CaCO3 dissolution in sediments of the Ceara Rise, western equatorial Atlantic

Geochmica

et Cosmochmica

Pergamon

Ac~a, Vol. 60. No. 7. pp. 233-263, 1996 CopyrIght 6; 1996 Elsevier Science Lid Printed in Ihc USA. All lights I-eset-ved 0016.7037/96 $15.00 + .OO

0016-7037( 95 )00383-5

CaC03 dissolution in sediments of the Ceara Rise, western equatorial Atlantic W. R. MARTIN and F. L. SAYLES Department of Marine Chemistry and Geochemistry. Woods Hole Oceanographic Institution. Woods Hole. MA 02543. USA (Received May 2, 1995; accepted in revised

form October 16. 1995)

have used porewater sampling by in situ techniques, including whole-core squeezing, as well as by shipboard sectioning and whole-core squeezing to estimate the rates of sedimentary organic matter oxidation and CaC03 dissolution at seven sites on the Ceara Rise in the western equatorial Atlantic Ocean. Porewater NO, profiles at all sites show a pattern indicative of active organic matter oxidation in the upper 15-20 cm of the sediments and in a buried, organic-rich layer. The organic C oxidation rate generally decreases with increasing water depth, from a value of 22 pmol/cm’/y at the shallowest site (3279 m) to 14 bmol/cm’/y at the deepest site (4675 m). Over this depth range, the bottomwaters vary from moderately supersaturated with respect to calcite to strongly undersaturated. High-resolution alkalinity profiles, measured in porewaters collected by in situ whole-core squeezing, yield estimated Ca” fluxes of 11 pmollcm2/y at a site located at the depth of the calcite saturation horizon, and 7.6 /Imollcm’/y at a moderately undersaturated site. Ca’+ fluxes calculated from profiles in porewaters collected by relatively coarse-resolution in situ sampling methods clearly indicate that there is CaCO? dissolution above the calcite saturation horizon. The dissolution of aragonite may contribute to the dissolution flux at the shallowest site. These Ca” fluxes, as well as fluxes estimated from a model of sedimentary organic matter oxidation and calcite dissolution, indicate that 36-66% of the CaCO? rain to the seafloor dissolves at sites at and above the calcite saturation horizon, while 52-75% of the rain dissolves at sites below this depth. When these results are incorporated into the oceanic CaCO, budget of Milliman ( 1993), they indicate that 35% of CaC03 production is preserved in the deep sea; they suggest a CaCO, accumulation rate that is 27% lower than that estimated by Milliman ( 1993). Our C,,,, oxidation/CaCO? dissolution model indicates that a large fraction of the CaCO? dissolution that is occurring on the Ceara Rise is attributable to the neutralization of metabolic acids produced during organic matter oxidation. The efficiency with which organic matter oxidation dissolves CaCO, (that is, the ratio, CaCO, dissolution attributable to organic matter oxidation:organic matter oxidation rate) generally increases as degree of undersaturation of bottomwaters increases. However, there are deviations from the general trend that can be attributed to site-to-site variations in the kinetics of organic matter oxidation and calcite dissolution. This result indicates that the dissolution of CaCOJ as a result of organic matter oxidation in the deep sea may mask the effects of variations in surface water CaCO, nroductivitv and bottomwater chemistry on the accumulation rate of CaCOJ in deep-sea sediments. .’ I ’

Abstract-We

1. INTRODUCTION

understanding of CaCO, dissolution in sediments is a prerequisite to correctly interpreting changes in sedimentary CaC07 concentrations and accumulation rates over time. In particular, it is essential to differentiate between the effects of changes in surface water CaCO, productivity, changes in the preservation/dissolution ratio that are driven by variations in the composition of the particle rain to the seafloor (the C,,,,/ CaCO? and SiOJCaCO, ratios), and changes in ocean circulation that influence the dissolution/preservation ratio, such as the corrosiveness of bottomwaters to CaC03 and bottomwater dissolved oxygen distributions (e.g., Arrhenius, 1988; Boyle, 1988; Curry and Lohmann, 1985; Farrel and Prell, 1989; Lyle et al., 1988; Pedersen et al., 1988). Many attempts at estimating the productivity of the oceans in the past have been based on determinations of accumulation rates of CaCO, as a function of age in sediment cores. The central assumption upon which these studies are based is that CaCOj dissolution in sediments lying above the saturation horizon or lysocline is negligible. In the absence of postdepositional dissolution, variations in accumulation rates of CaCO? are attributed solely to variations in the productivity of the overlying surface water (Berger, 1982b; Curry and Lohmann, 1985. 1990; Farrel and Prell, 1989; Francois et al.,

The response of the oceanic carbon cycle to changing inputs and water circulation is characterized by feedback loops, operating on timescales of a few thousand years, that influence atmospheric CO2 levels (Broecker and Peng, 1982). The uptake of C to form CaCO, in the surface ocean and its release upon dissolution of CaCOJ into the deep waters of the oceans are central components of the oceanic C cycle. While changes in the locus of CaCO? accumulation from deep to shallow associated with high sea level stands modify the oceanic carbonate cycle (Berger, 1982a; Milliman, 1993). changes in preservation/dissolution of CaCO, in the deep sea are required to explain the major shifts in atmospheric pC0, of glacial-interglacial transitions and the sedimentary record of these changes (Archer and Maier-Raimer, 1994; Berger, 1982b; Boyle, 1988; Broecker and Peng, 1987; Keir and Berger, 1985 ) . Since most CaCO? dissolution takes place at the surface of and within deep-sea sediments, rather than in the water column, identifying and quantifying the factors influencing sedimentary CaCOz dissolution are crucial to understanding the marine CaCO, cycle, its role in the marine carbon cycle, and controls on atmospheric CO? levels. In addition, an 243

