The contribution of water radiolysis to marine sedimentary life
Radiation experiments
We experimentally quantified radiolytic hydrogen (H2) production in (i) pure water, (ii) seawater, and (iii) seawater-saturated sediment. We irradiated these materials with α- or γ-radiation for fixed time intervals and then determined the concentrations of H2 produced. Sediment samples were slurried with natural seawater to achieve a slurry porosity (φ) of ~0.83, which is the average porosity of abyssal clay in the South Pacific Gyre34. The seawater source is described below. To avoid microbiological uptake of radiolytic H2 during the course of the experiment, seawater and marine sediment slurries were pre-treated with HgCl2 (0.05% solution) or NaN3 (0.1% wt/vol). To ensure that addition of these chemicals did not impact radiolytic H2 yields, irradiation experiments with pure water plus HgCl2 or NaN3 were also conducted. HgCl2 or NaN3 addition had no statistically significant impact on H2 yields5,6,10.
Experimental samples were irradiated in 250 mL borosilicate vials. A solid-angle 137Cs source (beam energy of 0.67 MeV) was used for the γ-irradiation experiments at the Rhode Island Nuclear Science Center (RINSC). The calculated dose rate for sediment slurries was 2.19E−02 Gy h−1 accounting for the (i) source activity, (ii) distance between the source and the samples, (iii) sample vial geometry, and (iv) attenuation coefficient of γ-radiation through air, borosilicate, and sediment slurry. 210Po (5.3 MeV decay−1)-plated silver strips with total activities of 250 μCi were used for the α-irradiation experiments. For α-irradiation of each sediment slurry, a 210Po-plated strip was placed inside the borosilicate vial and immersed in the slurry. Calculated total absorbed doses were 4 Gy and 3 kGy for γ-irradiation and α-irradiation experiments, respectively.
The settling time of sediment grains in the slurries (1 week) was long compared to the time span of each experiment (tens of minutes to an hour for α-experiments, hours to days for γ-experiments). Therefore, we assumed that the suspension was homogenous during the course of each experiment.
H2 concentrations were measured by quantitative headspace analysis via gas chromatography. For headspace analysis, 30 mL of N2 was first injected into the sample vial. To avoid over-pressurization of the sample during injection, an equivalent amount of water was allowed to escape through a separate needle. The vials were then vigorously shaken for 5 min to concentrate the H2 into the headspace. Finally, a 500-μL-headspace subsample was injected into a reduced gas analyzer (Peak Performer 1, PP1). The reduced gas analyzer was calibrated using a 1077 ppmv H2 primary standard (Scott-Marrin, Inc.). A gas mixer was used to dilute the H2 standard with N2 gas to obtain various H2 concentrations and produce a five-point linear calibration curve (0.7, 2, 5, 20, and 45 ppm). H2 concentrations of procedural blanks consisting of sample vials filled with non-irradiated deionized 18-MΩ water were also determined. The concentration detection limit obtained using this protocol was 0.8–1 nM H2. Relative error was less than 5%. Radiation experiments were performed at a minimum in triplicate.
Sample selection and experimental radiolytic H2 yields, G(H2)
Millipore Milli-Q system water was used for our pure-water experiments. For seawater experiments, we used bottom water collected in the Hudson Canyon (water depth, 2136 m) by RV Endeavor expedition EN534. Salinity of North Atlantic bottom water in the vicinity of the Hudson Canyon (34.96 g kg−1) is similar to that of mean open-ocean bottom water (34.70 g kg−1)44,45.
