Significance The geological record contains numerous examples of “greenhouse periods” and ocean acidification episodes, where the spreading of corrosive (CO 2 -enriched) bottom waters enhances the dissolution of CaCO 3 minerals delivered to the seafloor or contained within deep-sea sediments. The dissolution of sedimentary CaCO 3 neutralizes excess CO 2 , thus preventing runaway acidification, and acts as a negative-feedback mechanism in regulating atmospheric CO 2 levels over timescales of centuries to millennia. We report an observation-based indication and quantification of significant CaCO 3 dissolution at the seafloor caused by man-made CO 2 . This dissolution is already occurring at various locations in the deep ocean, particularly in the northern Atlantic and near the Southern Ocean, where the bottom waters are young and rich in anthropogenic CO 2 .

Abstract Oceanic uptake of anthropogenic CO 2 leads to decreased pH, carbonate ion concentration, and saturation state with respect to CaCO 3 minerals, causing increased dissolution of these minerals at the deep seafloor. This additional dissolution will figure prominently in the neutralization of man-made CO 2 . However, there has been no concerted assessment of the current extent of anthropogenic CaCO 3 dissolution at the deep seafloor. Here, recent databases of bottom-water chemistry, benthic currents, and CaCO 3 content of deep-sea sediments are combined with a rate model to derive the global distribution of benthic calcite dissolution rates and obtain primary confirmation of an anthropogenic component. By comparing preindustrial with present-day rates, we determine that significant anthropogenic dissolution now occurs in the western North Atlantic, amounting to 40–100% of the total seafloor dissolution at its most intense locations. At these locations, the calcite compensation depth has risen ∼300 m. Increased benthic dissolution was also revealed at various hot spots in the southern extent of the Atlantic, Indian, and Pacific Oceans. Our findings place constraints on future predictions of ocean acidification, are consequential to the fate of benthic calcifiers, and indicate that a by-product of human activities is currently altering the geological record of the deep sea.

Seafloor dissolution of CaCO 3 minerals will constitute a primary feedback to ocean acidification over timescales of centuries to tens of millennia (1). The overall dissolution reaction is as follows: CaCO 3 + CO 2 + H 2 O → Ca 2 + + 2 HCO 3 − , [1]where CaCO 3 denotes solid carbonate in bottom sediments, mainly as calcite. This process is termed geochemical carbonate compensation. Consequently, CO 2 entering the ocean, including that of anthropogenic origin, can be neutralized permanently by conversion to dissolved bicarbonate ions (HCO 3 −).

The upper oceans are wholly supersaturated with respect to calcite, despite the current acidification. Largely because of the increasing solubility of calcite with pressure, the deeper oceans become undersaturated, whereafter the rate of reaction 1 increases with oceanographic depth. The depth where undersaturation first occurs is the calcite saturation depth (CSD) (2). The preindustrial oceans contained CO 2 acquired from the atmosphere, from marine volcanism, and from oxic organic matter decay. At the same time, calcifying organisms precipitated CaCO 3 shells that settled to the seafloor upon their death. Combined, these processes and reaction 1 lead to decreasing CaCO 3 content in sediments below the CSD. The depth where the deposition rate of CaCO 3 is exactly balanced by reaction 1 is called the calcite compensation depth (CCD) (2), although operationally it is commonly defined as the depth where the sediment CaCO 3 content falls below 10%. The snowline (2) denotes the depth below which no CaCO 3 is found in the sediment; the snowline and CCD coincide at steady state (2).

