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biogeochemical cycles i carbon and oxygen

https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdfhttps://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdf

https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS3.pdf

Biogeochemical cycles I carbon and oxygen

9:00 – 10:30 Wednesday Dec 8

carbon cycle questions
Carbon Cycle Questions

What caused changes in CO2 between glacial and interglacial? Did CO2 force or respond to climate?

More on isotopes and what they can tell you

More detail on the ocean biogeochemical cycling

What is the long-term fate of CO2 we add to the atmosphere?

controls on co 2 vary with timescale
Controls on CO2 vary with timescale
  • Millions of years; volcanic CO2 supply and weathering uptake (tectonics)
  • Thousands of years (glacial/interglacial cycles)

Ocean determines atmospheric CO2

  • Decades-Centuries: many processes (growth/decomposition in forest, soil OM exchange, changes in ocean circulation, …..
  • Seasonal cycle
    • Ocean gas exchange
    • Terrestrial biosphere
slide4

Distribution of Carbon;

1015 grams =

1 Petagram (Pg)

ATMOSPHERIC CO2

640 X1015 g C

Response times are seasons to centuries

LIVING BIOMASS

830 X1015 g C

DISSOLVED ORGANICS

1500 X1015 g C

Response times are centuries to millennia

ORGANIC CARBON IN SEDIMENTS AND SOILS

3500 X1015 g C

CO2 DISSOLVED IN OCEANS

38,000 X1015 g C

LIMESTONE AND SEDIMENT CARBONATES

18,000,000 X1015g C

Response times are tens of thousands to millions of years

TRAPPED ORGANIC CARBON: NATURAL GAS, COAL PETROLEUM, BITUMEN, KEROGEN

25,000,000 X1015 g C

last glacial maximum lgm
Last Glacial Maximum (LGM)
  • Time of maximum ice sheet extent centered on 21 ka (19-23 ka)
  • Glacial world was significantly different from today:
    • Ice sheets/sea level
    • Temperature
    • Greenhouse gases
    • Aridity
    • Winds
    • Vegetation
    • Ocean circulation
  • Continental configuration & insolation were nearly identical to today…
    • pCO2 and ice volume are most likely factors affecting LGM climate
  • Major boundary conditions are well known and abundant paleoclimate data is available – a crucial test for climate models!
where did the c go during glacials
Where did the C go during glacials?
  • pCO2 changes from 190-280 ppm (30%) in a few 1000 y, this cannot be due to weathering or volcanic CO2.
  • Fast changes during Quaternary can only be explained by rapid C exchange among surface reservoirs!

A useful way to measure how C has moved among various reservoirs is using δ13C

Total change of atmospheric C: 90 ppm or nearly 200 PgC

how to get c into the deep ocean
How to get C into the deep ocean?
  • Physical changes (Solubility pump)
    • Temperature & salinity
    • Isolation of deep from surface waters (decreased ventilation)
  • Stronger biological pump
    • Fe fertilization
    • Increase of whole ocean nutrient content
    • Change in Redfield ratios (more efficient C pumping)
  • Changes in ocean [CO32-]
    • Increased CaCO3 weathering
    • Decreased coral reef growth
    • Change in C0rg:CaCO3 export to deep ocean
the solubility and biological pumps changes in ocean circulation
The solubility and biological pumpsChanges in ocean circulation

Surface waters equilibrate quickly; CO2 reacts with water

Falling particles move organic carbon into the deep ocean

Sinking waters in polar regions isolate water that has equilibrated at the surface (cold waters)

temperature salinity mechanism is easy to test
Temperature/Salinity mechanism is “easy” to test…
  • Cooler SSTs = increased CO2solubility; lowers CO2 by 30 ppm
  • But, also higher salinity, decreases

Solubility (+6.5)

Net T & S effect  -23.5 ppm (as opposed to 90 ppm observed) … not enough

was it stored on land
Was it stored on land?

