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Human land cover change has altered an estimated 50% of Earth’s land surface, mostly through conversion of forest ecosystems to agricultural uses⁵. The IPCC Fifth Assessment Report (AR5) provides global radiative forcing values for the opposing effects of land use on CO₂ and surface albedo that have similar magnitudes but opposite sign: +0.17–0.51 W m⁻² (1850–2000) versus −0.15 ± 0.10 W m⁻² (1750–2011)¹. The sign of the global average surface temperature response to these forcings due to human land cover change remains controversial^{6, 7}. Thus, it is not known whether large-scale human deforestation has contributed to global warming or global cooling^{8, 9}.

Forest ecosystems return to the atmosphere about 1% of the annual carbon uptake by the land biosphere in the form of chemically reactive BVOCs (ref. 10). The ecological and physiological roles of BVOCs are broad, ranging from abiotic and biotic stress functions to integrated components of carbon metabolism. In the present-day climate state, the dominant BVOCs emitted are isoprene (400–600 TgC yr⁻¹) and monoterpenes (30–130 TgC yr⁻¹; ref. 10). Tropical and temperate vegetation tends to emit isoprene whereas boreal vegetation tends to emit monoterpenes. New-generation global BVOC emission models suggest that the BVOC emissions were about 20–25% higher in the pre-industrial climate state and that the cropland expansion is the major driver of the decrease in emissions over the past century^{2, 3, 4}. This high sensitivity to human deforestation arises because of the strong BVOC emission dependence on ecosystem type: broadleaf forests are strong emitters whereas crops and grasses are weak emitters or even non-emitting¹⁰.

BVOC emissions undergo rapid oxidation in the atmosphere, generating the climate pollutants O₃ and biogenic secondary organic aerosol (SOA; ref. 11). The photochemical processing of BVOC emissions influences the oxidation capacity of the atmosphere, which affects the lifetime of CH₄ and the production of other secondary aerosols, sulphate and nitrate, whose formation rates depend on the availability of oxidants¹². O₃ and CH₄ are powerful greenhouse gases that warm the atmosphere. Aerosol particles (organics, sulphate and nitrate) predominantly scatter solar radiation back to space and lead to global cooling. The extent to which the past changes in BVOC emissions have altered global radiative forcing is not clear-cut at the outset because the perturbation involves multiple warming and cooling climate pollutants.

Here, a global carbon–chemistry–climate model is employed to investigate the effects of the historical cropland expansion between the 1850s and 2000s on the BVOC emissions and climate pollutants. A baseline simulation representative of the chemical and meteorological background atmosphere in the 2000s is performed. Land cover and anthropogenic emissions are prescribed using harmonized gridded datasets that were developed for the IPCC AR5 assessment^{5, 13}. Between the 1850s and 2000s the global crop cover fraction of vegetated land area more than doubles (14–37%) at the expense of grass (11–6%), shrub (6–3%), savanna (8–5%), deciduous (13–10%) and tropical rainforest (12–10%) (Supplementary Fig. 1). A sensitivity experiment is performed in which the 2000s baseline simulation is forced with the 1850s land cover dataset. The difference between these simulations allows faithful isolation of the effects of the 1850s–2000s land cover change on BVOC emissions and the climate pollutants and is defined as the case LAND. The 1850s land forcing also affects the climate pollutants via non-BVOC pathways, including hydrological cycle impacts on oxidant and aerosol sources and sinks, underlying surface albedo effects on aerosol–radiation interactions, and surface deposition rates. Therefore, a second pair of simulations is performed identical to above (2000s baseline and the 2000s baseline forced with the 1850s land cover dataset) but with the BVOC emissions held fixed at climatological monthly varying values for the 2000s. The difference between this simulation pair is defined as the case LAND-fixbvoc. In this case, the climate pollutants respond only to the non-BVOC mechanisms of land cover change.

I follow the IPCC’s reporting structure and adopt the historical global mean annual average radiative forcing metric, which is a powerful indicator of the equilibrium global average surface temperature response to the perturbation. The global model is used to compute the radiative forcings due to O₃, biogenic SOA, sulphate and nitrate for the LAND and LAND-fixbvoc cases. CH₄ radiative forcing is calculated off-line using the simulated changes in CH₄ chemical lifetime, which accounts for the indirect CH₄ effects on stratospheric water vapour and the longer-term O₃ response¹⁴. For two reasons this study does not consider aerosol–cloud interactions. First, the sign of the global radiation interaction between SOA and cloud in the present-day atmosphere is not robust across models, with published estimates ranging from +0.23 W m⁻² to −0.77 W m⁻² (refs 15, 16). Second, there is evidence that isoprene, the dominant BVOC in the temperate and tropical biomes where the cropland expansion has occurred, acts to inhibit new particle nucleation¹⁷.