244

W. R. Martin and F. L. Sayles

1990; Lyle et al., 1988). If, however, dissolution above the lysocline is significant, as a number of studies have indicated (Archer et al., 1989; Emerson and Bender, 1981; Hales et al., 1994; Sayles, 1981), then the estimation of productivity from CaCOj accumulation rates above the saturation horizon cannot be correct and variations in accumulation rates over time may reflect changes in preservation rather than in surface water productivity. Assessing the applicability of this approach to determining paleo-productivity requires a knowledge of dissolution above the saturation horizon and the factors that control the ratio of preservation to dissolution in this depth range of the oceans. Significant dissolution above the saturation horizon can also have an important impact on CaCO, cycling and mass balances in the marine environment. Milliman (1993) has constructed a mass balance for CaCO, in the oceans. In estimating accumulation in the deep sea, he assumed dissolution of 20% of the CaCO? rain above the lysocline. If dissolution rates such as we have calculated for the Ceara Rise (see below) and have been reported by others for locations above the saturation horizon (Hales et al., 1994; Sayles, 1981; Sayles and Curry, 1988) are representative, then dissolution of -40% of the rain would be a more appropriate figure. Such an increase in dissolution significantly changes CaCO, accumulation in the deep sea, reducing the Milliman ( 1993) estimate by -27%. As noted above, virtually all scenarios put forward to explain glacial-interglacial variations in atmospheric pC0, have significant impact on the CO:- ion concentration and hence, the cycling of CaCO? in the deep sea. The mechanism(s) providing this linkage have not, however, been identified (Sundquist, 1993). The importance of the relative proportions of organic carbon and CaCO, in particulate matter reaching the sediments of the seafloor to determining the fraction of CaCO, surviving to be buried has been recognized for many years (Bemer, 1977; Gieskes, 1970; Parker and Berger, 1971; Schink and Guinasso, 1977; and numerous references cited above). Archer and Maier-Raimer (1994) have postulated that glacial-interglacial changes in the ratio of organic carbon to CaCO, (C,,,/CaCO,) in particulate rain and corresponding changes in CaCO, preservation in the deep sea can account for the large variations in atmospheric pCOZ associated with these transitions. While the ratio C,,/CaC03 determines the maximum potential dissolution driven by CO, from organic matter degradation, the maximum is probably rarely realized. The extent and efficiency with which metabolic CO2 leads to CaCOJ dissolution depends on a number of factors which themselves are likely to vary in different environments. These factors include the saturation state of the overlying water, the depth profile of organic carbon remineralization rate, the C,,/ CaCOj of the sedimenting particulate matter, and the rate constant of CaC03 dissolution (in at least some environments, both calcite and aragonite). At present, the information on these variables and their interrelationship is insufficient to evaluate the role of varying particulate C,,,/CaCO, on CaCO, dissolution/preservation in the deep sea, and ultimately on atmospheric pC02. The study of CaCO, dissolution in the deep sea has been advanced in recent years by the deployment of in situ benthic flux chambers (Berelson et al., 1990; Jahnke et al., 1994) and

in situ microelectrode profilers (Archer et al., 1989; Cai et al., 1995; Hales et al., 1994). Deployments of the different instruments above the calcite saturation horizon have had mixed results: while microelectrode profiles have consistently and convincingly shown the occurrence of above-saturation horizon dissolution (Archer et al., 1989; Hales et al., 1994; Hales and Emerson, 1994), benthic flux chambers have detected dissolution in some, but not all continental margin areas but not in the deep sea (Jahnke, 1994; Jahnke et al., 1994). We report here the results of porewater sampling and modeling studies aimed at quantifying the dissolution of CaC03 in sediments of the Ceara Rise in the western equatorial Atlantic. The sediments of the Rise are exposed to bottomwaters that range from strongly undersaturated to moderately supersaturated with respect to calcite. Our results include the first effort to determine CaCO? dissolution rates using alkalinity profiles obtained by in situ whole-core squeezing, providing high-resolution alkalinity profiles at two sites: one essentially at the calcite saturation horizon and one in moderately undersaturated bottomwater. These profiles are used to infer dissolution fluxes at these sites and are joined with additional porewater data, defining the organic carbon oxidation rate, to evaluate the rate constant for calcite dissolution. We use this rate constant information with additional, coarse-resolution in situ porewater alkalinity, XOZ, and Ca2+ profiles to quantify calcite dissolution at seven sites that cover the range of bottomwater saturation conditions on the Ceara Rise. 2. STUDY SITES In March 1994. we occupied seven sites on the northern part of the Ceara Rise, located at about 5”N, 43”W in the western equatorial

Atlantic Ocean. The rise spans a depth range of about 3000 m to 4700 m. The rise is sediment-covered throughout this depth range, with estimated Holocene sedimentation rates of 2-2.5 cm/ky (Curry and Lohmann, 1990). Calcite is abundant in the upper 10 cm of the sediments at all of our sample sites: sediments are. 60-67% CaCO,

at all but the deepest site, where the CaCO, concentration drops to 35%. We estimate that the calcite saturation horizon occurs at about

4000 m, a depth that approximately correspondsto the regional lysocline (Biscaye et al., 1976). Bottomwaters are well oxygenated. at 260 pmollkg above the apparent NADWlAABW boundary (4000 m), decreasing to 240 ,umol/kg at 4700 m. The locations, depths, and calcite saturation conditions of sample sites are listed in Table 1.

3. SAMPLING METHODS Two wire-line instruments were used to collect sediments and porewaters. A multi-corer (operated by R. Jahnke, SMO), which takes four piston cores after landing on the seafloor, was used to collect sediments for shipboard sampling. One core on each deploy-

ment was fitted with a piston designed for shipboard whole-core squeezing using the methodsdescribedby Martin et al. ( 1991). Other cores were used for shipboard sectioninglcentrifugation,

resistivity

profiling, and solid phase sampling. Recovered cores typically had very clear overlying water and apparently intact interfaces, indicating that disturbance during coring was minimal. The WHIMP (Sayles, 1981) was used for in situ porewater sampling. It is equipped with an in situ whole-core squeezer(Bender et al., 1987; Sayles and Dickinson, 1991) for high resolution sampling near the sediment-water interface and with harpoon-type samplers for collecting porewaters on a coarser depth scale from 2-60 cm below the interface.