Crowley et al., 1997, JGR

slide12

Was it stored on land?Still unresolved, but in evidence so far is that LESS C was stored on land…Where did it go? The deep ocean is the only reservoir big enough and slow-exchaging enough

  • Decreased temperate forests
  • Increased northern tundra
  • Decreased tropical rain forests
  • Reduced growth due to low pCO2

Crowley et al., 1997, JGR

slide13

13C changes in benthic foraminifera should show this transfer

The 13C in benthic forams varied between the last glacial and today http://www2.ocean.washington.edu/oc540/lec01-28/

Comparison of the d13C records from equatorial (V19-30) and northeast Pacific (W8709A-8) cores spanning the last glacial cycle.

Based on this record, the glacial ocean 13C was roughly 0.4per mil  lighter during the LGM (indicating transfer of isotopically light C from land to ocean), and consistent with a smaller land biosphere. However, the decrease predicted by transferring 530 PgC is less, only -0.35 per mil; something else going on…

slide14

Remember:  Land pants (C3) have  13C of about -25 per mil  (R =0.975= 13C/1000 +1))          Ocean total CO2 (Holocene)  13C is about +0.50 per mil  (R=1.005)       LGM Ocean total CO2    = 0.50  (Holocene value) minus 0.35 per mil = 0.15 per mil

We can use the difference in 13C between ocean+atmosphere today and in the LGM to estimate the how much less land C there was on the LGM by mass balance:

Carbon mass balance:

Land]today  + OA]today = Land]glacial + OA]glacial

2,000 (land today)   + 3,6500 (35,100 in ocean, 500 in preindustrial atmosphere)  = Total = 38,600 Pg of carbon

13C mass balance:

(2000)(0.975)  +  (38,600)(1.0005) =  Land]glacial (0.975)  + OA]glacial(1.00015)

                                           = Land]glacial (0.975)  + [38,600 - Land]glacial](1.00015)

Solving for Land]glacial we get ~1500 Pg C  (or 500 Pg C less than today)

Other differences:                    Preindustrial                 LGM

Land :                2000                         1500     (from 13C in benthic forams)Atmosphere         500                          360     (from pCO2 in ice cores)Ocean             31,500                      35,740  (by difference)

slide15

NOTE: There are some problems here. 

  • The 500 Pg C difference between LGM and today in the biosphere calculated using 13C change is at the very low end of the range that has been estimated from paleovegetation maps (700-1300 PgC)
  • There are a number of potential problems with 13C in forams, mostly involved with
  • differences in 13C between coexisting benthic species (vital effects) coupled with selective dissolution
  • the tendency of benthic forams to use DIC that is in part derived from the decomposition of organic material in sediment pore waters.
  • (3) the distribution of C3 and C4 plants in the LGM was likely different (i.e. if C3 biomes were replaced with C4 vegetation, there in theory be a shift in 13C isotopes without a shift in biomass on land).
  • (4) the 13C record differs from one area of the ocean to the next - this likely reflects changes in paleo-ocean circulation/ biological pump (more on this later).
slide16

Carbon species in seawater

Dissolved CO2 pCO2 (or as it is more correctly expressed [H2CO3] ) is a minor constituent of seawater carbon ~1%

Bicarbonate ion (HCO3-) is ~90% of the carbon at ocean pH (8.2)Carbonate ion  (CO32-) is ~10% of the total carbon

Total Dissolved Inorganic C (TDIC) =H2CO3 + HCO3- + CO32-

Alkalinity (ALK) is the excess of cations over weak acid anionsIn seawater, and ignoring borate for the moment,

ALK is proportional to HCO3- + 2CO32-

Therefore, carbonate ion may sometimes be approximated as ALK - TDIC (in surface water)

The major chemical equilibrium we deal with is:CO2 +CO32- +H2O <==>2HCO3-

The equilibrium constant,

varies with temperature and salinity (and pressure)

slide17

TDIC (= *H2CO3   + HCO3- + CO32- )is influenced by three processes: (1) CO2 exchange with the atmosphere (2) photosynthesis/respiration (3) carbonate precipitation and dissolution

Alkalinity (Charge balance ~ HCO3- + 2CO32-) is influenced by: (1) carbonate precipitation and dissolution (2) organic matter formation and decomposition

(a small amount, through NO3- uptake and release)

seawater dic is primarily hco 3 and co 3 2
Seawater DIC is primarily HCO3- and CO32-

CO2(aq) increases at lower pH

slide19

Revelle Factor

CO2 increases by ~10%

when DIC increases by ~1%

Low latitudes have

Higher CO32-

And lower R factor

What does this mean?