In the 2000s baseline, the simulated global source strength of BVOC emissions is 607 TgC yr⁻¹ (isoprene = 404 TgC yr⁻¹; monoterpene = 132 TgC yr⁻¹; other VOCs = 70 TgC yr⁻¹), which increases to 819 TgC yr⁻¹ (isoprene = 539 TgC yr⁻¹; monoterpene = 184 TgC yr⁻¹; other VOCs = 96 TgC yr⁻¹) when forced with the 1850s land cover dataset. Thus, the 1850s–2000s land cover change perturbation reduces the global BVOC emission source strength substantially by ~35%. This value is larger than the 20–25% calculated for pre-industrial to present day in previous studies because it represents the response to the land cover forcing only and does not take into account other global change factors, such as temperature and vegetation productivity, that tend to increase BVOC emissions since the pre-industrial and partially offset the reductions from deforestation. The CH₄ chemical lifetime in the 2000s simulation is 10.7 yr, which increases by six months to 11.2 yr in the 2000s simulation forced with 1850s land cover.

The LAND case results in global radiative cooling through O₃ (−0.13 W m⁻²) and CH₄ (−0.06 W m⁻²) and warming through biogenic SOA (+0.09 W m⁻²), as shown in Fig. 1. The largest single impact of the cropland expansion is through O₃. Human deforestation is responsible for substantially lower levels of O₃, CH₄ and biogenic SOA in the present-day global atmosphere than would exist if the cropland expansion had not occurred. It is well established that, even in the NO_x-limited regime, the O₃ production efficiency is VOC-dependent. Remarkably, the net atmospheric chemistry effect is a global radiative cooling of −0.11 ± 0.17 W m⁻², which is of similar magnitude to the land cover change effects on surface albedo and CO₂ compared in Fig. 1.

Climate pollutants are calculated in this study. Barred black lines indicate uncertainty range. Uncertainties for climate pollutants are assigned a best estimate of ±100% based on a factor of two uncertainty in BVOC emissions^{10, 24}. Values and uncertainty ranges for surface albedo and carbon dioxide release are from IPCC AR5 (ref. 1). The time period for each perturbation signal is 1850s–2000s, except for the surface albedo estimate, which refers to 1750–2011.

The Yale-E2 global carbon–chemistry–climate model is applied in this study. Yale-E2 is built around the NASA GISS ModelE2 global climate model²³ and includes an interactive land biosphere model²⁴. The Yale-E2 framework fully integrates the land biosphere–oxidant–aerosol system so that these components interact with each other and with the physics of the climate model. This study applies 2° × 2.5° latitude by longitude horizontal resolution with 40 vertical layers extending to 0.1 hPa. The atmospheric composition has been well tested against observations²⁵. The vegetation submodel is embedded within a general circulation model that provides the key meteorological drivers for the vegetation physiology. The land-surface hydrology submodel provides the grid-cell-level soil characteristics to the vegetation physiology. The vegetation cover is described using eight ecosystem types: tundra, grassland, shrub, savanna, deciduous, tropical rainforest, evergreen and crop. On-line oxidants affect aerosol production and on-line aerosols provide surfaces for chemical reactions and influence photolysis rates. The model includes NO_x-dependent biogenic SOA production from the oxidation of BVOC emissions (isoprene, monoterpene and other VOCs; ref. 26) using a two-product scheme that describes partitioning of semi-volatile compounds between the gas and aerosol phases depending on their volatility and pre-existing carbonaceous aerosol availability. O₃-initiated isoprene oxidation recycles the hydroxyl radical (OH) but does not represent any additional OH enhancements in forest environments because new evidence suggests that elevated OH field measurements may have been the erroneous result of interference inside the instrument chamber²⁷. Alkyl nitrate species that are formed during the oxidation process are 100% available for OH-oxidation to re-release NO_x and peroxy radicals²⁸.

Leaf isoprene emission is calculated as a function of electron-transport-limited photosynthesis, intercellular and atmospheric CO₂ concentration, and canopy temperature²⁴. Leaf-level monoterpene emissions are simulated by defining an ecosystem-specific basal rate that is modified using a temperature-dependent algorithm²⁹. Other VOCs emissions are parameterized as 44% of the isoprene emission rate, where 32% of other VOCs are assumed to generate biogenic SOA precursors²⁶. The vegetation biophysics module computes the photosynthetic uptake of CO₂ coupled with the transpiration of water vapour and the BVOC emissions at the 30-min physical integration time step of the global model. The global BVOC emission simulation has been extensively evaluated for conditions representative of the present-day climatic state and shown to reproduce 50% of the variance across different ecosystems and seasons in a global database of 28 measured campaign-average fluxes, and to capture authentically the observed variance in the 30-min average diurnal cycle (R² = 64–96%) at nine sites where the flux magnitude is simulated to within a factor of two (ref. 24).

The atmosphere-only model simulations apply decadal average (1996–2005) monthly-varying sea surface temperatures and sea ice from the HadSST2 dataset³⁰. Anthropogenic emissions of trace gas and aerosol emissions for 2000 are from the historical inventory developed for IPCC AR5 (ref. 13). Long-lived greenhouse gas and CH₄ concentrations are prescribed for the 2000s: CO₂ = 370 ppm; N₂O = 316 ppb; CH₄ = 1,752 ppb; CFC-11 = 263 ppt; CFC-12 = 534 ppt; other halocarbons = 363 ppt. Integrations of 22 model years are completed for all simulations (baseline and sensitivity runs); the first two years of the simulations are discarded as spin-up and the remaining 20 years are averaged for analyses. The simulations include the effects of feedbacks from physical climate change on reactive chemical composition but do not allow the on-line O₃ and aerosol changes to feed back to the radiation and dynamics.

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