Porewaters collected by shipboard SectioningIcentrifugation were analyzed for NO; ; shipboard whole-core squeezing (WCS) was used to measure interfacial profiles of NO; and Oz. Cores sampled on board ship were handled as rapidly as possible to minimize artifacts.

Kinetics of CaCOi dissolution in marine sediment

Station

Lat/Long

Depth (ml

T (“a

S

I3

S’16’N, 44”09’W

3279

2.55

/&(I) Wdki9 34.913 2328

cc02 (ll WmolW 2179

245

Af3&2-(2) (pm&kg) 13

E

5”27’N, 44”Ol’W

3772

2.32

34.898

2328

2178

3

A

5”17’N,43”34’W

3990

2.19

34.884

2330

2181

-2

F

5”33’N, 43”36’W

4164

2.05

34.864

2338

2194

-9

H

5’46’N, 43”38’W

4267

1.94 34.850

2343

2202

-13

J

5”45’N, 43”24’W

4342

1.78

34.831

2347

221 I

-17

G

6” ION, 42”53’W

4675

1.49

34.975

2358

2228

-28

(I) Alkalinity and ZCO2 are interpolated from TTO stations 61 and 62 (2) Calculated using the calcite saturation constant of INGLE (1975) They were typically removed from the multi-corer and placed in a cold van at 4°C within 10 min of the return of the corer to the ship. Sectioning and centrifugation, as well as whole-core squeezing, were carried out in the cold van and completed within about 2 h. Centrifuge tubes were stored in ice between filling and centrifugation and between centrifugation and filtration. Despite these precautions, there was some evidence of the NO; artifact that has been noted previously (Berelson et al., 1990; Martin and McCorkle, 1993). as section-centrifuge samples from the upper centimeter of the sediments generally had higher concentrations than shipboard WCS samples and often deviated from concentration vs. depth trends. At two locations (sites A and H), where in situ WCS data were collected, NO; profiles from shipboard WCS were comparable to the in situ profiles (see Fig. 2), but showed some tendency toward higher concentrations over the 0.5-I cm depth interval. In situ WCS was carried out as follows. After the WHIMP landed on the seafloor, a piston core was taken. The volume of overlying water in the core was determined by placement of a light sensor that halted coring when a preset position was reached. When the core was raised out of the sediments, a cup slid underneath it, and the core was lowered into the cup, sealing the bottom. The WHIMP was then raised to 50 m above the seafloor, where whole-core squeezing was carried out. Overlying water, then porewater, was forced through the squeezing piston as it was pushed onto the sediments; samples were filtered in situ and stored in a series of sequentially filled sample containers, whose volume corresponded to sediment depth in increments of about 2 mm (with the exact volume:depth correspondence depending on porosity). These samples were analyzed for NO;, alkalinity, and ZCO,. Cl analyses were used to correct for dilution of porewaters by distilled water in the sampling lines. The dilution corrections were typically 6%. In addition to these porewater samples, deep porewaters (2-60 cm below the sediment-water interface) were obtained using the WHIMP’s in situ samplers. These samples were analyzed for NO.;, SiOZ, alkalinity, CCO?, and Ca2+, Porewaters collected by this technique are occasionally subject to dilution by bottomwater. SiOz data are particularly sensitive to this problem. When the SiO, concentration for a particular sample falls distinctly below the concentration vs. depth trend, alkalinity and ECO- data invariably also show evidence of dilution by bottomwater. Multiple deep porewater profiles were collected at three sites: A, B, and G. At B, thirteen out of fiftysix samples showed evidence of dilution by bottomwater; at A, seven of fifty-six; at G, three of twenty-eight. These samples have not been included in the data interpretation. The suite of porewater samples has been used to estimate organic matter decomposition and CaCO, dissolution rates in the following way. The primary tool for determining the organic matter oxidation rate at each site is a combined porewater NO: profile. Because of the artifact discussed above, all section/centrifuge data from the upper 1 cm are excluded; in some cases, NO; data from l-2 cm also deviate from the trend of the other data, and these are excluded as well. Where in situ WCS data were collected (A and H), these are used to define the sutlicial NO; profile; at the other sites, shipboard WCS data are used. The porewater NO; profile is completed by the addition of the deep. in situ porewater samples. This combined profile

is fit to obtain the NOj production rate as a function of depth below the sediment-water interface. The C:N ratio in the degrading organic matter is determined by joining the NO.; production rate with a model of organic matter oxidation/calcite dissolution in order to fit deep alkalinity, ZCO,, and Ca’+ profiles. The deep part of these profiles is not sensitive to variations in the rate constant for calcite dissolution. but is sensitive to the C:N ratio of the degrading organic matter. The reason for this is that, in the range of CaCO, dissolution rate constants that we observe, the dissolution rate deeper in the sediments is limited by the rate of production of metabolic acids, hence, by the organic matter oxidation rate. The principal effect of variations in the rate constant for dissolution is to change the gradients of ZCO, and alkalinity near the interface, where dissolution of sedimentary CaCO, competes with diffusion of bottomwater carbonate ion. The variation in C:N ratios adjusts the organic matter oxidation rate by a constant factor over the entire depth range of oxic C,, decomposition. The deep portions of ZCOZ and alkalinity profiles see the accumulated effect of this variation over the entire interval; therefore, they are quite sensitive to this sort of change in C:N. Because the effects of variations in C:N and dissolution rate constant are seen in distinct depth intervals, we are able to use our model with the deep part of the profiles to adjust C:N, and with high-resolution profiles near the interface, obtained by in situ WCS at site H, to adjust the dissolution rate constant. This rate constant is used, with bottomwater data and the organic carbon oxidation rate at each site, to predict the calcite dissolution rate. and the result is compared with Cal’ fluxes estimated from in situ porewater samples. 4. ANALYTICAL

METHODS

The analyses of porewater solutes and sediment properties employ standard methods, summarized below: Ca’+ (*0.2%)

