Minimal Influence of AAN Deposition on the North Atlantic.

The average CS-δ15N in the Hog Reef coral core is 4.17‰, with a SD of 0.26‰ (Fig. 2A). This is 1.6–1.7‰ higher than the thermocline nitrate δ15N of the surrounding ocean (Fig. 1B), consistent with both prior work on Bermuda (29) and the global correlation between the δ15N of the N supply to surface waters and CS-δ15N (19). Despite the low-δ15N of AAN deposition at Bermuda (12, 13, 15, 22), the CS-δ15N record is remarkably stable. There appears to be a long-term decline in CS-δ15N of 0.6‰ from the early 20th century to the 1980s, with several superimposed decadal oscillations of 0.5–0.6‰ (see smoothed record in Fig. 3). One might argue that the long-term δ15N decline, although weak, was caused by increasing AAN deposition. However, since the 1980s, CS-δ15N has increased back to the earlier 20th century values, a change that has not been observed in terrestrial or ice core δ15N records (Fig. 2C). This suggests that factors other than AAN have dominated variations in Bermuda CS-δ15N over this time period. The decadal variability aside, the stability of CS-δ15N across the entire record suggests that the AAN can only have been a minor contributor to the N supply to the remote North Atlantic surface ocean over the 20th century.

Fig. 2. The Bermuda CS-δ15N record and its comparison with other records. (A) The Bermuda CS-δ15N record (red circles and black line, the latter being a 10-y running mean), with the CS-δ15N record from Dongsha Atoll in the South China Sea shown for comparison (gray line) (10). The Bermuda CS-δ15N record was compiled from two coral cores: the continuous coral core from Hog Reef (AD 1880–2014; solid red circles) and the fossil coral core from the John Smith’s Bay (AD 1780–1840; open red circles). Due to a problem in core archiving, only short intervals in the bottom part of the John Smith’s Bay coral were available for CS-δ15N analysis. (B) Comparison of CS-δ15N changes to modeled surface ocean δ15N changes (28) in the North Atlantic and South China Sea over the 20th century. (C) Greenland Summit ice core nitrate δ15N record (green line) (25) and power functions (black lines) fitted to δ15N records from 24 remote North American watersheds (23). The watershed δ15N records were normalized to each record’s preindustrial value (23). In both A and C, δ15N decreases upward, so that interpreted greater AAN importance is upward. (D) US N fertilizer consumption and NO x emissions (red and orange lines), compared with China fertilizer consumption and coal consumption, the latter intended as a rough indicator of combustion N sources (black and gray lines). Data in D were compiled by the Earth Policy Institute.

Fig. 3. Decadal variation in the Bermuda CS-δ15N record and the history of the North Atlantic Oscillation. (A) The Bermuda CS-δ15N data (black circles) and its 10-y running mean (red line). (B) The station-based North Atlantic Oscillation (NAO) index (light gray line) and its 10-y running mean (red line) (57). The CS-δ15N record shows maximum positive correlation (r = 0.70) with NAO when CS-δ15N lags NAO by roughly a decade (SI Appendix, Fig. S3).

AAN deposition can lower the δ15N of ocean ecosystems and thus corals in two related ways. First, in each year, AAN deposition represents a N input to the euphotic zone (the photosynthetically active upper layer) that is lower in δ15N than the nitrate supply from below and in situ N 2 fixation, thus working to lower the δ15N of organic N produced by phytoplankton (Fig. 4A). Second, operating over multiple years, the lowered δ15N of organic N sinks into the shallow subsurface and is remineralized, gradually lowering the δ15N of the subsurface nitrate (Fig. 4A). This subsurface nitrate is the dominant annual supply of fixed N to the surface ocean, and thus, its δ15N decrease is communicated back to the surface ocean ecosystem. We first address the δ15N of subsurface nitrate in the Sargasso Sea and then consider the annual N isotope budget of its euphotic zone (the sunlit surface layer).