The 20 sediment samples used for the experiments were collected by scientific coring expeditions in three ocean basins (expedition KN223 to the North Atlantic46, expedition KN195-3 to the Equatorial and North Pacific47, International Ocean Discovery Program (IODP) Expedition 329 to the South Pacific Gyre34, MONA expedition to the Guaymas Basin48, expedition EN32 to the Gulf of Mexico49, and expedition EN20 to the Venezuela Basin50). To capture the dominant sediment types present in the global ocean, we selected samples typical of five common sediment types [abyssal clay (11 samples), nannofossil-bearing clay or calcareous marl (2 samples), clay-bearing diatom ooze (3 samples), calcareous ooze (2 samples), and lithogenous sediment (2 samples)]. The locations, lithological descriptions, and mineral compositions of the samples are given in Supplementary Tables 1, 2, 3, and Supplementary Fig. 1. Additional chemical and physical descriptions of the sediment samples used in the radiation experiments can be found in the expedition reports for the expeditions on which the samples were collected34,46.
Energy-normalized radiolytic H2 yields are commonly expressed as G(H2)-values (molecules H2 per 100 eV absorbed)1. As shown in Supplementary Fig. 2, for all irradiated samples (pure water, seawater, and marine sediment slurries), H2 production increased linearly with absorbed α- and γ-ray-dose. We calculated G(H2)-values for each sample and radiation type (α or γ) as the slope of the least-square regression line of radiolytic H2 concentration versus absorbed dose (Supplementary Fig. 2). The error on the yields is less than 10% for each sample. G(H2)-values for each sample and radiation type (α or γ) are reported in Supplementary Table 3.
Although radiolytic OH• is known to react with dissolved organic matter51, total organic content does not appear to significantly impact radiolytic H2 production, since the most organic-rich sediment (e.g., Guaymas Basin and Gulf of Mexico sediment) did not yield particularly high H2 (Supplementary Table 3).
Calculated radiolytic production rates of H2 and oxidants in the cored sediment of individual sites
We calculated radiolytic H2 production rates (PH2, in molecules H2 cm−3 yr−1) for the cored sediment column at nine sites with oxic subseafloor sediment in the North Pacific, South Pacific, and North Atlantic; and seven sites with anoxic subseafloor sediment in the Bering Sea, South Pacific, Equatorial Pacific, and Peru Margin (see Supplementary Fig. 3 for site locations). For these calculations, we used the following equation from Blair et al.2:
$$P_{{mathrm{H}}_2} = {sum} A _{{mathrm{m}},i}rho left( {1 – {{upvarphi}} } right)E_i{mathrm{G}}({mathrm{H}}_2)_i$$
(1)
where i is alpha, beta, or gamma radiation; Am is radioactivity per mass solid; φ is porosity; ρ is density solid; (E_i) is decay energy; and ({mathrm{G}}({mathrm{H}}_2)_i) is radiolytic yield.
We calculated radiolytic oxidant production rates for these sediment columns from the H2 production rates. Because H2 production and oxidant production are stochiometrically balanced in water radiolysis [2H2O → H2 + H2O2], the calculated radiolytic H2 production rates (in electron equivalents) are equal to radiolytic oxidant production rates (in electron equivalents).
The in situ γ- and α-radiation dosages in marine sediment are, respectively, 13 and 15 orders of magnitude lower than the dosage used in our experiments. Because the measured G(H2) for pure water in our γ-irradiation experiment (dose rate = 2.19E-02 Gy h−1) is statistically indistinguishable from previously published G(H2) values at much higher dose rates (ca. 1.00E+3 Gy h−1)5, we infer that the γ-irradiation G(H2) value is constant with dose rate over five orders of magnitude. Similarly, our experimental pure water H2 yields following α-particle irradiation from a 210Po-source (dose rate of 2.55E+03 Gy h−1) are indistinguishable from the yield obtained by Crumière et al.6 [G(H2) = 1.30 ± 0.13] for air-saturated deionized water exposed to a cyclotron-generated He2+ particle beam at higher dose rate (dose rate 1.62E+05 Gy h−1). The close similarity in H2 yields obtained in both experiments implies that (i) radiolytic H2 yield from α-particle irradiation is identical to that from cyclotron-generated He2+ particle irradiation, and (ii) this yield is constant over a two-orders-of-magnitude range dose rate. Therefore, we use our experimentally determined α- and γ-irradiation G(H2) values for the low radiation dose rate found in the subseafloor. Because the G(H2) of β irradiation has not been experimentally determined for water-saturated materials, we assume that the G(H2) of β-radiation matches the G(H2) of γ-radiation for the same sediment types. In pure water, their G(H2) values differ by only 17%1. Because β radiation, on average, contributes only 11% of the total radiolytic H2 production from the U, Th series and K decay in marine sediment, these estimates of total H2 production differ by only 2–5% relative to estimates where the G(H2) of β radiation is assumed equal to that for pure water or for α radiation of the same sediment types.