EGTA titration

Alkalinity (20.3%) TCOz (20.4%) NO? + NOT (20.2 pmol/l)

Automated Gran titration Gas Chromatography Copper/Cadmium reduction wmechnicon Autosampler Silicomolybdate blue complex w/Technicon Autosampler O2 microelectrode, calibrated by 5 mL Winkler titrations X-ray Diffraction

SiO? (20.5 PmoW Oxygen (23%) Aragonite (22%) % CaCO, (-cOS%) % OrgCarbon (?3%) Chloride (20.3%)

Total Carbon-Organic Carbon Total Carbon After HCl Leach Silver/Silver Chloride plating

(Tsunogai et al., 1968) (Gieskes and Rogers, 1973) Hach Carle GC (Glibert and Loder, 1977) (Glibert and Loder, 1977) (Martin et al., 1991)

(Milliman,

1974)

CHN Analyzer CHN Analyzer Buechler Chloridomter

W. R. Martin and F. L. Sayles

246 5. ORGANIC MATTER OXIDATIONICACO, DISSOLUTION MODEL

Our measurements will be interpreted in terms of a reaction describing the oxidation of organic matter and the consequent release of metabolic acids into porewaters, and reactions describing the fate of these acids. The acids can be neutralized by other porewater solutes, in particular. by CO:- and B(OH); diffusing into the region of rapid C,,, oxidation, or by dissolution of sedimentary CaCO?. 5.1. Model Description The model is constructed similarly to those of Boudreau ( 1987). Archer ( 1991), and Hales et al. ( 1994). The solutes considered and the reactions proposed to determine their porewater distributions are listed in Table 2. Each solute is described by a transport-reaction equation,

porewaters. The concentration of each solute in bottomwater is spccified, as is its gradient at the base of the modeled layer (45 cm at site G, 60 cm at all other sites). Acid-base reactions are assumed to be at equilibrium, allowing differential equations to be combined (Boudreau, 1987). yielding equations for alkalinity, X0,, H+ , B(OH);, and Ca’+. The solid phase portion of the model includes CaCO? and an overall sediment mass balance. Organic carbon (C,,) was excluded: we believe that the C,,, oxidation rate can be described more accurately using solute profiles, which are less affected by the details of bioturbation and by nonsteady-state effects; and. since C,,, is a trace component, it does not affect the sedimentary mass balance significantly. The solid phase CaCOX distribution and mass balance are described by: CaCO, : D,(l

1 ( +Dz 1- $ (@C)+ [Rma +P]=0,

-&(p,C,)

--&,(I

Mass balance:

in which C is the solute concentration, 4 is the porosity, D is the sedimentary diffusion coefficient for the solute, u is the porewater advection velocity (due to sediment burial and compaction only), RHB.Bare proton transfer reactions, and P are reactions causing net production or consumption of the solute. The brackets indicate that the reactions occur for some solutes, but not for others. Calcite dissolution is described by a 4Sth-order rate law (Table 2; based on Keir, 1980); CaCO, is not allowed to precipitate in supersaturated

0, PC, =

021

{ p,(l - +)k,C,(l

Reactions I. Solutes considered: CO2, HCO3-, CO32-, H+, OH-, B(OH)3. B(OH)4-, Ca2+ II. Reactions A. Organic Matter Oxidation (1) (CHzO),(NH3), + (x+2y)O2 -+ xc02 + yN03- + (x+y)HzO + yH+ B. Acid Neutralization Reactions (1) Hz0 + CO2 + CO32H+ + co& (2) CO2 + B(OH)dH+ + B(OH)4-

+ + + +

2HCO3HC03B(OH)3 + HCO3B(OH)3 + Hz0

(3) Hz0 + CO2 + CaC03

+

Ca2+ + 2HCO3-

+

Ca2+ + HC03-

H+ + CaC03 III. Reaction Rates A. C,,rg Oxidation =

Pco* = (c/N)c,,,g, PAlo;= w&,,,~ox (wx

+ w+)

B. H+ production due to organic N oxidation to NOg- = P~o3C. Acid/Base reactions are assumed to be instantaneous. Equilibrium described by the first and second dissociation constants for H2CO3 (Kl and K2); the dissociation constant for water (K,); and the dissociation constant for B(OH)3 (KB). D. Calcite dissolution c2) (kD is the rate constant for calcite dissolution; C, is the calcite concentration) R>l 0, P>(1 - @)k,C‘( 1 - Q),, , R < 1

P&* = PC,;- =

-+)wC,}-Pc,=O.

! R= [C”“][W] KJcalcife) Notes: (1) The reaction is written using the Redfield oxidation state for organic C for clarity; the assumption is not necessary for the model. (2) The rate law for calcite dissolution of KEIR (1980) is assumed.

- n)45,

R < 1

Kinetics of CaCO? dissolution in marine sediment

with the CaCOz and non-CaCO, rain rates, to derive a new solid phase CaCOi distribution. This new profile was used to generate new solute profiles; the procedure was repeated until the change in the solid phase CaCO, profile (the component that converged the slowest) was less than one part in IO’. The sets of differential equations were solved using the relaxation technique described by Press et al. ( 1992) with a depth-varying grid spacing, with Ax increasing linearly from 0.01 cm adjacent to the sediment-water interface to 1.0 cm as X~-Gn,..