The geological record contains numerous examples of deep-sea CaCO 3 dissolution events driven by natural acidification, for example, at ∼56 My BP, known as the Paleocene–Eocene Thermal Maximum (3⇓–5). During these events, CaCO 3 disappeared from deep-sea sediments, where it had previously accumulated, burial rates dropped, and the snowline shoaled. The same scenario has been predicted for the Anthropocene oceans (6, 7), but no estimates of increased deep-sea sediment CaCO 3 dissolution have been published. This lack of documentation might be attributed to the restricted penetration of anthropogenic CO 2 to shallow depths, but this premise is contradicted by observed changes in the carbonate chemistry of the deeper oceans (8⇓–10) and as reported here. Both numerical ocean models and the presence in abyssal waters of transient tracers produced almost entirely after the end of the 1940s, that is, chlorofluorocarbons and polychlorinated biphenyls, strongly imply the presence of anthropogenic CO 2 in the deep oceans. Anthropogenic CO 2 in deep and bottom waters is simply very difficult to measure because it is a small signal superimposed on a large natural background concentration. Early changes in CaCO 3 content of sediments caused by anthropogenic acidification are also extremely difficult to detect through changes in either solid mass or dissolution indices.

Given that sediment monitoring is unlikely to yield immediate evidence of anthropogenic calcite dissolution at the seafloor, we employ a different approach and compare the rate of dissolution at the seafloor under preindustrial (∼1800 AD) and modern (2002 AD) benthic conditions. In accord with recent work (11, 12) on the dissolution kinetics of calcite beds, the dissolution rate (r) at any depth between the CSD and the CCD is given by the following (11): r = k ∗ ( [ CO 3 2 − ] eq − [ CO 3 2 − ] SW ) , [2]where [CO 3 2−] eq is the calcite-equilibrium carbonate ion concentration (Fig. 1B), [CO 3 2−] SW is the carbonate ion concentration in bottom waters, and k* is the overall CO 3 2− mass transfer coefficient (11): k ∗ = k s β ( k s + β ) − 1 , [3]where k s is the sediment-side mass transfer coefficient, which characterizes dissolution and transport (diffusion) of the carbonate ion within the sediment, and β is the water-side mass transfer coefficient (13). In this formulation, k* tends toward the value of the smallest rate-limiting mass transfer coefficient, β or k S , without ever reaching it, as predicted by theory and validated by observations. For further explanation with regard to the derivation of Eqs. 2 and 3, please see SI Appendix.

Fig. 1. Chemical and physical parameters for calcite dissolution. (A) Overall mass transfer coefficient k* for CaCO 3 dissolution, (B) current bottom-water saturation concentration [CO 3 2−] eq at in situ temperature, salinity, and pressure, and (C) present-day calcite compensation depth (CCD), regionally averaged as described in Materials and Methods, where the light gray identifies areas where the CCD is deeper than the depth of the seafloor, that is, where calcite can undergo net sediment accumulation. The CCD is not computed in the Arctic and Southern Oceans (south of 60°S).

With any Anthropocene (14) rise in the CCD, r is supplemented by dissolution of previously deposited CaCO 3 between the new and old CCD positions (Materials and Methods). Hence, the difference between the calculated r over the deep seafloor for preindustrial and current conditions reflects the impacts of anthropogenic acidification. To perform this calculation globally, we require the spatial distributions of k*, [CO 3 2−] eq , [CO 3 2−] SW , and the CaCO 3 contents of surface sediments, whose distribution is available from a new database (15, 16) and displayed in SI Appendix, Fig. S2A.

The distribution of k* at the seafloor is shown in Fig. 1A. The water-side mass transfer coefficient, β, appears in Eq. 3 because the seafloor is covered by a water layer through which solute transport occurs via molecular diffusion (13), termed the diffusive boundary layer (DBL). The existence of the DBL has been amply illustrated by previous research (17⇓–19). β is calculated as the ratio of the diffusion coefficient of CO 3 2− ( D CO 3 2 − ), at in situ temperature (T) and salinity (S P ), to the thickness of the DBL (Z DBL ), as β is effectively independent of pressure (Materials and Methods). Whereas D CO 3 2 − does not vary widely, Z DBL depends on the flow velocities at the seafloor (13), but an ocean-wide distribution of Z DBL has never been reported in the literature. Herein, we use a global bottom-current velocity (U) model to derive in situ β values (Materials and Methods). SI Appendix, Fig. S3C provides an ocean-wide distribution of β and shows that this parameter is high on the east side of continents, beneath the Equator, and on the northern fringe of the Southern Ocean, all areas with enhanced bottom currents. Conversely, β is very small over much of the abyssal ocean due to relatively sluggish flow. The magnitude of k s is derived from available experiments and calculated as a function of the CaCO 3 content in surface sediments (Materials and Methods). The distribution of k s at the seafloor is shown in SI Appendix, Fig. S2B. Overall, the rate of deep-sea CaCO 3 dissolution, r, is largely controlled by β, rather than k s (11, 12), except in regions of high bottom currents or where sediments are CaCO 3 -poor, such as the North Pacific or the Southern Ocean (SI Appendix, Fig. S5).