To calculate H2 production rates for the entire sediment column at seven South Pacific sites and two North Atlantic sites, we measured downcore sediment profiles of U, Th, and K (i) 187 sediment samples from IODP Expedition 329 Sites U1365, U1366, U1367, U1368, U1369, U1370, and U137134,52, and (ii) 40 samples from KN223 expedition Sites 11 and 12 (ref. 46). Total U and Th (ppm) and K2O (wt%) for these sites are reported in the EarthChem SedDB data repository. We measured U, Th, and K abundances using standard atomic emission and mass spectrometry techniques (i.e. ICP-ES and ICP-MS) in the Analytical Geochemistry Facilities at Boston University. Sample preparation, analytical protocol, and data are reported in Dunlea et al.52. The precision for each element is ~2% of the measured value, based on three separate digestions of a homogenized in-house standard of deep-sea sediment.
To calculate H2 production rates for the sediment columns at North Pacific coring Sites EQP10 and EQP11 (ref. 47), we used radioactive element content data from Kyte et al.53, who measured chemical concentrations at high resolution in bulk sediment in core LL44-GPC3. Because Site EQP11 was cored at the same location as LL44-GPC3 (ref. 53) and the sediment retrieved at all three sites is homogeneous abyssal clay, we assume the radioactive element abundances measured in core LL44-GPC3 to be representative of Sites EQP10 and EQP11 (ref. 47). Calculated radiolytic H2 production rates for South Pacific sites are listed in Supplementary Table 4 and for North Atlantic and North Pacific sites in Supplementary Table 5.
For Bering Sea Sites U1343 and U1345 (ref. 54), sedimentary U, Th, and K content measurements are unavailable. Since sediment recovered at these two sites is primarily siliciclastic with a varying amount of diatom-rich clay, we use U, Th, and K concentration values reported for upper continental crust by Li and Schoonmaker for these Bering Sea sites55. Finally, we calculate downhole radiolytic H2 production rates for ODP Leg 201 Sites 1225, 1226, 1227, and 1230 (ref. 35). Sediment compositions for these sites include nannofossil-rich calcareous ooze (Site 1225), alternation of nannofossil (calcareous) ooze and diatom ooze (Site 1226), and siliciclastic with diatom-rich clay intervals (Sites 1227 and 1230). Because sedimentary U, Th, and K measurements are not available for Leg 201 sites, we used average U, Th, and K concentration values measured in North Atlantic46 and South Pacific Sites34,52 with corresponding lithologies.
We use isotopic abundance values reported in Erlank et al.56 to calculate the abundance of 238U, 235U, 232Th, and 40K from the measured ICP-MS values of total U, Th, and K concentration. We then converted radionuclide concentrations to activities using Avogadro’s number and each isotope’s decay constant2. We refer to Blair et al. for a detailed explanation of activities and radiolytic yield calculations2.
Calculation of global radiolytic H2 and oxidant production rates in marine sediment
We calculated global radiolytic H2 production in ocean sediment by applying Eq. (1) (ref. 2) globally. As with the rates at individual sites, we calculated global radiolytic oxidant production (in electron equivalents) from global H2 production and the stochiometry of water radiolysis [2H2O → H2 + H2O2].