Fig. 4. (A) The euphotic zone fixed nitrogen budget in the modern Sargasso Sea and the depth profiles of temperature, nitrate concentration, and nitrate δ15N near Bermuda during phases of (B) positive NAO and (C) negative NAO. On an annual basis, AAN deposition (with a δ15N of −5‰) represents less than 5% of N supply to the euphotic zone in the modern Sargasso Sea (SI Appendix, Fig. S2), with the vertical nitrate supply dominating the total fixed N supply and N 2 fixation (with a δ15N of −1‰) contributing a small fraction of N supply. The thermocline waters from which nitrate is supplied, in turn, exchange water with the deeper and higher-latitude waters, which is simplified in A as “mixing” with underlying deep water. During a positive phase of NAO (B), a thin STMW layer is generated in the subsurface, increasing the nitrate concentration and nitrate δ15N in the thermocline. During a negative phase of NAO (C), a thick STMW layer is generated in the subsurface, decreasing the nitrate concentration and nitrate δ15N in the thermocline. The temperature and nitrate concentration depth profiles are from ref. 51. The nitrate δ15N depth profile under positive NAO state is from Fig. 1, whereas the nitrate δ15N depth profile under negative NAO state (dashed yellow line) is based on a calculation (SI Appendix, Figs. S4 and S5).

In the modern Sargasso Sea, the subsurface nitrate δ15N at 200 m depth is 2.5‰ (Fig. 1) (16, 18), lower than the mean ocean nitrate δ15N of 5‰ (32). It has been suggested that N 2 fixation in the euphotic zone and subsequent sinking and remineralization of the resulting organic N is dominantly responsible for the deviation of thermocline nitrate δ15N from the mean ocean value (33, 34). However, because AAN at Bermuda has an even lower δ15N than newly fixed N, it could not be ruled out that AAN is an important contributor to the low δ15N of the nitrate in the Sargasso Sea thermocline (11). Similarly, it has been suggested that the low nitrate δ15N (of ∼3‰) in the Mediterranean Sea (35) is dominantly due to anthropogenic inputs (36).

However, the modern CS-δ15N is similar to the CS-δ15N in the late 19th century. Moreover, the fossil coral from the south shore of Bermuda (John Smith’s Bay; Fig. 1A) going back to approximately AD 1780 confirms this low δ15N (Fig. 2A), suggesting that the low nitrate δ15N in the thermocline water has been a persistent feature of the Sargasso Sea since before the Industrial Revolution. Because anthropogenic perturbation of the atmospheric reactive N cycle was very limited before the Industrial Revolution (7, 25), the stability of Bermuda CS-δ15N argues that natural processes such as N 2 fixation in the surface ocean, not AAN deposition, are the dominant drivers of the low δ15N in the nitrate of the modern Sargasso Sea thermocline. Qualitative expectations are that relatively rapid AAN deposition would be needed over many years to significantly change the δ15N of thermocline nitrate, due to the substantial burden of fixed N in the thermocline and its continuous exchange with the rest of the ocean’s vast nitrate reservoir (“mixing” in Fig. 4A). Thus, this finding is not particularly surprising.

In comparison with the δ15N of thermocline nitrate, the annual N isotope budget of the euphotic zone is more directly affected by a given rate of AAN deposition, and yet the stability of CS-δ15N argues that it also changed little over the 20th century. A simple isotope mixing calculation of this budget indicates that the 0.2‰ decrease in Bermuda CS-δ15N from the early 20th century to the year 2000 translates to only a 2% increase in the annual fixed N supply to the Sargasso Sea euphotic zone due to rising AAN deposition (SI Appendix, Fig. S2). One potential caveat in this calculation is that the rate of N 2 fixation in the North Atlantic might have decreased over the 20th century in response to increasing AAN input. In our isotope mixing model, when including suppression of N 2 fixation that perfectly compensates for AAN deposition, AAN deposition represents a maximum of 5% of total fixed N supply to the euphotic zone for the year 2000 (SI Appendix, Fig. S2). Thus, it appears that compensation by N 2 fixation is not dominantly responsible for the stability of the Bermuda CS-δ15N record; rather, AAN deposition increases must have been small relative to the nitrate supply from below (Fig. 4A).