Solute Diffusion coefficeints in seawater (cmz/yr, at 2” C)

Solute 02 (1) N03CO2 (1) HC03-

cop B(OH)d (2) H+ OHCa2+

241

D (cm2/yr) 390 330 310 180 140 180 1800 850 120

5.2. Transport

Parameters

Diffusion coefficients in seawater for the modeled solutes, at 2°C are listed in Table 3. They were corrected for the effects of sediment tortuosity using measurements of porosity and formation factor (F: the ratio of sediment resistivity to the resistivity of overlying water, assumed to be similar to porewater) as described previously by Martin et al. ( 1991). The diffusion coefficients of the solutes in sediments are related to those in seawater by

(I) from BROECKERAND PENG (1974) (2) Assumed to equal D for HC03Source of others : LI AND GREGORY(1974) Because we assume that pY is constant. the first term in the mass balance equation is zero. Thus, three boundary conditions are required. At x = 0,

D,

= g

Porosity and resistivity were determined at four sites (B, E, F, G). In each case, a resistivity profile was measured, using a probe with 4 mm electrode spacing ( Andrews and Bennett, 198 1) , on one portion of a recovered core at the in situ temperature. This portion of the core was discarded during subsequent sectioning, and the remainder was saved for porosity determination (Manheim et al., 1974). Porosity and formation factor profiles were similar at B, E. and F (Fig. 1) Porosity decreased, and formation factor increased, more slowly at G; however, this core had a severely sloped sedimentwater interface (a change in elevation across the interface of -2 cm), so that the profiles were uncertain. Because of this uncertainty, an exponential fit to the B, E, and F data was used to describe 4(x) at all sites. The resulting uncertainty at G does not affect conclusions of the modeling effort. Fits of F = 4-” at B, F, and G yielded an average value for n of 2.65 ? 0.54 (Fig. 1). The fit at E yielded a higher value of 4.0, which was not included in the average because the resistivity profile only extended to 4 cm at E. Thus. sediment diffusivity and its variation with depth in the sediments were estimated by Dscd = D,,t$’ ‘, with 4 = .712 + 0.156e’-“-5”“. Solid phase CaCO, profiles on the Ceara Rise show only slight gradients in the upper 20 cm of the sediments, and %CaCO? is greater than 60% at all or our sites except G, where it is 30-40%. Most organic carbon oxidation and CaCO, dissolution are in the upper 5 cm of the sediments (see discussion below). Under these circumstances. the details of the bioturbation mixing regime do not greatly

R,M t RN w=mAtx=.r,,,,,,

C, is the CaCOT concentration (mollg dry sediment), M is the molecular weight of CaCO,, and R is the saturation state of the porewaters with respect to calcite. Da is the bioturbation mixing coefficient; note that, in this formulation, bioturbation mixes sedimentary solids, rather than bulk sediment. p. is the density of sedimentary solids (assumed to be constant at 2.6 g/cm), 4 is porosity, and w is the bulk sediment burial velocity (cm/y). R, is the rain rate of CaCO, (mol/cmk,/y) and R, is the rain rate of noncarbonate sedimentary components (g/cmL/y). The solution phase and solid phase portions of the model were solved separately, using an iterative procedure. First, a solution for the solutes was derived using the C,, oxidation rate derived from NO j profiles (discussed below ) and an arbitrary solid phase CaCO, distribution. This solution yielded a starting depth profile of porewater saturation state with respect to calcite, which was used, along

C

0.6

Porosity (cm$cmi,,) 0.7 0.8 0.9

F (R,&,w) 1.0

1.0 I

1.5 I

2.0

2.5 I 0.0 ,

-0.1-I

-0.14

-0.12

-0.10 -0.08 1% (01

FIG. 1. (A) Porosity and (B) formation factor data from sites B. E. F, and G. The line in (C) is calculated with n = 2.65 in log (l/F) = q.

-0.06

248

W. R. Martin and F. L. Sayles

(4 10

NO, (umol/l) 15 20 2s

30

0, (pmol/l) 160

200

240

280

NO; (umol/l) 0, ~llmol~l) 160

NO, (umol/l) 10 15 20 25 30

200

240

280

0, (umol/l)

O

NOi WOW

150 190 230 270

10

16

20

24

28

-1-r”““”

3

20

g,

30

2 s

8

40

12

0

50

I

3

60

10

NO, (umol/l) 15 20 25

30 160

0, (~mol/l) 200 240

280

FIG. 2. Porewater NOj and O2 profiles, with model fits. For each site, the left-hand panel shows data collected by various techniques (open symbols: in situ, harpoon-type samples, with the different symbols at each site representing different instrument lowerings; closed circles: whole-core squeezer samples; x and +: porewaters collected by shipboard sectioningkentrifugation). The right-hand panels show data collected by whole-core squeezing. In each case, the zero line shows the position chosen for the sediment-water interface. At sites A and H. NO.; data collected by shipboard WCS and by in situ WCS are compared.

affect CaCO? dissolution, We have adopted a moderate value for the surficial bioturbation mixing rate (0:) of 0.3 cmL,/y (Boudreau, 1994). and describe its depth dependence by the simple form, & = D;e-‘*‘L”

with L = 5 cm. Thus, DB (5 cm) = 0.1 cm’/y and DB (10 cm) = 0.005. Preliminary ““Pb data indicate that excess ““Pb reaches about 6 cm at sites B and G. The sedimentation rate is determined by the model, given CaCO, and non-CaCO? rain rates to the seafloor. Values for non-CaCO, rain

Kinetics of CaCO, dissolution in marine sediment

249

NO; (umol/l)

(b)

10 1s 20 25 30

0, (umolll)

15

140 180 220 260

NO;(WOW 20

25

30

NO; (umol/l)

10 15 20 25 30 35

0, (umol/l) 120

0

160 200

240

10

-g 20 ;

30

B 40 Site J

50 PF]

60

NO; (umolll)

5 10 15 20 25 30 35

O?WOW 120 160 200 240

60 FIG. 2. (Continued)

rates were based on Curry and Lohmann ( 1990). The CaCO, rain rate used for model runs was chosen so that the CaCO, accumulation rate predicted by the model matched the CaCO, accumulation rates estimated by Curry and Lohmann (1990) and/or Fran9ois et al. ( 1990) (see discussion below) These values, and the model-derived sedimentation rates, are given in Table 4. The sedimentation rates predicted by the model ranged from I .6-2.3 cm/kyr when the rate constant for calcite dissolution was 0.7%lday; for comparison, Curry and Lohmann ( 1990) estimated rates of 2-2.5 cmlkyr, while Francois et al. (1990) estimated 1.8-2.5 cm/kyr over the last 12.1 kyr. The porewater advection rate due to sediment burial and compaction was approximated using the relationship, UT-