Our global radiolytic H2 production calculation spatially integrates calculations of sedimentary porewater radiolysis rates that are based on (i) our experimentally constrained radiolytic H2 yields for the principal marine sediment types, (ii) measured radioactive element content of sediment cores in three ocean basins (North Atlantic46, North Pacific53, and South Pacific34,52), and (iii) global distributions of sediment lithology57, sediment porosity58, and sediment column length59,60.
To generate the global map of radiolytic H2 production, we created global maps of seafloor U, Th, and K concentrations, density, G(H2)-α values, and G(H2)-γ-and-β by assigning each grid cell in our compiled seafloor lithology map (Supplementary Fig. 4) its lithology-specific set of input variables (Supplementary Table 6). Because our model assumes that lithology is constant with depth, U, Th, and K content, grain density, and G(H2)-values are constant with depth.
The G(H2)-values (α, β, and γ radiation), radioactive element content (sedimentary U, Th, and K concentration), density, porosity, and sediment thickness are determined as follows.
Radiolytic yield [G(H2)] for α,β-&-γ radiation
Radiolytic yields for the main seafloor lithologies are obtained by averaging experimentally derived yields for the respective lithologies (Supplementary Table 6). We assume that G(H2)-β values equal G(H2)-γ values.
Sediment lithology
For these calculations of radiolytic chemical production, we generally used seafloor lithologies and assumed that sediment type is constant with sediment depth. For seafloor lithology, the geographic database of global bottom sediment types57 was compiled into five lithologic categories: abyssal clay, calcareous ooze, siliceous ooze, calcareous marl, and lithogenous (Supplementary Fig. 4). Some areas of the seafloor are not described in the database57. These include (i) high-latitude regions (as the seafloor lithology database extends from 70°N to 50°S)57 and (ii) some discrete areas located along continental margins (e.g., Mediterranean Sea, Timor Sea, South China Sea, Supplementary Fig. 4). We used other data sources to identify seafloor lithologies for these regions. We added an opal belt (siliceous ooze) in the Southern Ocean between 57°S and 66°S61,62. The geographic extent of this opal belt was based on DeMaster62 and Dutkiewicz et al.61. We defined the areas of the seafloor from 50°S to 57°S, from 66°S to 90°S, and in the Arctic Ocean as mostly composed of lithogenous material, based on (i) drillsite lithologies in the Southern Ocean [ODP: Site 695 (ref. 63), Site 694 (ref. 63), Site 1165 (ref. 64), Site 739 (ref. 65)], the Bering Sea and Arctic Ocean [International Ocean Discovery Program (IODP): Sites U1343 and U1345 (ref. 54), Site M0002 (ref. 66), ODP: Site 910 (ref. 67), Site 645 (ref. 68)] and between 50°S and 57°S [Deep Sea Drilling Project (DSDP): Site 326 (ref. 69), Ocean Drilling Program (ODP): Site 1138 (ref. 70), Site 1121 (ref. 71)], and Dutkiewicz et al.61.
In the North and South Atlantic, sediment type can be very different at depth than at the seafloor. For these regions, we departed from our assumption that sediment lithology is the same at depth as at the seafloor. Subseafloor lithologies at ODP Sites [1063 (ref. 72), 951 (ref. 73), 925 (ref. 74), and 662 (ref. 75)] and IODP Sites [U1403 (ref. 76) and U1312 (ref. 77)] indicate that sediment in the Atlantic Ocean basin is generally 30–90% biogenic carbonate content and detrital clay78, even where the seafloor lithology is abyssal clay57. Therefore, regions in the Atlantic Ocean described as abyssal clay in the seafloor lithology database57 were characterized as calcareous marl for our calculations (Supplementary Fig. 4). Because abyssal clay catalyzes radiolytic H2 production at a higher rate than calcareous marl, this characterization may underestimate production of radiolytic H2 and radiolytic oxidants in these Atlantic regions.