CO2 +CO32- <==>2HCO3-

Increasing CO2 drives the reaction to the right, reducing CO32- but making more HCO3-

There is a lot of DIC in the ocean, converting one form to another does not change the total amount much; relative change is small

slide20

WHAT WILL BE THE IMPACT ON OCEAN CHEMISTRY AND ATMOSPHERIC CO2?

The change in land carbon actually added carbon to the atmosphere in the LGM;  some of that CO2 would dissolve immediately in the surface ocean, and ultimately be reflected in increased CO2 in deep waters.  The increased CO2 would cause dissolution of

carbonates in the deep sea (over a timescales of thousands of years).

DEEP WATER CHANGES IN CARBONATE CHEMISTRY

                          Interglacial Ocean         LGM                             LGM                                                          (before Calcite)      (after calcite dissolution)Alkalinity                2270 (meq/kg)           2270                    2322 (2270 + 52)Total CO2  (TDIC)   2085 (mmol/kg)         2115                     2141 (2115 + 26)CO32-                      129 (mmol/kg)            112129pCO2                       280 (matm)                 336                             296

DpCO2                                               +56                             +16

Adding or removing CO2 does not change alkalinity much (why not?)

500/35,600 is a 0.14% increase in atmosphere/ocean C –How much goes into the ocean (vs. atmosphere) depends on the Revellefactor. Adding a 500 Pg CO2 means about a 50 ppm rise in CO2 (with RF of 0.1)

Because the CO32-   is lower, the deep waters are undersaturated and CaCCO32-    will dissolve until equilibrium is re-established.

slide21

If we add 500 Pg C to the atmopshere, how much will by the surface ocean and how much will remain in the atmosphere?

Revelle factor (DpCO2/pCO2)/(DDIC/DIC) ~10

If you equilibrate with just the surface ocean (~1020 PgC)

DpCO2 = pCO2* 10 *(DDIC/DIC); DpCO2 = 6(DDIC)

For the deep ocean (38,000 PgC = DIC); DpCO2 = 0.11DDIC

But mass balance says DDIC = 500PgC – DpCO2

So

for pCO2 = 480 (LGM) and DIC = 1020;

DpCO2 (1+1/6) = 500; DpCO2 = 430 PgC

For DIC = 38,000 (i.e. equilibrate with whole ocean),

DpCO2 (1+1/.11) = 500; DpCO2 = 50 PgC

slide22

Negative feedback – precipitation rate of CaCO2 in the ocean

(the depth of the lysocline). Buffers changes in deep ocean CO3--

Solubility Ksp = [Ca+2][CO32-]; Ksp is dependent on pressure, temperature

(increases with pressure – so that carbonate formed in the surface ocean will dissolve at depth)

Le Chatlier’s rule – if you decrease[CO32-] in deep water in contact (equilibrium) with CaCO3 in sediments, you will dissolve carbonate until equilibrium is reestablished)

slide23

The bottom line:  A smaller biosphere in the LGM means HIGHER CO2 (by about 16 ppm if the biosphere lost 500 PgC to the atmosphere/ocean).  An even smaller biosphere (as has been proposed by those making estimates from paleoecology) means an even higher LGM pCO2)

SUMMARY WITH TEMPERATURE/SALINITY CHANGES:

Terrestrial C decrease                   +15 ppmOcean cooling                               -30  ppmOcean salinity increase                 +6.5 ppm

Total                                              -8.5 ppm

SOMETHING ELSE IS NEEDED TO EXPLAIN GLACIAL-INTERGLACIAL CO2 CHANGE!

biological pump
Biological ‘pump’
  • 12C preferentially taken up by phytoplankton
  • surface waters (and shells) enriched in 13C

12C enriched from oxidation of organic matter

slide25

d13C of DIC in seawater

Surface water

Photosynthesis preferentially removes 12C, leaves behind water enriched in 13C

Deep water – also along ‘conveyor’

