Biogeochemical cycles I carbon and oxygen

Biogeochemical cycles I carbon and oxygen9:00 – 10:30 Wednesday Dec 8

https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdfhttps://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS3.pdf

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

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

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 youMore detail on the ocean biogeochemical

cyclingWhat is the long-term fate of CO2 we add to the

atmosphere?

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

ATMOSPHERIC CO2

640 X1015 g C

LIVING BIOMASS

830 X1015 g C

DISSOLVED ORGANICS

1500 X1015 g C

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 X1015 g C

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

25,000,000 X1015 g C

Distribution of Carbon;

1015 grams =

1 Petagram (Pg)

Response times are seasons to centuries

Response times are centuries to millennia

Response times are tens of thousands to millions of years

Long term control- balance of weathering rate and CO2

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?• 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?

• 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 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…

• Cooler SSTs = increased CO2 solubility; 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

Crowley et al., 1997, JGR

Was it stored on land?

Crowley et al., 1997, JGR

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

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

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 0.4 per 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…

http://www2.ocean.washington.edu/oc540/lec01-28/

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)

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

(1) differences in 13C between coexisting benthic species (vital effects) coupled with selective dissolution

(2) 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).

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- + CO3

2-

Alkalinity (ALK) is the excess of cations over weak acid anions In 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 + CO3

2- +H2O <==> 2HCO3-

The equilibrium constant,

varies with temperature and salinity (and pressure)

232

2

3

COpCOKHCOKc

H

TDIC (= *H2CO3 + HCO3- + CO3

2- ) is influenced by three processes:

(1) CO2 exchange with the atmosphere (2) photosynthesis/respiration

(3) carbonate precipitation and dissolution

Alkalinity (Charge balance ~ HCO3- + 2CO3

2- ) 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 HCO3- and CO3

2-

CO2(aq) increases at lower pH

Revelle Factor

Low latitudes haveHigher CO3

2- And lower R factor

CO2 increases by ~10% when DIC increases by ~1%

What does this mean?

CO2 + CO32- <==> 2HCO3

-

Increasing CO2 drives the reaction to the right, reducing CO3

2- 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

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 ofcarbonates 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) 112 129

pCO2 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 Revelle factor. 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 CaCCO3

2- will dissolve until equilibrium is re-established.

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

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[CO3

2-] in deep water in contact (equilibrium) with CaCO3 in sediments, you will dissolve carbonate until equilibrium is reestablished)

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 ppm Ocean cooling -30 ppm Ocean salinity increase +6.5 ppm

Total -8.5 ppm SOMETHING ELSE IS NEEDED TO EXPLAIN GLACIAL-INTERGLACIAL CO2 CHANGE!

Biological ‘pump’

•12C preferentially taken up by phytoplankton• surface waters (and shells) enriched in 13C

12C enriched from oxidation of organic matter

Ocean 13C

d13C of DIC in seawater

Surface waterPhotosynthesis preferentially removes 12C, leaves behind water enriched in 13C

Deep water – also along ‘conveyor’Remineralization of organic matter adds 12C enriched material, lowering d13C

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 CO3

2- ion content of deep water

Possible mechanism: Increased nutrient utilization (or supply)?

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• 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…

14C in DIC and DOC in the

Deep Conveyor

Williams and Druffel, 1987; Bauer et al. 1992;Druffel and Bauer, 2000

SS

SOce

NCP

-525 to -390‰

-600 -400 -200 0 2000

1000

2000

3000

4000

5000

6000

² 14C (‰)

Depth (m)

DICDOC

NCPSoceSS

Bomb14C

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

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

Radiocarbon

D14CF (approx) age (years)

Dissolved inorganic carbon (DIC) Deep Atlantic -80 0.92 670Deep Pacific -225 0.775 2048Dissolved organic carbon (DOC) Deep Atlantic -325 0.675 3157Deep Pacific -525 0.475 5980

2050 – 670 = 1380 yr

5980 – 3160 = 2820 yr

Possible mechanisms…1. Stronger overturning of Antarctic intermediate waters could have

delivered more nutrients to surface waters & increased biological pump

2. Polar alkalinity hypothesis **Remember: CO2 + CO3

2- + H2O 2HCO3-

– Today: NADW dissolves little CaCO3 and upwells in S. Ocean with low [CO3

2- ],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 CO3

2- 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…

• 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

= HC

O3- +

2CO

32- +

OH- -

H+ …

= CO2(aq) + H2CO3 + HCO3- + CO3

2-

The answer likely lies in the Southern Ocean

Two mechanisms for changes in S. ocean nutrient utilization:Physical changes could isolate deep waters from surface, limiting CO2 degassingBiological changes due to increased Fe (and Si?) fertilization by dust (increased Corg:CaCO3 export)

