p0 = 1000 hPa)

Equator SN

300 K

350 K

200 K

300 K

260 K

100 hPa

10 hPa

1000 hPa

100 hPa

1000 hPa

1 hPa

260 K260 K

Source: ECMWF - ERA40

34


Atmospheric Structure

35

Example: Radiosounding, Meteorological Observatory Lindenberg, 09. Juli 2011; 12h

Temperaturpotentielle Temperatur

Höh

e (m

)

0

2500

5000

7500

10000

12500

15000

Temperatur (°C)−60 −40 −20 0 20 40 60 80 100

Hei

ght (

m)

Temperature (°C)

T

Hei

ght (

m)

θ

Pressure[hPa]

Temperature [°C]

Height[m]

pot. Temp.[K]

1002 27,0 115 26,8

1000 25,8 132 25,8

925 19,0 810 25,6

850 12,0 1528 25,6

752 4,8 2542 28,4

http://weather.uwyo.edu/upperair/sounding.html

observed inferred

35



Lunch Break

we meet again at 14:00

36


Solar radiation is the primary forcing of the Earth system. Its magnitude and its variation is governed by radiation laws.

• Radiation is characterized by wavelength λ (“lambda”, unit: m) or frequency ν (“nu”, unit: Hz = 1/s)

• Relationship between wavelength and frequency:

• Wave-particle (“photon”) dualism, with photon energy (“quantum”):

• Every object with a temperature emits radiation according to the Stefan-Boltzmann law:

• Peak emission occurs at a wavelength λmax given by Wien’s displacement law:

4. Solar radiation

37

c = � · ⌫

E = h⌫

R = � · T 4

�max

T = bT temperature in Kelvinc speed of light 3 108 m s-1

h Planck constant 6.63 10-34 Jsσ Stefan-Boltzmann constant 5.67 10-8 W m-2 K-4

b constant 2.8978 10-3 m K E

37


Radiative Spectra

38

The emission spectra for T = 5760K and T = 255K (relative intensity) are well-separated. This results in the separation between solar vs. terrestrial radiation.

0

0.2

0.4

0.6

0.8

1.0

0.1 1 10 1000

0.2

0.4

0.6

0.8

1.0

0.1 1 10 100

Rel

ativ

e in

tens

ity

Rout = 342 W m-2

EarthSolar radiation(low entropy)

Reflected solar and reemitted terrestrial radiation

(high entropy)

Rin = 342 W m-2

Wavelength (µm)

Irreversible processes

Tsun= 5760 K

Wavelength (µm)

Rel

ativ

e in

tens

ity Tearth = 255 K

scattered

re-emitted

38


The magnitude of solar radiation at a given place on Earth is determined by the solar luminosity, the distance between the Sun and the Earth, and the orientation of the Earth’s surface.

SunEarth

Solar Radiation on Earth

Solar luminosity:L0 = 3.9 1026 W

I =L0

4π · d2

Radiative flux atdistance d:

DistanceEarth-Sun:

d = 150 109 m

I = 1379 W m-2

39

Solar “constant”:Mean influx: I/4 = 345 W m-2

39


Variations in Solar Radiation

Variations in solar radiation are caused by:

1. Solar activity (e.g. luminosity)2. Orientation of surfaces to solar radiation3. Aspects of Earth’s orbit (e.g. distance)

Sun Earth

4040


Variations in Solar Activity

• Millions of years:• increase of solar activity over time• 4.5 billion years ago, 70% of today’s luminosity

• Thousands of years • interannual: “sunspot” cycle

approx. 11 years

science.nasa.gov/ssl/pad/solar/sunspots.htmwww.sel.noaa.gov

41

starting in 1750

41


Orientation of Surfaces

Sun

January

July

vernal

equinox

autumnal

equinox

distance to perihelion

147,000,000 km

summer

solstice

distance to aphelion

152,000,000 km

winter

solstice

April

October Arctic circle (winter)

Tropic of cancer

Tropic of capricorn

Antarctic circle (summer)

Equator

tilt

(obliquity)

North pole

South pole

42

E

I =L0

4π · d2

42



Sun

sola

r rad

iatio

n

AB

C

Sun

Sun

A:

B:

C:

= 90°

0° < < 90°

= 0°

declinationangle

Earth

Equator

North pole

South pole

Arctic circleTropic of cancer

Tropic of capricorn

Antarctic circle

43

Definition of zenith angle θ: Sun

sola

r rad

iatio

n

AB

C

Sun

Sun

A:

B:

C:

= 90°

0° < < 90°

= 0°

declinationangle

Earth

Equator

North pole

South pole

Arctic circleTropic of cancer

Tropic of capricorn

Antarctic circle

E

cos ✓ = sin lat · sin � + cos lat · cos � · coshourCalculation of the zenith angle:

lat: ! latitude in degree δ: ! declination angle (-23.5° at winter solstice, 0° at equinox, 23.5° at summer solstice; otherwise more ! complicated to calculate)hour:! time in degree (0° = noon)

Is = I? · cos✓Accounting for solar inclination:

43



• Insolation-weighted daily average solar zenith angle as a function of latitude:

44

Hartmann, Fig. 2.8

44



• Seasonal and latitudinal variation of solar irradiation in W m-2

45


45


Orbital Variations

Three orbital parameters of the Earth vary on time scales of thousands of years: eccentricity (related to distances of perihelion and aphelion), timing of perihelion (precession), and tilt (obliquity)

46

Sun

January

July

vernal

equinox

autumnal

equinox

distance to perihelion

147,000,000 km

summer

solstice

distance to aphelion

152,000,000 km

winter

solstice

April

October Arctic circle (winter)

Tropic of cancer

Tropic of capricorn

Antarctic circle (summer)

Equator

tilt

(obliquity)

North pole

South pole

evidence from ocean and ice cores supports orbital causes for ice ages

altered orbital parameters alter seasonal insolation patterns, particularly near the poles

46


5. Global energy balance

47

team view of the closure for the TOA radiation budget. The TOA imbalance in the original CERES products is reduced by making largest changes to account for the uncertainties in the CERES instrument absolute calibration. They also use a lower value for solar irradiance taken from the recent TIM observations (Kopp et al. 2005).

Several atlases exist of surface f lux data, but they are fraught with global biases of several tens of watts per meter squared in unconstrained VOS observation-based products (Grist and Josey 2003) that show up, especially when net surface flux fields are globally averaged. These include some based on bulk flux formulas and in situ measurements, such as the Southampton Oceanographic Centre (SOC) from Grist and Josey (2003), WHOI (Yu et al. 2004; Yu and Weller 2007), and satellite data, such as the HOAPS data, now available as HOAPS version 3 (Bentamy et al. 2003; Schlosser and Houser 2007). The latter find that space-based precipitation P and evapora-tion E estimates are globally out of balance by about an unphysical 5%. There are also spurious variations over time as new satellites and instruments become part of the observing system.

Zhang et al. (2006) find uncertainties in ISCCP-FD surface radiative fluxes of 10–15 W m−2 that arise from uncertainties in both near-surface temperatures and tropospheric humidity. Zhang et al. (2007) computed surface ocean energy budgets in more detail by com-bining radiative results from ISSCP-FD with three

surface turbulent f lux estimates, from HOAPS-2, NCEP reanalyses, and WHOI (Yu et al. 2004). On average, the oceans surface energy flux was +21 W m−2 (downward), indicating that major biases are present. They suggest that the net surface radiative heating may be slightly too large (Zhang et al. 2004), but also that latent heat flux variations are too large.

There are spurious trends in the ISCCP data (e.g., Dai et al. 2006) and evidence of discontinuities at times of satellite transitions. For instance, Zhang et al. (20007) report earlier excellent agreement of ISCCP-FD with the ERBS series of measurements in the tropics, including the decadal variability. However, the ERBS data have been reprocessed (Wong et al. 2006), and no significant trend now exists in the OLR, suggesting that the previous agree-ment was fortuitous (Trenberth et al. 2007b).

Estimates of the implied ocean heat transport from the NRA, indirect residual techniques, and some coupled models are in reasonable agreement with hydrographic observations (Trenberth and Caron 2001; Grist and Josey 2003; Trenberth and Fasullo 2008). However, the hydrographic observations also contain significant uncertainties resulting from both large natural variability and assumptions associated with their indirect estimation of the heat transport, and these must be recognized when using them to evaluate the various flux products. Nevertheless, the ocean heat transport implied by the surface fluxes provides a useful metric and constraint for evaluating

products.

THE GLOBAL MEAN ENERGY BUDGET. The results are given here in Table 1 for the ERBE period, Table 2 for the CERES period, and Fig. 1 also for the CERES period. The tables present results from several sources and for land, ocean, and global domains. Slight differences exist in the land and ocean masks, so that the global value may consist of slight-ly different weights for each component.