Remineralization of organic matter adds 12C enriched material, lowering d13C

Ocean 13C

Efiiciency of the biological pump can be reflected in the difference in 13C between surface and deep water. There is therefore (or should be) a relationship between 13C and CO32- ion content of deep water

a proxy for the biological pump
A proxy for the biological pump?
  • Surface – deep water d13C (preserved in foram shells) is a measure of the strength of the biological pump
  • Glacial periods = Larger difference = stronger pump
  • More C stored in deep sea
  • But some problems:
    • Other sources of d13C variability
    • Foram d13C is complicated…
    • Increased C pumping should decrease deep ocean [CO3]2-, but no evidence for shallower lysocline
ocean circulation at the lgm
Ocean circulation at the LGM
  • Changes in Atlantic circulation have been linked to past climate changes (glacial-interglacial and abrupt)
  • In modern Atlantic , a net oceanic heat transport from North to South. If we perturb this transport, we alter climate

Modern ocean circulation can be visualized using Wally Broecker’s ocean conveyor…

14 c in dic and doc in the deep conveyor

Bomb

14C

SS

NCP

SOce

14C in DIC and DOC in the Deep Conveyor

A measure of the time since deep water equilibrated with the atmosphere

-525 to -390‰

Williams and Druffel, 1987; Bauer et al. 1992;

Druffel and Bauer, 2000

slide30
The ‘age’ of carbon increases from deep Atlantic to deep Pacific (this is where the ‘conveyor’ idea came from)

2050 – 670 = 1380 yr

5980 – 3160 = 2820 yr

possible mechanisms
Possible mechanisms…
  • Stronger overturning of Antarctic intermediate waters could have delivered more nutrients to surface waters & increased biological pump
  • Polar alkalinity hypothesis

**Remember: CO2 + CO32- + H2O  2HCO3-

    • Today: NADW dissolves little CaCO3 and upwells in S. Ocean with low [CO32- ],leaving S. ocean surface waters (and overlying atmosphere) with high CO2
    • Glacial: Southern source waters with high CO2 (more corrosive) expanded , dissolved more CaCO3 ,and returned more CO32- to Antarctic surface waters.
      • Broecker and Peng, 1989 proposed that this could explain ~ 40 ppm decrease in atmospheric CO2, , but more recent sediment data does not support this…
it is likely that the carbonate system plays an important role though
It is likely that the carbonate system plays an important role though…
  • pCO2 in surface water is a function of both DIC & Alk
  • Changes in mean inventory of either would impact surface water, and hence, atmospheric pCO2

= HCO3- + 2CO32- + OH- - H+ …

= CO2(aq) + H2CO3 + HCO3- + CO32-

the answer likely lies in the southern ocean
The answer likely lies in the Southern Ocean
  • Co-evolution of Antarctic temperature & atmospheric CO2
  • Nutrients are currently underutilized
  • Southern ocean ventilates large volumes of ocean interior
  • Two mechanisms for changes in S. ocean nutrient utilization:
    • Physical changes could isolate deep waters from surface, limiting CO2 degassing
    • Biological changes due to increased Fe (and Si?) fertilization by dust (increased Corg:CaCO3 export)
summary
Summary
  • It is likely that glacial-interglacial CO2 changes require a variety of mechanisms to explain.
  • The current frontrunners include:
    • T & S changes (-20 to 30 ppm)
    • Southern ocean mechanisms (major contributor)
  • Certain mechanisms (i.e. changes in whole ocean [CO32-] )seem unlikely due to disagreement with available proxy data (which is admittedly scarce)
  • Much work remains to be done to resolve this!
carbon cycle part ii
Carbon Cycle Part II

What is the fate of CO2 we add to the atmosphere by fossil fuel burning and land use?

slide37

http://www.esrl.noaa.gov/gmd/obop/mlo/programs/esrl/ccg/img/img_global_co2.jpghttp://www.esrl.noaa.gov/gmd/obop/mlo/programs/esrl/ccg/img/img_global_co2.jpg

slide38

Where does the other ~40% go???

