biogeochemical cycles i carbon and oxygen 9:00 – 10:30 wednesday dec 8
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Biogeochemical cycles I carbon and oxygen
9:00 – 10:30 Wednesday Dec 8
https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdfhttps://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS3.pdf
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…
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
COpCOK
HCOKc
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, 2000Williams and Druffel, 1987; Bauer et al. 1992;Druffel and Bauer, 2000
SSSS
SOceSOce
NCPNCP
-525 to -390‰
-600 -400 -200 0 2000
1000
2000
3000
4000
5000
6000
² 14C (‰)
Depth (m)
DICDOC
NCPSoceSS
-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 0.92 670
Deep Pacific -225 0.775 2048
Dissolved organic carbon (DOC)
Deep Atlantic -325 0.675 3157
Deep 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
= H
CO3- +
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 Asia
Tropical 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
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
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)
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
5
10
10
5
1850 1900 1950 2000
deforestation
fossil fuel emissions
Sink
Sour
ce
CO2 f
lux
(PgC
y-1) 7.7±0.5
1.1±0.7
2000-2009(PgC)
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
5
10
10
5
1850 1900 1950 2000
4.1±0.1
fossil fuel emissions
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
http://scrippsco2.ucsd.edu/graphics_gallery
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
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.
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
d13
C (p
er m
il)
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 Dioxide
Conway, et al. [1994]
http://www.aos.princeton.edu/WWWPUBLIC/andyj/gv04.mpg
http://scrippsco2.ucsd.edu/graphics_gallery
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
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 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
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
5
10
10
5
1850 1900 1950 2000
atmospheric CO2
fossil fuel emissions
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)
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
5
10
10
5
1850 1900 1950 2000
2000-2009(PgC)
atmospheric CO2
ocean
land
fossil fuel emissions
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 incubations
Fluxes 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 1300
time (years)
b
20% mortality
eventsTL
W c
arb
on b
alan
ce (
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
Gig
aton
s of
C p
er y
ear
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
Gig
aton
s of
C p
er y
ear
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 sediments
increased weathering
Slide from Hansen
http://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