The benthic distribution of [CO 3 2−] eq in Eq. 2 is calculated by dividing the stoichiometric solubility constant of calcite, K* sp at in situ conditions, by the Ca2+ concentration of the oceans. Fig. 1B illustrates the resulting map of [CO 3 2−] eq . [CO 3 2−] eq is commonly near 75 µmol⋅kg−1 on top of oceanic ridges (∼2,500 m) but increases to values as high as ∼140 µmol⋅kg−1 on the abyssal plains (∼6,000 m).

Next, we need to estimate the position of the preindustrial CCD to calculate the amount of dissolution below this depth. As stated earlier, the CCD and snowline are coincident at oceanic steady state. In addition, there is no evidence that the calcite snowline has, as yet, migrated measurably due to anthropogenic dissolution (7, 20). Therefore, we set the preindustrial CCD to the current snowline depth that we estimate from the CaCO 3 contents of sediments for each basin (Materials and Methods). With these local preindustrial CCD values, the flux (F) of CaCO 3 at that depth, and corresponding grid point, can be estimated (2), and the dissolution rate below the CCD, set to the value of F, mapped.

We note that calcifying organisms exhibit various responses to elevated pCO 2 conditions, due to the influence of other climate change-related effects, such as the warming of waters that counteracts acidification (21⇓⇓–24). In the absence of unequivocal evidence of an immediate increase or decrease in the total calcification rate in the pelagic oceans, we assume that the flux of calcitic particles reaching the CCD (F) has remained invariant since the preindustrial era. This is supported by what took place in similar acidification events in the geological record, such as the Paleocene–Eocene Thermal Maximum, where no obvious plankton carbonate productivity reduction is readily detectable (25). Thus, the “current” CCD was computed from the fixed (known) F and the present-day [CO 3 2−] SW , and is represented in Fig. 1C.

Finally, the preindustrial distribution of bottom-water [CO 3 2−] SW (Fig. 2A) is computed using estimated preindustrial, deep-ocean dissolved inorganic carbon (DIC) concentrations (26) and the present-day total alkalinity (TA), which is assumed to have remained invariant so far over the Anthropocene (Materials and Methods). The current bottom-water [CO 3 2−] SW is estimated from the DIC, normalized to the calendar year 2002 (26) (Fig. 2B). The anthropogenic decrease of bottom-water [CO 3 2−] SW is currently limited mostly to the northern Atlantic Ocean and the Southern Ocean, where surface waters enter the deep-ocean reservoir (27). Accordingly, North Atlantic [CO 3 2−] SW has decreased from ∼110 to ∼95 µmol⋅kg−1, whereas the North Pacific [CO 3 2−] SW has remained approximately constant at ∼70 µmol⋅kg−1. The difference between the preindustrial and current bottom-water [CO 3 2−] SW is illustrated in Fig. 2C, which highlights areas of deep acidification in the northwestern North Atlantic and Southern Ocean.

Fig. 2. Ocean acidification and bottom-water [CO 3 2−] SW decrease. (A) Preindustrial bottom-water [CO 3 2−] SW , (B) current bottom-water [CO 3 2−] SW , and (C) difference between preindustrial and current bottom-water [CO 3 2−] SW (Δ[CO 3 2−] SW = current bottom-water [CO 3 2−] SW − preindustrial bottom-water [CO 3 2−] SW ), below 300 m. Uncertainties are indicated by the red outline on the color bar, corresponding to 1 SD of 16.9 µmol⋅kg−1 for Δ[CO 3 2−] SW .