Also, what happens to CO2 from deforestation (not counted here)

Source: Ralph Keeling, SIO

slide39

Deforestation: Clearing of forests (formerly in the US, now in the tropics)

Responsible for ~40% of total C emissions since 1850

In 1990s 0.5 to 2 GtC/year (8-25% of total emissions)

slide40

2000-2009

(PgC)

10

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

5

Sink

10

1950

2000

1900

1850

Time (y)

Human Perturbation of the Global Carbon Budget

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide41

1000

Latin America

800

S & SE Asia

Tropical Africa

600

400

CO2 emissions (Tg C y-1)

200

0

-200

1850

1860

1870

1880

1890

1900

1910

1920

1930

1940

1950

1960

1970

1980

1990

2000

2010

Time (y)

Emissions from Land Use Change (2000-2009)

R.A. Houghton 2010, personal communication; GFRA 2010

slide42

Global Fire Emissions Database (GFED) version 3.1

1400

1200

1000

Fire Emissions from

deforestation zones(Tg C y-1)

America

800

Africa

Asia

600

Pan-tropics

400

200

0

01

99

1997

07

05

2003

2009

Time (y)

Fire Emissions from Deforestation Zones

van der Werf et al. 2010, Atmospheric Chemistry and Physics Discussions

use of remote sensing to determine area deforested leads to reduced estimates of co 2 emissions
Use of remote sensing to determine area deforested leads to reduced estimates of CO2 emissions

Estimates for the 1990’s

Ref. 106 ha a-1PgC a-1

Houghton (FAO) 15.5 2.2(±0.8)

DeFries 5.6 0.9(±0.4) E

Van der Werf et al. 2009 Nature Geoscience

slide44

2000-2009

(PgC)

10

fossil fuel emissions

7.7±0.5

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

5

Sink

10

1950

2000

1900

1850

Time (y)

Human Perturbation of the Global Carbon Budget

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide45

2000-2009

(PgC)

10

fossil fuel emissions

7.7±0.5

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

5

Sink

10

1950

2000

1900

1850

Human Perturbation of the Global Carbon Budget

Time (y)

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide46

2000-2009

(PgC)

10

fossil fuel emissions

7.7±0.5

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

atmospheric CO2

4.1±0.1

5

Sink

10

1950

2000

1900

1850

Time (y)

Human Perturbation of the Global Carbon Budget

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide48

Suess Effect: Fossil fuel-driven depletion of atmospheric D14C

SUESS HERADIOCARBON CONCENTRATION IN MODERN WOOD, SCIENCE, 122 (3166): 415-417 1955

Jacobson [2000]

slide49

Fossil fuel has d13C of -21 to -27 per mil

If all emissions are taken up by the biosphere, the d13C of atmospheric CO2 should not change. Dissolution in the ocean preferentially removes 12C more than 13C, so we would expect a decline in 13C of atmospheric CO2

http://scrippsco2.ucsd.edu/graphics_gallery

land ocean sinks from 13 c
Land/Ocean sinks from 13C
  • The basic equation

C3 ~ 20‰C4 ~ 4.4‰O ~ 2‰

  • A terrestrial sink makes the atmosphere heavier ( more enriched in d13C)
  • An ocean sink has little effect on atmospheric 13C
  • A C4 sink looks like ocean to the atmosphere
  • As with CO2, the atmosphere shows signs of terrestrial uptake, mainly at high latitudes.
  • But the “disequilibrium” problem makes the interpretation of 13C very challenging.
slide51

Suess Effect: The Pre-Bomb Depletion of Atmospheric D14C by Fossil Fuels Also Applied to the Depletion of Atmospheric d13C by Fossil Fuels

d13C (per mil)

CO2 (ppm)

SUESS HERADIOCARBON CONCENTRATION IN MODERN WOOD SCIENCE 122 (3166): 415-417 1955

Francey et al. [1999]

the terrestrial sink from the n s co 2 gradient
The Terrestrial Sink from the N-S CO2 gradient

NOAA/CMDL Latitudinal Distribution of Carbon Dioxide

  • The observed gradient is shallower than expected from the distribution of fossil fuel and land use in atmospheric models.
  • Tans et al. 1990
  • W-E mixing is so rapid that trace gas gradients are very difficult to detect.
  • Need a gradient to infer regional sources/sinks

