Isotopic evidence for large gaseous nitrogen losses from tropical rainforests

Abstract

The nitrogen isotopic composition (15N/14N) of forested ecosystems varies systematically worldwide. In tropical forests, which are elevated in 15N relative to temperate biomes, a decrease in ecosystem 15N/14N with increasing rainfall has been reported. This trend is seen in a set of well characterized Hawaiian rainforests, across which we have measured the 15N/14N of inputs and hydrologic losses. We report that the two most widely purported mechanisms, an isotopic shift in N inputs or isotopic discrimination by leaching, fail to explain this climate-dependent trend in 15N/14N. Rather, isotopic discrimination by microbial denitrification appears to be the major determinant of N isotopic variations across differences in rainfall. In the driest climates, the 15N/14N of total dissolved outputs is higher than that of inputs, which can only be explained by a 14N-rich gas loss. In contrast, in the wettest climates, denitrification completely consumes nitrate in local soil environments, thus preventing the expression of its isotope effect at the ecosystem scale. Under these conditions, the 15N/14N of bulk soils and stream outputs decrease to converge on the low 15N/14N of N inputs. N isotope budgets that account for such local isotopic underexpression suggest that denitrification is responsible for a large fraction (24-53%) of total ecosystem N loss across the sampled range in rainfall.

Fig. 1.

Conceptual model of controls on forest ¹⁵N/¹⁴N; soil N and plant N are boxed. I, nitrogen input flux to the forest; H and G, hydrologic leaching and gaseous fluxes, respectively; U and R, plant fluxes (uptake and return, respectively). ε_U, ε_H, and ε_G are effective (i.e., expressed) isotope effects for plant uptake and external hydrologic and gaseous losses, respectively [ε (in ‰) = (¹⁴k/¹⁵k − 1)·1,000, where k is the rate constant]. Under steady-state conditions, the N inputs into plants must balance the losses from it. Hence, the same flux-weighted δ¹⁵N that plants take up (U) is returned to the soil (R), resulting in no net change in soil δ¹⁵N by plant uptake processes. By contrast, losses of N to the external environment that fractionate against soil N isotopes (i.e., ε_H or ε_G) can lead to an elevation in soil δ¹⁵N relative to inputs. Consequently, at steady state, soil N isotopes can be modeled independently of plants: δ¹⁵N_soil = δ¹⁵N_inputs + ε_H·(H/(H + G)) + ε_G·(G/(H + G)). In this study, we test which of the following best explains the soil δ¹⁵N variation across a rainfall sequence of tropical forests: (i) the δ¹⁵N of inputs, (ii) N isotope discrimination during hydrologic leaching (ε_H), or (iii) discrimination during gaseous N loss (ε_G).

Fig. 2.

Nitrogen isotope ratios and chemistry of small streams and atmospheric inputs across the rainforests. (a and b) In stream water, the ¹⁵N/¹⁴N of TDN (a) and of NO₃⁻ (b) are shown. (c) Stream water N species. (d and e) In atmospheric N inputs (precipitation and cloud water), the ¹⁵N/¹⁴N of TDN (d) and of NO₃⁻ (e) are shown; a δ¹⁵N of 0‰ for N₂ fixation (11) is indicated in d. (f) Precipitation N species. Open gray symbols represent individual observations from 2001–2004 (see Supporting Materials and Methods); filled symbols connected by solid lines represent means ± 1 SE.

Fig. 3.

Nitrogen isotope ratios and chemistry in rainforest soils. (a) ¹⁵N/¹⁴N of the top 50 cm of bulk soil (14). (b and c) ¹⁵N/¹⁴N of TDN (b) and of NO₃⁻ (c) in soil solution from a 35-cm depth. (d) Soil solution N species. Symbols are the same as in Fig. 2. (e) Relationship between soil solution nitrate ¹⁵N/¹⁴N and ¹⁸O/¹⁶O across all sites. (f) ¹⁵N/¹⁴N of NO₃⁻ vs. the natural logarithm of soil NO₃⁻ concentration in a series of cores designed to eliminate processes other than denitrification, such as nitrification (by addition of N serve) and plant uptake (by removal of plant roots) (see Supporting Materials and Methods).

Fig. 4.

Isotopic expression of denitrification across sampling scales. The x axis is f, the fraction of nitrate remaining, which is the ratio of measured NO₃⁻ to the estimated initial NO₃⁻ concentration. We estimated the initial NO₃⁻ concentration and ¹⁵N/¹⁴N as follows: for the incubation experiment (see Supporting Materials and Methods and Fig. 3f), we used the highest extractable NO₃⁻ concentration and its corresponding ¹⁵N/¹⁴N; for soil extracts, we used the mean of the five highest NO₃⁻ concentrations and their corresponding ¹⁵N/¹⁴N; for lysimeters, we used the mean of the five highest soil water NO₃⁻ concentrations and their corresponding ¹⁵N/¹⁴N; for streams, we used the mean of the five highest stream water NO₃⁻ concentrations and their corresponding ¹⁵N/¹⁴N. The organism-level isotope effect line was calculated using an approximate form of the Rayleigh equation, δ¹⁵NO₃⁻ = δ¹⁵NO₃⁻ _initial − ε·(ln(f)) (20), and an intrinsic ε of 20‰ (23, 24).

Fig. 5.

Partitioning of N losses by isotope balance. (a) The fraction of total N loss by gaseous N using stream and lysimeter isotope data. The following equation was solved for the fraction of total N loss by gaseous N using stream and lysimeter isotope data: δ¹⁵N-TDN_deposition (i.e., 0.47‰) = δ¹⁵N_gas·(f_gas) + δ¹⁵N_TDN·(1 − f_gas), where f is the fraction of loss. In the “underestimate,” δ¹⁵N_gas was estimated as the weighted-mean δ¹⁵N of NO₃⁻ in stream water minus an ε of 20‰ (23, 24), and δ¹⁵N-TDN was taken from the stream water; in the “best estimate,” δ¹⁵N_gas was estimated as the weighted-mean δ¹⁵N of NO₃⁻ in deep soil waters minus an ε of 13.2‰, and δ¹⁵N-TDN was taken from the soil waters (Fig. 2f). (b) Estimated rates and forms of gaseous N loss. Fluxes were calculated by applying a linear regression for the relationship between rainfall and evapotranspiration to estimated stream flow (31). Also shown are in situ fluxes of NO_x and N₂O from ref. and P. A. Matson, personal communication. The N₂ flux is estimated as the difference between total gas flux based on the N isotope balance and the N₂O and NO_x fluxes measured by chambers.