Results With the above information, we calculated the distribution of both preindustrial (Fig. 3A) and current (Fig. 3B) calcite dissolution rates at the seafloor. In both figures, areas with low k* values (Fig. 1A), such as the abyssal plains on each side of the North Atlantic midocean ridge or in the North Pacific, are typified by low calcite dissolution rates, whereas sediments subject to faster currents (larger β values) or rich in CaCO 3 (larger k S values) exhibit greater rates. The difference between current and preindustrial calcite dissolution rates, as plotted in Fig. 3C, corresponds to the rate of anthropogenic calcite dissolution at the seafloor. The anthropogenic dissolution is most pronounced and widespread in the northwestern Atlantic Ocean. There are also areas of significant dissolution in the southern Atlantic and Indian Oceans, resulting from the presence of tangible anthropogenic CO 2 in these deep waters (9, 26), and a consequent rise in the CSD. These prominent dissolution areas reflect the loci of anthropogenic CO 2 introduction into the deep oceans (North Atlantic and the Antarctic margins) and/or the presence of faster bottom currents (larger β values). In the “hot” area of the northwest Atlantic, the anthropogenic component now accounts for 40–100% of the total dissolution rate. Fig. 3. Geographic calcite dissolution rate distribution and anthropogenic impact. (A) Preindustrial and (B) current calcite dissolution rate (r) at the SWI, and (C) the difference between preindustrial and current calcite dissolution rate below 300 m (Δr = current r − preindustrial r), that is, the anthropogenic CO 2 -driven calcite dissolution rate. Hence, this excludes possible calcite dissolution on the continental shelves, particularly in the North Pacific, where the CCD shoals to less than 300 m. Uncertainties are indicated by the red outline on the color bars, corresponding to the SD, equal to 0.05 mol⋅m⋅y−1 for r and to 0.07 mol⋅m⋅y−1 for Δr. The seafloor area between the current and the preindustrial CCD contains the calcitic sediments subjected to the greatest increase in their dissolution rate. In this fringe of high-end anthropogenic dissolution, the CaCO 3 budget at the sediment–water interface (SWI) switches from gain (net accumulation) to deficit (net depletion) when the rising CCD becomes shallower than the seafloor, causing the dissolution of falling calcitic particles before their burial in the sediment. Fig. 3C also suggests the existence of numerous small “hot spots” of dissolution in the Indian and Pacific Oceans, corresponding to topographic highs or islands, for example, Guam and Pitcairn Islands, where the CCD has risen above the seafloor. Because of anthropogenic CO 2 uptake by the oceans, both the saturation (CSD) and compensation (CCD) depths should shoal, leading to dissolution of previously deposited carbonate minerals (28), also known as chemical burndown. Fig. 4 presents our estimates of the upward displacements of the CSD and CCD to date for the Atlantic, plotted with the regionally averaged surface-sediment calcite content water-depth profiles (see SI Appendix, Table S1 for other basins). These results constitute strong signs of anthropogenic dissolution just below the current (risen) CSD, illustrating dissolution of previously deposited calcite, in particular in the North Atlantic where shoaling of the CSD and CCD is greatest. Fig. 4. Atlantic Ocean sediment calcite content profiles and calcite marker horizons. Calcite fraction in dry sediments (X calcite ), ±1 SD, for the basins defined by the maps below the plots. The number of measurements (n) comprised within each basin is also reported. The preindustrial calcite compensation depth (CCD) and saturation depth (CSD) are represented with dashed horizontal lines (on the Left of each profile), whereas the solid horizontal lines correspond to current values (on the Right of each profile). Both the CCD and CSD are reported along with their uncertainties.