• Co-evolution of Antarctic temperature & atmospheric CO2

• Nutrients are currently underutilized

• Southern ocean ventilates large volumes of ocean interior

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 [CO3

2-] )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

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

http://scrippsco2.ucsd.edu/graphics_gallery

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

Source: Ralph Keeling, SIO

Where does the other ~40% go???Also, what happens to CO2 from deforestation (not counted here)

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)

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

Human Perturbation of the Global Carbon Budget

Sink

Sour

ce

Time (y)

5

10

10

5

1850 1900 1950 2000

1.1±0.7deforestation

CO2 f

lux

(PgC

y-1)

2000-2009(PgC)

Emissions from Land Use Change (2000-2009)

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

-200

0

200

400

600

800

1000

1850

1860

1870

1880

1890

1900

1910

1920

1930

1940

1950

1960

1970

1980

1990

2000

2010

Latin AmericaS & SE AsiaTropical Africa

CO2 e

miss

ions

(Tg

C y-1

)

Time (y)

Fire Emissions from Deforestation Zones

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

Fire

Em

issio

ns fr

om

defo

rest

atio

n zo

nes (

Tg C

y-1)

Global Fire Emissions Database (GFED) version 3.1

0

200

400

600

800

1000

1200

1400

1997 99 01 2003 05 07 2009

AmericaAfricaAsiaPan-tropics

Time (y)

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

Van der Werf et al. 2009 Nature Geoscience

Ref. 106 ha a-1 PgC a-1

Houghton (FAO) 15.5 2.2(±0.8)DeFries 5.6 0.9(±0.4) E

Estimates for the 1990’s



5

10

10

5

1850 1900 1950 2000

7.7±0.5

deforestation

fossil fuel emissions

Sink

Sour

ce

Time (y)

CO2 f

lux

(PgC

y-1)

1.1±0.7

2000-2009(PgC)


Time (y)


5

10

10

5

1850 1900 1950 2000

deforestation


Sink

Sour

ce

CO2 f

lux

(PgC

y-1) 7.7±0.5

1.1±0.7

2000-2009(PgC)



5

10

10

5

1850 1900 1950 2000

4.1±0.1


deforestation

atmospheric CO2

Sink

Sour

ce

Time (y)

CO2 f

lux

(PgC

y-1) 7.7±0.5

1.1±0.7

2000-2009(PgC)

Suess Effect: Fossil fuel-driven depletion of atmospheric D14C

Jacobson [2000]

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


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

Land/Ocean sinks from 13C• The basic equation

DC3 ~ 20‰DC4 ~ 4.4‰DO ~ 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.

Suess Effect: The Pre-Bomb Depletion of Atmospheric D14C by Fossil Fuels

Also Applied to the Depletion of Atmospheric d13C by Fossil Fuels

Francey et al. [1999]

350

340

330

320

310

300

290

280

198019601940192019001880186018401820180017801760174017201700

-7.8

-7.6

-7.4

-7.2

-7.0

-6.8

-6.6

-6.4

d13C

( pe r

mil)

CO2 (p

pm)

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

The Terrestrial Sink from the N-S CO2 gradient

• 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

NOAA/CMDL Latitudinal Distribution of Carbon DioxideConway, et al. [1994]

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




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

Land/Ocean sinks from 13C• The basic equation

DC3 ~ 20‰DC4 ~ 4.4‰DO ~ 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 d13C Isotopic Disequilibrium

Atm. d13C (‰)

time

Gba Gab

Isotopic Disequilibrium

tb

tb = Mean Residence Time

-6.5

-8.0

http://scrippso2.ucsd.edu/plots

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

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

Fossil fuel burningSlope is -1.1 mole O2 consumed per mole CO2 producedBiosphere uptake, loss will have the same slope

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

Ocean uptake – why is the slope zero?

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

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

IPCCBased on on Keeling 1996

GLOBAL BIOGEOCHEMICAL CYCLES, VOL. 19, GB4017, doi:10.1029/2004GB002410, 2005Bender et al. Atmospheric O2/N2 changes, 1993–2002: Implications for the partitioning of fossil fuel CO2 sequestration

Ocean average uptake about 2 PgC/yr

Sabine et al.