ERBE period results . For the ERBE period, Table 1 presents results from KT97 for comparison with those

FIG. 1. The global annual mean Earth’s energy budget for the Mar 2000 to May 2004 period (W m–2). The broad arrows indicate the schematic flow of energy in proportion to their importance.

314 MARCH 2009|

Trenberth et al. (2009)

Ts ≈ 288 K

47


Global Energy Balance

• The global energy balance in steady state balances absorbed solar radiation with emitted terrestrial radiation.

• Simplest energy balance yields the radiative temperature of a planet

• Planetary albedo describesthe fraction of reflected radiationresulting from ice, clouds, …• example surface albedo values:

ice ≈ 0.6; desert ≈ 0.35; rainforest ≈ 0.14; ocean ≈ 0.06

• Earth’s radiative temperature: Tr ≈ 255 K

48

I

4(1 − ap) − σT 4

r = 0

T temperature K

αp planetary albedo 0,3

I solar constant 1367 W m-2

σ Stefan-Boltzmann const 5.67 10-8 W m-2 K-4

48


Planetary Albedo

Venus Earth Mars

albedo = 0.71 albedo = 0.3 albedo = 0.16

49

E

d = 108 106 km d = 249 106 kmd = 150 106 km

49


Greenhouse Effect

• A very simple, two-layer model of the surface-atmosphere to demonstrate the importance of the greenhouse effect

50

atmosphere

surface 0 = Rs,in(1� a)� �T 4s + �T 4

a

0 = �T 4s � 2�T 4

a

Greenhouseforcing

E

50


6. DynamicsMotion results mostly from buoyancy, i.e., density differences, and related effects (especially in the vertical).

51

State B: State C:State A:

ρl

ρh

heavier fluid with density

ρh

lighter fluid with density ρl

lighter fluid rises due to buoyancy

lighter fluid spreads (if possible)

51


DynamicsDifferences in density are caused by temperature differences (due to heating and cooling) and salinity differences (ocean, mostly due to evaporation - precipitation)

52

1.15

1.20

1.25

1.30

1.35

1.40

−20 −10 0 10 20 30

Salin

ity (‰

)

40

0

20

30

10

0 10 20 30Temperature (°C)Temperature (°C)

Den

sity

(kg

m-3

) Density - 1000 (kg m-3)

Atmosphere:ideal gas law

Ocean:empirical formula

52


Dynamics

A simple upper bound to large-scale heat transport by motion can be inferred from a simple 2-box model and thermodynamics.energy balances of two boxes:

Carnot limit describes limit of converting heat into work:

here: ! Pmax: !generation of kinetic energy!! J: ! ! heat transport

53

Pmax

= J · Th

� Tc

Th

= J · �T

Th

Jin,h Jout,cJin,cJout,h

TcTh

Atmospheric heat engineJ

P

Tropics Extratropics

Solarradiation

Terrestrialradiation

0 = Jin,h

� Jout,h

� J0 = J

in,c

� Jout,c

+ J

E

53


Dynamics

Steps to a estimate:- use a simple approximation for emission of radiation (instead of Stefan-Boltzmann law):

- with energy balances yields an expression for ∆T (note: heat flux attempts to deplete gradient):

- results in a quadratic function of Pmax:

- yields a maximum estimate of:

- estimate: ∆Jin ≈ 98 W m-2; Jopt ≈ 24 W m-2 [obs: 23 W m-2]

54

Jout

(T ) ⇡ J0 + kT

�T =�Jin � 2J

kPmax

= J · �Jin

� 2J

kTh

Jopt

=�J

in

4Pmax

=(�J

in

)2

8kTh

54


Dynamics

-200

-100

0

100

200

300

400

Rad

iativ

e Fl

ux (W

m-2)

-90 -60 -30 0 30 60 90

-5.0

-2.5

0.0

2.5

5.0

7.5

10.0

Wat

er F

lux

(mm

d-1)

-90 -60 -30 0 30 60 90

Latitude

-50

0

50

100

150

200

250

Ener

gy F

lux

(W m

-2)

-90 -60 -30 0 30 60 90

NorthSouth

top of atmosphere

surface

solar

terrestrial

net

net

solar

terrestrial

latent heat

sensible heat

precipitation

evaporation

net

55

• Top: zonal means of solar and terrestrial radiation at the top of the atmosphere