Radioactive element content
For four of the five lithologic types in our global maps (abyssal clay, siliceous ooze, calcareous ooze, and calcareous marl), we average U, Th, and K concentrations from sites in the North Atlantic46, North Pacific53, and South Pacific34,52. The average U, Th, and K concentration values are consistent with data reported in Li and Schoonmaker55 for the characteristic U, Th, and K content found in abyssal clay and calcareous ooze. For lithogenous sediment, we use U, Th, and K concentration values reported for upper continental crust by Li and Schoonmaker55. Lithology-specific radioactive element values are given in Supplementary Table 6 and used to calculate Am,i in Eq. (1).
Density
Characteristic density values for calcite, quartz, terrigenous clay, and opal-rich sediment were extracted from the Proceedings of the Integrated Ocean Drilling Program Volume 320/321 and are assigned to calcareous ooze, lithogenous sediment, abyssal clay, and siliceous ooze, respectively79.
Global porosity
For global porosity, we use a seafloor porosity data set by Martin et al.58 and accounted for sediment compaction with depth by using separate sediment compaction length scales for continental-shelf (0–200 m water depth; c0 = 0.5 × 10−3), continental-margin (200–2500 m; c0 = 1.7 × 10−3), and abyssal sediment ( >3500 m; c0 = 0.85 × 10−3)80,81. Once the porosity was 0.1%, the depth integration was halted.
Global sediment thickness
We calculated global depth-integrated radiolytic H2 production by summing the seafloor production rates over sediment depth in one-meter intervals (Fig. 3 in main text). Sediment thickness is from Whittaker et al., supplemented with Laske and Masters where needed82,83.
Ocean depth
For porosity calculations, water depths were determined using the General Bathymetric Chart of the Oceans84, resampled to a 5-arc minute grid, i.e. the resolution of the Naval Oceanographic Office’s Bottom Sediment Type (BTS) database “Enhanced dataset”57.
Dissolved H2 concentration profiles
H2 concentrations from South Pacific Sites U1365, U1369, U1370, and U1371, and the measurement protocol, are described in ref. 1. H2 concentrations from North Atlantic KN223 Sites 11, 12, and 15, and North Pacific Site EQP11 were determined using the same protocol and are posted on SedDB (see “Data availability”). The detection limit for H2 ranged between 1 and 5 nM H2, depending on site, and is displayed as gray vertical lines in Fig. 2 of the main text. H2 concentrations for Equatorial Pacific Site 1225 and Peru Trench Site 1230 were measured by the “headspace equilibration technique”, which measures steady-state H2 levels reached following laboratory incubation of the sediment samples85,86.
For comparison to these measured H2 concentrations, we use diffusion-reaction calculations to quantify what in situ H2 concentrations would be in the absence of H2-consuming reactions. The results of these calculations are represented as solid circles (•) in Fig. 2 of the main text. Temporal changes in H2 concentration due to diffusive processes and radiolytic H2 production in situ are expressed by Eq. (2):
$$frac{{partial {mathrm{H}}_2(x,t)}}{{partial t}} = frac{D}{{varphi F}}frac{{partial ^2{mathrm{H}}_2(x,t)}}{{partial x^2}} + P(x)$$
(2)
with
D: the diffusion coefficient of H2(aq) at in situ temperature
(varphi): porosity
F: formation factor
x: depth
Z: sediment column thickness
({mathrm{H}}_2): hydrogen concentration
P: radiolytic H2 production rate
t : time.
With constant radiolytic H2 production, P(x) = P with depth,
and at steady-state,
$$frac{{partial ^2{mathrm{H}}_2(x)}}{{partial x^2}} = – frac{{Pvarphi F}}{D}.$$
(3)
We integrate Eq. (3) over the length x twice,
$${mathrm{H}}_2(x) = – frac{1}{2}frac{{Pvarphi F}}{D}x^2 + Ax + B$$
(4)
where A and B in Eq. (4) are constants of integration. We use two boundary conditions to derive the value of these constants.
Boundary condition 1: concentration of H2 at the sediment-water interface, x = 0, is zero due to diffusive loss to the overlying water column.
Boundary condition 2: concentration of H2 at the basement–sediment-water interface, x = Z, is zero due to diffusive loss to the underlying basement.