Discussion On average, we find that F, the flux of CaCO 3 that must reach the CCD to explain the observed CaCO 3 contents in surficial sediments (Materials and Methods) is 0.17 ± 0.07 mol⋅m−2⋅y−1. This value is equivalent to CaCO 3 fluxes to the deep ocean derived from previous studies based on excess total alkalinity (TA*) budgets (29, 30). Integrating over the surface of the seafloor, we find a global calcite downward flux at the CCD of 54 ± 17 × 1012 mol⋅y−1 (i.e., 0.7 ± 0.2 Gt C⋅y−1), and a current global seafloor CaCO 3 dissolution flux of 32 ± 12 × 1012 mol⋅y−1 of CaCO 3 (i.e., 0.4 ± 0.1 Gt C⋅y−1). Both of these values are similar to and more precise than the observation-based estimates presented in ref. 29. Our global dissolution estimate represents only 22 ± 8% of the total estimated CaCO 3 dissolution in the marine environment each year, that is, 144 × 1012 mol⋅y−1, or 1.7 Gt C⋅y−1 (31). The remainder of the dissolution occurs primarily in the water column and, possibly, after burial through metabolic dissolution or later diagenesis at shallow to mid depths. As stated and justified earlier, we assumed that the calcite rain rate has remained invariant since preindustrial times. Although a decrease in the precipitation rate of calcite might be expected under more acidic conditions (32), thus reducing the calcite export from the surface ocean and the rain rate to deep-sea sediments, no such trend is clearly observed today (33). Testing the sensitivity of the computed sediment dissolution rates (r) to a reduction in the calcite downward flux (F), we find that when F is decreased by 10%, the global dissolution rate increases by 2.5%. Likewise, when F is decreased by 20%, the global dissolution rate increases by 3.1%. Thus, counterintuitively, if less calcite is currently delivered to the seafloor than during the preindustrial era, more calcite is being dissolved at the SWI because a weaker F causes the CCD to respond faster to changes in [CO 3 2−] SW , as explained in ref. 2, and in the case of the present acidification, to rise faster. Sediments that were above the preindustrial CCD and are now below the current CCD are subject to the greatest dissolution rate increase, as can be seen when comparing Eqs. 12–14 in Materials and Methods. In other words, if F has remained constant, ∼2% of the seafloor is currently comprised between the preindustrial and current CCD, but this fraction increases to 9.3% when F is reduced by 20% relative to its preindustrial value. Our results provide a tangible indication that significant anthropogenic calcite dissolution is currently occurring at the seafloor. This dissolution will eventually change the CaCO 3 accumulation patterns and rates in the oceans, while mitigating runaway ocean acidification. Any future model of oceanic acidification needs to reproduce our observations to assure its validity. The consequences of this anthropogenic dissolution to the ecology of benthic calcifiers has not been determined with certainty, but could be as substantial as it is for pelagic calcifiers (34). Chemical burndown of previously deposited carbonate-rich sediments has already begun and will intensify and spread over vast areas of the seafloor during the next decades and centuries, thus altering the geological record of the deep sea. The deep-sea benthic environment, which covers ∼60% of our planet, has indeed entered the Anthropocene.

Acknowledgments We thank all who contributed to the creation of GLODAPv2. We also thank John H. Trowbridge and the two anonymous reviewers for valuable discussions and criticism. This research was funded by Natural Sciences and Engineering Research Council of Canada Discovery Grants (to A.M. and B.P.B.). D.S.T. and B.K.A.’s contributions to this study were funded by the US National Science Foundation (NSF) Grant OCE-0960820, and R.M.K. was funded by NSF Grant PLR-1425989. O.S. also acknowledges the Department of Earth and Planetary Sciences at McGill University for financial support during his residency in the graduate program.

Footnotes Author contributions: O.S., B.P.B., and A.M. designed research; O.S. performed research; O.S., B.P.B., A.M., C.J., D.S.T., B.K.A., and R.M.K. contributed new reagents/analytic tools; O.S., B.P.B., C.J., D.S.T., B.K.A., and R.M.K. analyzed data; and O.S., B.P.B., A.M., C.J., D.S.T., B.K.A., and R.M.K. wrote the paper.

The authors declare no conflict of interest.

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