Conway, et al. [1994]

http://www.aos.princeton.edu/WWWPUBLIC/andyj/gv04.mpg

slide54

Fossil fuel has d13C of -21 to -27 per mil

If all emissions are taken up by the biosphere, the d13C of atmospheric CO2 should not change. Dissolution in the ocean preferentially removes 12C more than 13C, so we would expect a decline in 13C of atmospheric CO2

http://scrippsco2.ucsd.edu/graphics_gallery

land ocean sinks from 13 c1
Land/Ocean sinks from 13C
  • The basic equation

C3 ~ 20‰C4 ~ 4.4‰O ~ 2‰

  • A terrestrial sink makes the atmosphere heavier ( more enriched in d13C)
  • An ocean sink has little effect on atmospheric 13C
  • A C4 sink looks like ocean to the atmosphere
  • As with CO2, the atmosphere shows signs of terrestrial uptake, mainly at high latitudes.
  • But the “disequilibrium” problem makes the interpretation of 13C very challenging.
the d 13 c isotopic disequilibrium
The d13C Isotopic Disequilibrium

Gba

Gab

time

-6.5

tb

Atm. d13C (‰)

Isotopic Disequilibrium

-8.0

tb = Mean Residence Time

slide57

Decline in O2 is faster than increase in CO2

Stoichiometry says O2/CO2 for fossil fuel burning/biosphere should be ~-1.1

http://scrippso2.ucsd.edu/plots

slide58

Seasonal cycle in O2 in the southern hemisphere reflects marine biosphere activity and faster equilibration of the surface ocean for O2 compared to CO2

http://scrippso2.ucsd.edu/plots

slide59

Observation: O2 decline in the atmosphere is faster than expected from CO2 increase alone

Ocean uptake – why is the slope zero?

Fossil fuel burning

Slope is -1.1 mole O2 consumed per mole CO2produced

Biosphere uptake, loss will have the same slope

Outgassing – as the ocean warms, what happens to the solubility of O2?

Land uptake

Slope is +1.1 mole O2 produced per mole C removed from the atmosphere by plants

IPCC

Based on on Keeling 1996

slide60

GLOBAL BIOGEOCHEMICAL CYCLES, VOL. 19, GB4017, doi:10.1029/2004GB002410, 2005

Bender et al. Atmospheric O2/N2 changes, 1993–2002: Implications for the partitioning of fossil fuel CO2 sequestration

Ocean average uptake about 2 PgC/yr

slide63

2000-2009

(PgC)

10

fossil fuel emissions

7.7±0.5

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

atmospheric CO2

4.1±0.1

5

Sink

ocean

2.3±0.4

ocean

(5 models)

10

1950

2000

1900

1850

Time (y)

Human Perturbation of the Global Carbon Budget

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide64

2000-2009

(PgC)

10

fossil fuel emissions

7.7±0.5

5

Source

deforestation

CO2 flux(PgC y-1)

1.1±0.7

atmospheric CO2

4.1±0.1

land

5

2.4

(Residual)

Sink

ocean

2.3±0.4

(5 models)

10

1950

2000

1900

1850

Time (y)

Human Perturbation of the Global Carbon Budget

Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

slide65

Forest Cover in Massachusetts 1830 to 1985

  • Processes on Land that could be taking up the residual carbon:
  • Regrowth of some forests that were previously cut
  • Thickening of forests because of forest fire suppression
  • Increase of woody vegetation in dry regions due to better water use efficiency
  • Fertilization of forests by increased CO2

Foster, Motzkin and Slater 1998

slide66

A negative feedback: CO2 fertilization

Effect has been assessed with FACE (Free Air CO2 Enrichment) studies

No studies yet in the tropics – most are in temperate forests or low stature vegetation (crops). Strong response in some lianas.

b factors used in models are generally larger than 0.2; most models currently overpredict C storage in the future

slide67

Carbon storage potential depends on the residence time of carbon

How long will it take for respiration to catch up to increased production?

Plant +rhizosphere respiration

Photosynthesis

C respired

leaf

Microbial respiration

Time since C fixed

Allocation

Respiration

Fire

stem

Plant and

Root

Respiration

< years

years-

centuries

Litter and SOM

decomposition

> centuries

storage

root

Microbial

community

Stabilized SOM

Loss by leaching, erosion, weathering consumption

slide68

Photosynthesis ~30

Total ecosystem respiration ~30

(5 – 12 yr)

Total autotrophic respiration ~23.7

(0.01-1 yr)

Total heterotrophic

respiration ~6.3

(25-55 yr)

Time lag between photosynthesis and decomposition

Mean age of dying wood (model)*

Wood 2.0 70-115 yr*

Litter 3.3 2-3 yr incubations

Fluxes from Chambers et al. 2004 Ecol. Applications.

