Concluding Remarks

The apportionment of sources and sinks between ocean and land using <513C and CO2 records in the atmosphere is mainly dependent on isotopic disequilibrium terms, which have an impact on the atmospheric trend as large as the net fluxes themselves. At the global scale, isotopic

Table 14.1 Estimates of Carbon Emissions due to Biomass Burning (TgCyr ') and their Isotopic ¿I3C Signature (%<.PDB)

for the Major Ecosystems (TgC = r~sG+C)

Table 14.1 Estimates of Carbon Emissions due to Biomass Burning (TgCyr ') and their Isotopic ¿I3C Signature (%<.PDB)

for the Major Ecosystems (TgC = r~sG+C)

Average carbon emission by fires (TgCyr-1)

Average fire return interval (yr)


Average SI3C of burned biomass (%oPDB)

Tropical forest


250 (100-500)

Levin (1994)

(200-1200) (Amazonia alone)

Cochrane and Laurence (2002)

Potter et al. (2001)

Mixed forest

38 (11-60) USA


Leenhouts (1998)


17 (10-49) China


Wang et al. (2001)


8 Europe (mostly Mediterranean)


Mouillot et al. (2002)


22 Australia


Trabaud (1994)


Gill etal. (1997)

Boreal forest

40 (10-100) Canada


Chen et al. (2000)


50 (10-100) Russia

50 (25-100)

Shvidenko and Nilsson (2000)


Conard and Ivanova (1997)


1100 (500-1200) Africa


Barbosa et al. (1999)



82 India


Hao and Liu (1994)


550 South America


Seilerand Crutzen (1980)


100 (70-150) Australia



66 Central Asia






Andreae (1991)





Today's atmosphere is set to —8%o and ice core data have been used in this calculation. Averages and minimum/maximum values (in brackets when available) are presented.

disequilibria have an uncertainty range of 30% for the air-sea disequilibrium (range 47-60 GtC %oyr-1) and 50% for the soil-respired disequilibrium (range 19.8-33.6 GtC %o yr^1). Further, we have quantified other important disequilibrium and isofluxes, which are generally ignored in double deconvolutions:

(1) The replacement of C3 by c4 vegetation, which amounts to 50-100% of the soil respiration disequilibrium but has an opposite sign ( — 15 to —22 GtC%oyr~1).

(2) The (small) disequilibrium induced by biomass burning processes (1.7GtC%0yr-1).

(3) The disequilibrium of rock weathering processes (10.5 GtC %o yr-1).

(4) The correction to soil-respired disequilibrium due to non-C02 gas emissions and oxidation within the atmosphere (17 GtC %oyr-1).

At the regional level, some isotopic disequilibrium can become proportionally larger than the isoflux of net CO2 fluxes, especially at high northern latitudes for the aging of soil-respired CO2, and in the Tropics for the shifts from c3 to c4 vegetation. At even smaller scales, the ecosystems are probably never approaching isotopic equilibrium, neither over long timescales because of disturbances (biomass formed never has the same age as soil-respired carbon), nor on very short timescales (the isotopic composition of photosynthates does not equal that of respiration). Over time, interannual climate-induced fluctuations in the biospheric fractionation factor alter the interference of land and ocean sinks anomalies using atmospheric records of S13C and CO2 by less than 0.1 GtCyr-1. On the other hand, the c3 shift to c4 land use-induced disequilibrium should not change strongly from one year to the next; neither should the disequilibrium due to rock weathering. The soil-respired and air-sea isotopic disequilibrium should have a rate of interannual variability similar to one of gross fluxes, that is 10-20% globally, according to terrestrial and ocean model calculations. Priorities in future research should be to include all disequilibria in inverse atmospheric transport models to re-analyze regional fluxes. As recently performed for C0180 by Cuntz et al. (2003a,b) , it should also be important to incorporate 'on-line' the discrimination of 13C02 by canopy photosynthesis and its subsequent transport in the atmosphere in coupled land-surface-atmosphere models.

For 5180-C02, the atmospheric transport and the biospheric <5180-C02 isofluxes determine almost completely the atmospheric signal (see Table 14.2). S180-C02 has therefore a high potential to deduce the CO2 gross fluxes of the terrestrial biosphere, but the lsO isotopic exchange between water and C02 makes the 5180-C02 cycle more complex than the 513C cycle. On the other hand, ¿lsO in atmospheric C02 is determined by the gross fluxes of the carbon cycle compared to S13C that is determined

Table 14.2 Global <5180-C02 Isofluxes of Different Processes®


Isoflux (GtC %o yr-')






- — 1450

Ocean net flux


Ocean gross flux



Fossil fuel


Biomass burning







Carb. anhydrase


a Estimated from Peylin (1999) and Cuntz rf a£ (2003a,b).

a Estimated from Peylin (1999) and Cuntz rf a£ (2003a,b).

by the net fluxes. Isotopic disequilibria play therefore an important role in the i513C cycle but are of secondary importance for ¿>180-C02. The main unknown in the <5180-C02 cycle is leaf discrimination (A^) due to large uncertainties in its determining variables: mainly the relevant CO2 mixing ratio inside the leaf (Ca), the leaf water isotopic composition at the site of evaporation (Sf), and the carbonic anhydrase activity. Global estimates for leaf water isotopic composition range between 3 and 8.8%o VSMOW (Gillon and Yakir, 2001) and lead to an uncertainty in A^ of about 25%. This puts a limit to the accuracy at which the assimilation can currently be retrieved in double lsO deconvolutions. The stratosphere-troposphere exchange is the other important uncertain term in the ¿¡'^O-COy cycle. The stratosphere contributes most probably an isoflux of about 400 GtC %o yr~ 1 to the troposphere, which is approximately one fourth of the assimilation isoflux. Not taking it into account in double lsO deconvolutions would lead to a similar bias in the derived terrestrial biosphere CO2 gross fluxes to that caused by the uncertainty in A^. The contribution of the mass-independent enrichment of <$180-C02 in the stratosphere has not yet been incorporated in global models because of the low vertical resolution of the models and it was speculated that this could lead to the observed differences between modeled and measured <5180-C02 at observatories of marine background air (Cuntz et al., 2003b).

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