• Difference between solar and terrestrial radiation (net) describes net heat transport

• Middle: surface energy balance components

• Bottom: differences in precipitation and evaporation imply differences in moisture transport by atmospheric dynamics

• Release of latent heat by precipitation causes moist convection and generation of motion, esp. in the tropics, which results in strong coupling between energy fluxes, hydrologic cycling, and generation of motion

ECMWF reanalysis

55


Biogeochemical Cycles: Overview

• Global cycles of water and carbon

• Residence times vs. spatial and temporal variations

• Planetary evolution

5656


Global Hydrologic Cycle

Estimates of reservoirs and fluxes of the global hydrologic cycle in 1012 m3 and 1012 m3 yr-1

57

Atmosphere

Ocean Land

Lithosphere

1 348 000

13

22 780

2258062

Soils

Groundwater

361324

99 62…

E

57


Global Carbon Cycle

Estimates of reservoirs and fluxes of the global carbon cycle in 1012 kg C and 1012 kg C yr-1

58

Atmosphere

Ocean Land

Lithosphere

38 000

750

1500 Soils

9090

120 60

E

500Vegetation

60

58


Biogeochemical Cycles

Residence times include information about how fast a stock can change for a given flux.

59

Precipitable water(0 - 50 kg m-2)

Atm. CO2 concentrations(350-380 ppmv)

59



Catling (2004)

60

The atmospheric composition of Earth has likely changed substantially throughout its history, mostly due to biotic activity.

Human Society

Biosphere

FossilBiomass

MarineTerrestrial

Ene

rgy

Cap

ture

(EJ/

y)

anox

ygen

ic p

hoto

synt

hesi

s

oxyg

enic

pho

tosy

nthe

sis

land

pla

nts

pale

olith

ic fir

e us

e

neol

ithic

revo

lutio

n

~3 Ga ~1.5 Ga ~350 Ma ~10,000 BCE ~1850 CE ~2000 CE

10-1

100

101

102

103

104

indu

stria

l rev

olut

ion

Figure 1. Energy capture in the biosphere and human society. Data and sources in Table S1.

O2 persistently >15% of the atmosphere since 370 Ma (Belcher and McElwain, 2008; Scott and315

Glaspool, 2006).

The rise in atmospheric oxygen and increase in food supply brought about by land plants has

allowed a flourishing of animal complexity from aerobic pathways – including the emergence of us

humans. Today, the total global energy flux through heterotrophic biomass, based on a 10% conver-

sion efficiency of 100 PgC yr�1 with energy density 40 kJ gC�1, is ⇠ 400 EJ yr�1, roughly half on320

land and half in the ocean. Natural fires additionally consume ⇠ 55 EJ yr�1 (1.4 PgC yr�1) (Eliseev

et al., 2014), and human-induced fires ⇠ 45 EJ yr�1 (1.1 PgC yr�1) (Haberl et al., 2007), giving a

total biomass burning flux today of ⇠ 100 EJ yr�1 (⇠ 2.5 PgC yr�1) (Randerson et al., 2012), or

⇠ 2.5% of the energy and carbon captured in photosynthesis.

3 Revolutions in human history325

Like all animals humans are heterotrophs. Our biological metabolism relies on the products of pho-

tosynthesis. At the same time humans are exceptional among animals in creating and maintaining

a social metabolism via breeding and cultivating plants and animals, in constructing buildings and

large infrastructure systems and in producing numerous artifacts (Fischer-Kowalski and Hüttler,

1998). The social metabolism inevitably extends total human energy capture and material use be-330

yond the biological requirements. In modern industrial societies the amount of energy and materials

used to produce and reproduce domesticated livestock and all artifacts typically is an order of mag-

nitude larger than the basic biological metabolism of the human population itself. For the following

comparison between human energy use and the primary productivity of the entire biosphere, it is

therefore important to keep in mind the different trophic levels involved, autotrophs versus het-335

10

Lenton et al (in discussion, ESDD)

Over Earth evolution, the energy capture (power and dissipation) by life (and humans) likely increased substantially.