(MgC ha-1yr-1)

Root/SOM 1.0 3-10 yr incubations

Litter and wood

decomposition

Root

respiration

Root and soil organic

matter decomposition

* Vieira et al. 2006

slide69

Storage potential in soil and wood with CO2 fertilization

Rate of increase and time lag between increase in inputs and increase in outputs determine rate of C storage

Inputs increase with pCO2:

1+b*(ln(pCO2/278);

b = 0.2

Just increasing productivity is not enough to explain permanent plot observations of C gain of

~ 0.5 Mg C ha-1 a-1

Mg C ha-1 yr-1

See also Chambers and Silver 2006

slide70

Detecting Forest Disturbance with Multispectral Imagery

Spectral mixture analysis (SMA) for forested areas using image-derived endmember spectra for green vegetation (GV), non-photosynthetic vegetation (NPV), soil, and shade in a linear mixture model

~250 ha blowdown

Landsat sub-image from 2001 image – west bank of Rio Negro north of Manaus

slide71

Developing relationships between remote sensing metrics and field-based mortality rates

Each point represents a randomly placed 400 m2 inventory plot.

slide72

Carbon Balance and Catastrophic Mortality

100 ha run

Above carbon balance line sink, below line source

Large loss of carbon immediately following large mortality event

Afterwards a small sink for many decades

Overall carbon balance in aand b equal (0)

only background mortality

TLW carbon balance (Mg C ha-1 yr-1)

20%

mortality

events

carbon budget 1750 2008
Carbon Budget (1750-2008)

Large uncertainty

Land uptake (solve by difference)

Release by

Land use

Term we

know pretty

well

Dissolves in oceans

Gigatons of C per year

Fossil fuel

emission

Increase in

atmospheric

CO2

Terms we

know well

Added to atmosphere Where it goes

carbon budget 2000 2008
Carbon Budget (2000-2008)

Release by

Land use

Large uncertainty

Land uptake (solve by difference)

Term we

know pretty

well

Dissolves in oceans

Gigatons of C per year

Fossil fuel

emission

Increase in

atmospheric

CO2

Terms we

know well

Added to atmosphere Where it goes

what will be the fate of fossil fuel co 2
What will be the fate of fossil fuel CO2?
  • Revellefactor (see previous calculations – short term, add 500 Pg, increase atmosphere +56 ppm; long term 15 ppm)
  • Controls on different timescales

dissolution in surface ocean (pH concerns)

transport by biological pump into deep ocean

thermohaline circulation into deep ocean

dissolution of CaCO3 in ocean sediments

increased weathering

slide76

The biggest uncertainty in prediction of future climate is what we do:

Slide from Hansen

http://www.columbia.edu/~jeh1/SierraStorm.09Jan2007.pdf

energy increase from greenhouse gases is 2 5 watt m 2
Energy increase from greenhouse gases is 2.5 Watt/m2

A Christmas tree mini-light bulbs is 2.5 Watts

Imagine bulbs hung on a 1-meter grid everywhere around the globe

Bulbs burn 24 hours a day

CO2 responsible for about 50% of this radiative forcing; the rest is methane, nitrous oxide and other hydrocarbons including CFCs

CO2

CH4

N2O

slide80

Temperature has risen by 1.4 °F

(1 °F in the last 30 years)

9 of the hottest years of the century occurred in last 10 years (18 in the last 20 years)

slide81

We live in a time of abrupt climate change

Projections of global average surface temperature show we are heading for a climatic state far outside the range of variation of the last 1000 years.

We are already out of the range of CO2 for the last 800,000 years

slide82

Orr et al.

pH will change as pCO2 increases

predictions of future land c balance
Predictions of future Land C balance

All models use CO2 fertilization (negative) and warming/enhanced decomposition (positive) feedbacks; Differences between models reflect different predictions in climate as well as parameterization of these feedbacks

C Uptake

C Loss

Friedlingstein et al. 2006

slide84

Feedbacks – net short-term effect will be Positive (as CO2 increases, capacity to absorb CO2 decreases