&Wx

4 (Bemer. 1980). 5.3. Bottomwater

Saturation State

The model-based calculations of the rate constant for calcite dissolution that are presented below depend directly on estimates of the saturation state of bottomwater with respect to calcite. These esti-

mates depend on measurements of bottomwater alkalinity and CCOz and on the set of equilibrium constants used. We used comparisons of bottomwater SiOz and NO;, measured during our cruise, with data from TTO stations from the region to determine that ‘lTO60.61, and 62 show the same trends in bottomwater chemistry as our data. Our alkalinity and IZCOz data show the same trends with depth as the ‘IT0 stations, but are offset: our alkalinities are higher by 60 t 6 pmol/kg, while our ZCO,s are lower by 26 2 9 /Imol/kg. Because of this apparent calibration problem, we used values interpolated from ‘IT061 and ‘IT062 as the true bottomwater values, and adjusted our porewater data by subtracting the difference between our bottomwater values and the TTO values. Bottomwater values used are given in Table 1. We have chosen to use the & (calcite) of Ingle ( 1975), the first and second acidity constants for CO, of Mehrbach et al. ( 1973) (using the fitting functions given by Dickson and Miller0 ( 1987)), and the boric acid acidity constant of Lyman ( 1956). All constants are on the NBS pH scale. Pressure coefficients are those of Miller0 ( 1982) and UNESCO ( 1987). The resulting Ksp values are in agreement with the in situ porewater measurements of Sayles ( 1985). This set of constants yields a conservative estimate of above-saturation horizon dissolution, since it results in the shallowest saturation horizon among the available choices of constants. In particular, the ACO:- values in Table 1 are lower

250

W. R. Martin and F. L. Sayles

Sedimentation Rates - Model site B E A A (5) F H J G

& (1) (pmoVcm2/yr) 16.0 16.0 16.0

R,,,, (2) (pg/cm2/yr) 360 360 300

kD = 0.7 c3) 2.3 2.1 1.9

Sedimentation rate (ctiyr)

16.0 16.0 16.0 17.5

400 300 500 500

1.8 I .7 1.8 1.6

kD= 10c4) 2.1 1.8 I .7 1.4 1.4 1.3 1.2 0.9

(1) The rain rate of CaC03 to the sea floor used in model runs (2) The non-CaC03 rain rate to the sea floor used in model runs (3) The model sedimentation rate when k,, = 0.7 %/day (4) The model sedimentation rate when k,, = 10 %/day (5) The model sedimentation rate at A when kD = 150 %/day than those obtained by using the Mucci ( 1983) about 8 pmollkg.

K,, (calcite

) by

6. RESULTS 6.1. Organic Matter Decomposition The general features of porewater NO.7 profiles (Fig. 2) are similar at all sites, with concentrations increasing from the bottomwater value to a maximum at about 10 cm depth and decreasing in a quasilinear fashion below the maximum. The two deepest sites (J and G) are somewhat different from the others. The profile at J has a broader and deeper maximum and a slower rate of decrease below the maximum. At Cl, the NO? concentration decreases much more rapidly below the maximum than at the other sites, and the linear decrease stops in the depth range, 40-48 cm. NOT concentration vs. depth profiles similar to these have been observed in other locations in the equatorial Atlantic Ocean (Froelich et al., 1979; Jahnke et al., 1989). They are indicative of two, spatially separated zones of organic matter oxidation: one in the upper lo-20 cm of the sediments, where oxic decomposition and nitrification are occurring, and one deeper in the sediments, where any remaining O2 is consumed and denitrification occurs. The profiles require that there is a buried layer, rich in organic matter relative to the surrounding sediments, that is consumed at a rate determined by the rate of O2 and/or NO, diffusion from the surface sediments (Wilson et al., 1985). There is good evidence for the existence of such a layer in sediments at the base of the Ceara Rise. A giant gravity core from 4556 m (55GGC, Curry and Lohmann, 1990) shows a very distinct layer centered at 50 cm depth, in which the organic carbon concentration reaches 0.8%, compared to values of 0.45-0.5% in the sediments surrounding the layer. Given the difference in locations between our site and the GGC site, the depth of this layer corresponds well with that required by our NO.7 profile at site G. There is some, less striking evidence for an organic-rich layer at a depth consistent with our NO? profile at site B in a gravity core from a site at 3 164 m (7 IGGC, Curry and Lohmann, 1990). In this gravity core, the organic carbon concentration increased from about 0.26 to 0.38% near 50 cm depth, and a

single analysis at 68 cm showed an elevated concentration of 0.48%. Our NO, profile at site B (3279 m), if the two-zone oxidation model is correct, requires an organic-rich layer somewhere below 60 cm. Thus, solid phase data corroborate the two-zone oxidation model at site G and are consistent with it at B. The two-zone oxidation model requires that the nitrification rate decrease with increasing depth below the sediment-water interface. In addition, it was found that a nitrification rate described by a single exponential function could not describe both the surlicial NO; and O2 gradients and the deep NO; profiles. Thus, a double exponential function was adopted, with

PNOi= j,,e7’ ’ + j,e

-'lr .