60



goddess of spring (Retallack, 2000b), is undeni-able for the annual spring fall and autumn rise ofatmospheric CO2 with northern hemisphere leafsprouting and shedding (Figure 2(a)). Thisexplanation is especially demonstrated by themuted and out-of-phase annual fluctuation of CO2

in the southern hemisphere (Mooney et al., 1987),where there is less fertile land, more evergreenplants, and different seasons. The questionaddressed here is whether the Proserpina principleoperates on geologically significant timescales,and so far, such a simple idea does not conflictwith the history of life and paleoclimate outlinedhere.On evolutionary timescales, it is the biochemi-

cal evolution of lignin, pyrethrin, caffeine, andother substances that deter herbivory, digestion,and decay, which affect rates of carbon burial insediments as the principal long-term control onatmospheric CO2 levels. The role of trees andtheir soils in Late Paleozoic carbon seques-tration, cooling, and glaciation is widely accepted(Berner, 1997; Algeo and Scheckler, 1998; seeChapter 5.06). The role of humans in globalwarming is also becoming well known (Vitouseket al., 1997b). According to the Proserpinaprinciple, we may not have been the onlyorganisms to have had significant effects onclimate. There remain many other instances ofglobal change less clearly related to changes inlife and soils, in part because the numerouspaleosols of appropriate age have not yet beenstudied in detail. Asteroid impacts, volcaniceruptions, and methane clathrate dissociation

events also affect life and the carbon cycle,producing transient greenhouse events (Retallack,2001b). Ocean currents and mountain buildingalso are likely to play a role in carbon sequestra-tion (Raymo and Ruddiman, 1992;Ramstein et al.,1997). Soils and their ecosystems play animportant role in the carbon cycle today, andthe history of that role now decipherablefrom paleosols appears ripe for modeling andother quantitative comparisons with other likelycontrols on global paleoclimate change.

ACKNOWLEDGMENTS

Nathan Sheldon, Hope Jahren, and Tim Whitehave been sounding boards for the ideas presentedhere. I also thank J. I. Drever and H. D. Hollandfor helpful reviews.

REFERENCES

Algeo T. J. and Scheckler S. E. (1998) Terrestrial–marineteleconnections in the Devonian: links between the evolutionof land plants, weathering processes and anoxic events. Roy.Soc. London Phil. Trans. B353, 113–130.

Allen J. R. L. (1974) Geomorphology of Siluro-Devonianalluvial plains. Nature 249, 644–645.

Amundson R., Stern L., Baisden T., and Wang Y. (1998) Theisotopic composition of soil and respired CO2.Geoderma 82,83–114.

Bakker R. T. (1985) The Dinosaur Heresies.WilliamMorrow,New York.

Barley M. E., Pickard A. L., and Sylvester P. J. (1997)Emplacement of a large igneous province as a possible causeof banded iron formation 2.45 billion years ago. Nature 385,55–58.

Figure 16 The Proserpina principle relates variation in atmospheric CO2 concentration with coeval evolutionaryand ecological events on a variety of timescales. Carbon sequestering evolutionary innovations and ecologicaltransitions (closed symbols) alternate with carbon oxidizing evolutionary innovations and ecological transitions

(open symbols). The CO2 curve is a composite of those shown in Figure 2 and by Kasting (1992).

Soils and Global Change in the Carbon Cycle over Geological Time600

Retallack (2004)

Evolution of atmospheric CO2 and biotic events

6161



© Nature Publishing Group1965

Lovelock (1965) on habitability and life

Lovelock (1975): atmospheric composition

6262


Biotic Activity in the Earth System

63

Kleidon (in press)

Solarradiation

Differentialheating

Motion

Interiorcooling

Hydrologiccycling

Geochemicalcycling

Motion

Bioticactivity

Humanactivity

Photochemicalconversion

Thermalconversion

Photochemicalconversion

Radiativeconversion(potential)

cycling activity activity

Interior Earth

175000

123000

4600

20

< 22

21528 8

1000

< 12% < 93%< 12%

25Thermalconversion

All estimates in 1012 W

Atmosphere

Fossilfuels

17

The strong effect of biotic activity on the Earth system can be identified by the rate by which life generates chemical free energy by photosynthesis.

Estimates of planetary drivers and energy conversion limits for different Earth system processes. Black arrows indicate flows of energy, while dashed arrows indicate the effects of mass exchange limitations. All estimates are given in units of 1012 W.

63


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Documents

Module “Atmosphere” - IMPRS-gBGC · IMPRS-BGC: Module “Atmosphere” – Day 1 – slide Kleidon 02/2016 Motivation for this Course Goal: provide the basic background of climate