The lack of a clear decline in CS-δ15N at Bermuda was not expected. A model study of the effect of the recent history of AAN deposition rate on the δ15N of the organic N produced in the upper ocean predicts a decline of 0.8‰ from the early 20th century to the year 2000 in the Sargasso Sea (28), greater than the observed 0.2‰ decline in the Bermuda CS-δ15N record over this time period (Fig. 2B). The 0.8‰ decline in the model occurred despite the inclusion of N 2 fixation suppression by AAN; a larger δ15N decline would have been simulated without it (28). The Bermuda record also contrasts with a CS-δ15N record from Dongsha Atoll in the South China Sea (10), which shows a δ15N decline of 0.7‰ by the year 2000 (Fig. 2A); the South China Sea CS-δ15N change is consistent with the model-based prediction for that region (Fig. 2B). In the South China Sea, the same isotope mixing model as applied to Bermuda suggests that by the year 2000, AAN deposition had increased the total N supply to the surface ocean by 17–23% (10) (SI Appendix, Fig. S2).

The difference of the Bermuda CS-δ15N record from ocean model predictions raises the possibility that ocean models have overestimated open-ocean AAN deposition. In general, the variation in atmospheric N deposition across models, and thus the reliance of the community on model ensembles (37), points to the substantial uncertainties in the model estimates. In terms of systematic bias, overestimation of AAN deposition has been suggested and related to ocean/atmosphere cycling of both ammonium/ammonia and dissolved organic N (27), for which there is isotopic and chemical evidence at Bermuda (13, 14). The Bermuda CS-δ15N record thus provides impetus for further investigation of the processes responsible for AAN delivery to the open ocean (27).

An alternative set of explanations for the discrepancy relates to oceanic transports of fixed N. Specifically, the global ocean models used for the simulations may underestimate the rate of nitrate supply to the euphotic zone from below, thus artificially raising the proportional importance of AAN deposition in the N budget of the euphotic zone. The existence of such a bias toward underestimating the subsurface nitrate supply is supported by a documented discrepancy between ocean models and observations for net primary productivity in the subtropical gyres of both the North Atlantic (at BATS) and the North Pacific (38). Such a bias might apply in the subtropical gyres but not in the South China Sea because the gyres do not experience the wind-driven upwelling that occurs in the South China Sea, so that more complex mechanisms of nutrient supply are proportionally more important in the gyres. Global ocean models are unlikely to adequately simulate these more complex mechanisms (17, 39).

The lack of clear CS-δ15N decline in the Bermuda coral core also contrasts with the strong δ15N declines observed in remote North American terrestrial records (>3‰ on average) and Greenland ice cores (>10‰) (23, 25, 26) (Fig. 2C). On an annual basis, both terrestrial and marine ecosystems receive most of their fixed N from internal recycling (40, 41). In the ocean, circulation and mixing supply subsurface nitrate that itself is the product of remineralization of sinking organic N. On land, soil organic N is decomposed and remineralized, making the N available to plants (40, 42); the freshwater bodies from which the terrestrial δ15N records were generated also receive N from soil organic matter remineralization (23). The rate of this recycled N supply appears to be comparable between land and ocean (40, 43), so a land-to-ocean difference in this term is probably not the primary driver of the distinction in the AAN δ15N signal between the terrestrial δ15N records and the coral CS-δ15N record (Fig. 2).

One cause for the greater response of the terrestrial (lake) δ15N records is probably simply their greater proximity to the anthropogenic N sources and thus their receipt of more AAN deposition (44). However, given the starkness of the distinction, it may have additional causes. For example, as discussed above, the δ15N of thermocline nitrate is buffered by continuous exchange with the global ocean nitrate reservoir and thus is expected to change only very gradually due to anthropogenic N accumulation in the subsurface (28). In soils, however, the source of the remineralized N flux from soil organic matter may be restricted to a small fraction of the total soil N, with the rest cycling so slowly as to be considered virtually unavailable (45). If so, the remineralized N flux to plants and runoff may decline in δ15N relatively rapidly as AAN accumulates in this small, actively cycled fraction of soil N. Putting aside this specific proposal, the contrast between the land and coral δ15N records promises to advance our understanding of the differences in N dynamics between land and ocean.