With these boundary conditions, (A = frac{1}{2}frac{{Pvarphi F}}{D}Z) and B = 0
and
$${mathrm{H}}_2(x) = frac{1}{2}frac{{Pvarphi F}}{D}(xZ – x^2).$$
(5)
In cases where we expect radiolytic H2 production rates to significantly vary with depth due to changes in lithology, we adapted the boundary conditions and applied a two-layer diffusion model to account for this variation.
Calculation of Gibbs Energies for the Knallgas reaction
For H2 concentrations above the detection limits at South Pacific IODP Expedition 329 sites (Supplementary Fig. 5)34, we quantified in situ Gibbs energies (ΔGr) of the Knallgas reaction (H2 + ½O2 → H2O). In situ ΔGr values depend on pressure (P), temperature (T), ionic strength, and chemical concentrations, all of which are explicitly accounted for in our calculations:
$$Delta G_{mathrm{r}} = Delta G^circ _{mathrm{r}}left( {T,P} right) + 2.3,RT,{mathrm{log}}_{10}Q$$
(6)
where:
ΔGr: in situ Gibbs energy of reaction (kJ mol H2−1)
ΔG°r(T,P): Gibbs energy of reaction under in situ T and P conditions (kJ mol H2−1)
R: gas constant (8.314 kJ−1 mol K−1)
Q: activity quotient of compounds involved in the reaction.
We use the measured composition of the sedimentary pore fluid to determine values of Q.
For a more complete overview of in situ Gibbs energy-of-reaction calculations in subseafloor sediment, see Wang et al.87.
Calculation of organic oxidation rates (net rates of O2 reduction and DIC production)
We calculated the vertical distribution of net O2 reduction rates at nine sites where the sediment is oxic from seafloor to basement and the vertical distribution of DIC production rates at seven sites where the subseafloor sediment is anoxic (see Supplementary Fig. 3 for site locations). Dissolved O2 concentrations are from Røy et al.47 and D’Hondt et al.88. DIC concentrations are from ODP Leg 201 (ref. 35), and the Proceedings of the IODP Expedition 323 (Sites U12343, U1345)54 and IODP Expedition 329 (Site U1371 (ref. 34)).
The net rates are calculated using the MatLab program and numerical procedures of Wang et al.89, modified by using an Akima spline, rather than a 5-point running mean, to generate a best-fit line to the chemical concentration data. Details of the calculation protocol for O2 production rates and DIC production rates are respectively described in the supplementary information of D’Hondt et al.88 and in Walsh et al.90. The DIC reaction rates and their first standard deviations calculated for the seven sites are given in Supplementary Table 7.
To facilitate comparisons of radiolytic chemical rates to net DIC production rates, rates are converted to electron equivalents (2 electrons per H2, 4 electrons per O2, 4 electrons per organic C oxidized).
Estimation of sediment ages
We estimated sediment ages for Sites U1343 and U1343 using the sediment-age model of Takahashi et al.54, which is based on biostratigraphic and magnetostratigraphic data. Because detailed chronostratigraphic data are not available for the remaining sites (Equatorial Pacific sites (1225 and 1226), Peru Trench Site 1230 and Peru Basin Site 1231, South Pacific sites U1365, U1366, U1367, U1369, U1370, and U1371, North Pacific sites EQP9 and EQP10, and North Atlantic sites KN223-11 and KN223-12), we used the mean sediment accumulation rate for each of these sites (Supplementary Fig. 3) to convert its sediment depth (in meters below seafloor) to sediment age (in millions of years, Ma). Mean sediment accumulation rate was calculated by dividing sediment thickness by basement age91 (Supplementary Table 8). For Sites 1225, 1226, 1230, 1231, U1365, U1366, U1367, U1369, U1370, and U1371, sediment thickness was determined by drilling to basement34,35. For Sites EQP9, EQP10, KN223-11, and KN223-12, sediment thicknesses were determined from acoustic basement reflection data. More