Total uptake since 1900 = 118 Pg



5

10

10

5

1850 1900 1950 2000

atmospheric CO2


deforestation

ocean2.3±0.4

oceanSink

Sour

ce

Time (y)

CO2 f

lux

(PgC

y-1)

(5 models)

4.1±0.1

7.7±0.5

1.1±0.7

2000-2009(PgC)



5

10

10

5

1850 1900 1950 2000

2000-2009(PgC)

atmospheric CO2

oceanland


deforestation

(Residual)

Sink

Sour

ce

Time (y)

CO2 f

lux

(PgC

y-1)

2.3±0.4(5 models)

4.1±0.1

7.7±0.5

1.1±0.7

2.4

Foster, Motzkin and Slater 1998

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

66

A negative feedback: CO2 fertilization

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

Effect has been assessed with FACE (Free Air CO2 Enrichment) studiesNo studies yet in the tropics – most are in temperate forests or low stature vegetation (crops). Strong response in some lianas.

Photosynthesis

leaf

stem

root

Allocation

storage

Microbial community

Stabilized SOM

Litter and SOMdecomposition

RespirationPlant andRootRespiration

Fire

Loss by leaching, erosion, weathering consumption

< years

years-centuries> centuries

Time since C fixed

C re

spire

d

Plant +rhizosphere respiration

Microbial respiration

Carbon storage potential depends on the residence time of carbon How long will it take for respiration to catch up to increased production?

Wood 2.0 70-115 yr*

Total heterotrophic respiration ~6.3

(25-55 yr)

Total autotrophic respiration ~23.7

(0.01-1 yr)

Total ecosystem respiration ~30(5 – 12 yr)

Litter 3.3 2-3 yr incubations

Litter and wooddecomposition

Root and soil organic matter decomposition

Rootrespiration

Root/SOM 1.0 3-10 yr incubationsFluxes from Chambers et al. 2004 Ecol. Applications.(MgC ha-1yr-1)

Photosynthesis ~30

Mean age of dying wood (model)*

* Vieira et al. 2006

Time lag between photosynthesis and decomposition

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

1900 1930 1960 19900.00

0.02

0.04

0.06

0.08

0.10

Vegetation sink

Soil sink

Mg

C ha

-1 y

r-1

1+b*(ln(pCO2/278);b = 0.2

Inputs increase with pCO2:

Just increasing productivity is not enough to explain permanent plot observations of C gain of ~ 0.5 Mg C ha-1 a-1

See also Chambers and Silver 2006

Detecting Forest Disturbance with Multispectral Imagery

~250 ha blowdown

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

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

Each point represents a randomly placed 400 m2 inventory plot.

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

-2.0

-1.5

-1.0

-0.5

0.0

0.5

1.0

a

Carbon Balance and Catastrophic Mortality

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 a and b equal (0)

100 ha run

-2.0

-1.5

-1.0

-0.5

0.0

0.5

1.0

800 900 1000 1100 1200 1300time (years)

b

20% mortality

eventsTLW

car

bon

bala

nce

(Mg

C h

a-1 y

r-1)

only background mortality

Fossil fuel emission Increase in

atmospheric CO2

Release by Land use

Terms we know well

Dissolves in oceans

Added to atmosphere Where it goes

Giga

tons

of C

per

yea

r

Term we know prettywell

Carbon Budget (1750-2008)

Large uncertainty

Land uptake (solve by difference)

Fossil fuel emission Increase in

atmospheric CO2

Release by Land use

Terms we know well

Dissolves in oceans

Added to atmosphere Where it goes

Giga

tons

of C

per

yea

r

Term we know prettywell

Carbon Budget (2000-2008)

Large uncertainty Land uptake

(solve by difference)

What will be the fate of fossil fuel CO2?

• Revelle factor (see previous calculations – short term, add 500 Pg, increase atmosphere +56 ppm; long term 15 ppm)

• Controls on different timescalesdissolution in surface ocean (pH concerns)transport by biological pump into deep ocean

thermohaline circulation into deep oceandissolution of CaCO3 in ocean sedimentsincreased weathering

Slide from Hansenhttp://www.columbia.edu/~jeh1/SierraStorm.09Jan2007.pdf

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

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 globeBulbs 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

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)

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

We live in a time of abrupt climate change

Orr et al. pH will change as pCO2 increases

Friedlingstein et al. 2006

C Uptake

C Loss

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

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

Documents

Biogeochemical cycles I carbon and oxygen