The fitting procedure used ( Levenburg-Marquardt, modified from Press et al., 1992) does not constrainj, effectively. Thus, an initial analysis of WCS O2 and NO.; data was carried out to determine a value of j, that would fit most of the data. An optimal j, was determined using a series of fits, each using a fixed j, and allowing the other ji to vary; different j, values were used for each fit. j, = 2 cm-’ was found to be optimal. The only exception was site E, where a value of 3 cm-’ was needed. With these j, values, the equation was fit to the porewater data using jO, j,, j,, and the gradient at the maximum sampled depth (except at site G, where the gradient at 45 cm was used) as fitting parameters. The fits (Fig. 2, Table 5) defined both the integrated nitrification rate and the distribution of nitrification as a function of depth in the sediments. Site-to-site differences in the j, values reflect variations in the distribution of nitrification with depth below the sedimentwater interface. The sites are generally similar, with two exceptions: site E, where organic matter oxidation is required to occur closer to the interface than at the other sites (larger j, and j, ); and site H, where organic matter oxidation appears to occur deeper in the sediments (relatively small j, and large jz ) . In order to use these results to define the organic carbon oxidation rate, the C:N ratio of the degrading organic matter must be constrained. Porewater 0, data can be used, in conjunction with NO? data, for this purpose. This procedure has the drawback of requiring an assumption about the oxidation state of organic carbon, which may not be well known (Anderson and Sarmiento, 1994; Takahashi et al., 1985). In our case, there is an additional limitation, that the O2 data extend only to 2.5-3 cm below the sediment-water interface, while oxic decomposition occurs over a much longer depth interval.

Best Fits to NO3- Profiles Site B E A F H J G

JO pm01 cm-3Dwyrl 3.93 10.4 3.10 3.02 0.528 3.56 2.39

JI

J2

cm-l 2.0 3.0 2.0 2.0 2.0 2.0 2.0

pm01 cm-3Dwyr l 0.082 0.099 0.144 0.191 0.514 0.11 I 0.091

J3 cm-l 0.189 0.206 0.238 0.298 0.414 0.186 0.170

251

Kinetics of CaCO? dissolution in marine sediment Our shipboard WCS data (Fig. 3) indicate an average ratio of OZ consumption to NO1 production of 12.5, implying a C:N ratio of 8.8 (Anderson and Sarmiento, 1994 oxidation states) to 10.5 (Redfield oxidation states). We have taken an alternative approach to constraining the C:N ratio, which is based on the observation that porewater alkalinity, X0,, and Ca*+ profiles are insensitive to the rate constant for calcite dissolution below about 5 cm depth, but are sensitive to the organic carbon oxidation rate; given the nitrification rate determined by fits to NO; profiles, the C,, oxidation rate is determined by the C:N ratio of the degrading organic matter. Thus, we use in situ porewater profiles, bottomwater saturation state calculations, and our organic matter oxidationlcalcite dissolution model to constrain the C:N ratio. We assume that the C:N ratio is constant with depth in the sediments. In most instances, the procedure yields an uncertainty of about 2 1 mol/mol. The exceptions (Fig. 4) are site F, where scatter in the porewater data limits the estimate to an uncertainty of 42, and site G, where the model does not explain some features of the porewater data well (see below ) The C:N ratios estimated in this way vary from 8- 11 mol/mol, with no apparent spatial pattern. The average value is 9.5 t 1.2. This value is within the range of estimates based solely on WCS data, which cover a depth range of O-3 cm below the sediment-water interface. The agreement between C:N derived exclusively from shallow data and from deeper porewater data indicates that, within our uncertainty of about 1 mol/mol, there is no evidence for variation of C:N with depth below the interface at our study sites. The C:N derived from fits to deep porewater data should apply to the depth range of our in situ WCS profiles as well. Porewater O2 profiles, determined by shipboard WCS, were not used directly in the determination of organic matter oxidation rates. The O2 profiles implied by the nitrification rates obtained from fits to NO; profiles, C:N ratios determined as described above, and an assumption of the Redfield oxidation state for organic carbon are shown in Fig. 2. The model profiles fit the gradient at the surface in each profile except at site E (wherej, = 3 cm-’ was used to obtain the NOj fit). However, there is a break in slope apparent in the shipboard WCS profiles at most sites that is not matched in model profiles. We believe that this break is due to a difference in design between the squeezing piston in the shipboard WCS vs. that in the in situ WCS. While the in situ WCS piston has five holes for channelling pore solution to sample containers, approximately evenly spaced on the face in contact with the sediments, the shipboard WCS piston used on this cruise has only a single, central hole. This design results in a longer path length of the pore fluid across the piston face, increasing the possibility of mixing during sampling as well as the pressure required for squeezing. This increased pressure could cause dissolution of small amounts of air trapped in o-ring grooves, resulting in decreasing porewater O2 gradients as squeezing pressure increases. It is because of the possibility of this artifact, and the fact that the WCS 0, profiles only cover a small part of the depth range of oxic Core decomposition, that we have not used these OZ data to constrain Corgoxidation rates. The integrated NO7 production rates and C:N ratios allow us to estimate the organic matter oxidation rates at our study sites (Fig. 5 ) . The highest rate is observed at E (3772 m) ,

-80 -60

ope = 10.51 zk0.32

0

2

I I 4 6 ANO; (pmol/l)

I 8

I 10

Rc;. 3. AOz is the porewater O2 concentration, minus the bottomwater concentration. ANO; is the same quantity for NO;. The data are from shipboard WCS, with each site represented by a different symbol. Two C:N values were estimated from the slope of the bestfit line, using the diffusion coefficients from Table 3 and the (I) Redfield or (2) Anderson and Sarmiento (1994) oxidation state for c,, where we calculate 26 pmol/cm’/y. Apart from site E, the data show a general, gradual decrease in C,, oxidation rate from 22 pmol/cm’/y at site B (3279 m) to 14 at G; site H had the lowest value, 10 pmol/cm2/y. 6.2. Calcite Dissolution Our interpretation of porewater NO; profiles implies that organic matter decomposition uses O2 as an electron acceptor throughout the depth interval sampled (O-60 cm, except at G, O-45 cm). Thus, the production of alkalinity and Ca2+ in near-interface porewaters would provide evidence of CaCOl dissolution (Emerson and Bender, 1981). In situ porewater profiles from all sites (Fig. 4), including B and E, which lie above the calcite saturation horizon, clearly indicate alkalinity and Ca’+ production in the upper 15 cm of the sediment column. Unless there is precipitation of CaCO? at the sedimentwater interface, these data indicate that there is net CaCOj dissolution above the calcite saturation horizon. High-resolution profiles of alkalinity and X02 were collected by in situ WCS at sites A (ACO:= -2) and H ( ACO:- = - 13 ) . Alkalinity profiles increase regularly with depth, and are consistent with profiles predicted by the Core oxidation/calcite dissolution model (Fig. 6) _ CC02 profiles are more problematic: at site A, the deepest points in the WCS profile fall below model predictions; at H, the data are somewhat scattered and most fall below the model lines. At H, the agreement between prediction and measurements would be much better if the interface were placed 2 mm deeper; although such a placement would be consistent with alkalinity and NO? profiles (Fig. 2), it is not consistent with the SiOZ profile, which appears to require the interface shown. In general, the uncertainty in placement of the interface is significant, as the decision is subjective, based on observed breaks in slope of porewater profiles; the uncertainty is approximately equal to the sample spacing, 2 mm in this case.

Depth (cm)

-_-_- -:+----,_-- __-2._--

D

,,: -* __-, ---G c

Depth (cm)

ir

42 0b

b3

i;-

E

Y

8 =:

D

Depth (cm)

Depth (cm)

Depth (cm)

Kinetics of CaCO, dissolution in marine sediment

2400

Alkahnity 2800

3200

60

2200

~CO, 2600

253

3000

60

Alkalinity (peq/kg) 24co 2800 3200

2200

ZCO, (~mdkg) 2600 3000

60

Alkalinity (psfkg) 2400 2800 3200

2300

XCO, (pin&kg) 2700 3100

0 10 z g B Q

20

z u 5 & a

30, 40 50

20 30 40 50 60

FIG.

4. (Continued)

WCS alkalinity data could, in principle, be subject to artifacts due to chromatographic effects, primarily involving H + , occurring as deeper porewaters pass through surface sediments. These effects would not be expected to alter X02 profiles, as they are not sensitive to changes in H + _ Thus, if the artifact is present, the slope of an alkalinity vs. XOI, plot should deviate from the trend defined by other in situ porewater data. At sites A and H, the trend in alkalinity vs. CC02 for the WCS samples is indistinguishable (within the limits imposed by the scatter in the ZCOl data) from that of the other porewater data (Fig. 7) : there does not appear to be an artifact affecting WCS alkalinity data at these sites. We have used the WCS alkalinity data at site H to constrain the calcite dissolution rate constant. First, we estimate the alkalinity flux from the porewater profile. In order to set limits that are independent of dissolution models, we have estimated a range of possible fluxes: the lower limit is set by a linear fit to the top two porewater points; the upper limit

by an exponential fit to the entire profile. The resulting range is 14-25 peqlcm21y. Then, we generate a curve of alkalinity flux vs. kl,, the rate constant for calcite dissolution. The curve depends on several things: the saturation state of bottomwater, the organic matter oxidation rate and its distribution within the sediment column, the rate law for calcite dissolution, transport parameters, and the accuracy of the transport/reaction model. Clearly, the derived rate constant is model-dependent. The range of values derived from this treatment is 0.7-40%/day (Fig. SB). The k,, that corresponds to the mean of our alkalinity flux estimates is 8%/ day; allowing for the uncertainty in organic carbon oxidation rates implied by our 21 mol/mol uncertainty in C:N, the range corresponding to the mean alkalinity flux estimate is 5.5-lOS%/day. The range in k,, 0.7-40%lday, is consistent with other estimates based on in situ data, using both benthic flux chambers and microelectrode profiling (Table 6). The one exception is the range of values estimated by

and F.

W. R. Martin

254 A

2.0

C

B

C : N (mol/mol)

Net NO; Prod.(~moUcmz/yr) 1.0

L. Sayles

3.0

4.0

6

8

10

C,, Oxidation rate (~molJcm*/yr) 10 15 20 25 30

12

3200

3200

3600

3600

l’

I

4800

4800

4800 1

FIG. 5. (A) Net NOi production rate at each site, calculated as the integral of PNo;~(x) with the best fit j, values (Table 5). (B) The C:N ratio estimated at each site. (C) The organic carbon oxidation rate at each site, calculated as the

product of the integrated NO; production rate and the C:N ratio. Jahnke et al. ( 1994); the low alkalinity fluxes they determined required k,, values less than O.S%/day. The uncertainty associated with this approach to estimating k, is much greater at site A than at H. Since A is very near the calcite saturation horizon, nearly the entire alkalinity flux can be attributed to neutralization of metabolic acids ( >99%, compared to 76-92% at H; see discussion below). As a consequence, the rate of calcite dissolution is limited by the rate of production of metabolic acids, and only depends weakly on the rate constant for dissolution. Thus, uncertainties in the alkalinity flux and in the organic carbon oxidation rate result Alkalinity 2300

2450

x0, 2600

2150

2300

2600

2150

2300

Alkalinity 2300

2450

2450

x0, 2450

in considerable uncertainty in the estimated k,, value. The alkalinity flux estimated from a linear fit to the upper points in the profile at A is 20 peq/cm’/y. Using this value and an estimated uncertainty of +25%, the range of possible kD values is 5 to > lOOO%/day (Fig. 8A). The k,, associated with the mean flux estimate is 150%/day, and the range in this mean value associated with a 5 1 mol/mol uncertainty in C:N is 40-650%/day. Several considerations lead us to adopt a relatively small range for kr, in the calculations that follow. Model fits to alkalinity profiles (Fig. 6) show that, at site H, the range of kr, values, 0.7- lO%/day, explains the data equally as well as the larger range, 0.7-40. Two fits to the site A data, for klj = 20 and 300, show that the alkalinity profile does not allow us to distinguish between these kn values. Further, when we apply the dissolution model to the most undersaturated sites (J and G), we find that use of kl, values in excess of lO%/day would require a significantly larger CaCO, rain to the seafloor than at the other sites, a circumstance that would be difficult to explain given the proximity of the sites. Thus, in the absence of k,, estimates at each site, we use a small range, 0.7-lo%/ day, that is consistent with our best estimates of k,, i.e., those from site H. This approach gives conservative estimates of CaCO? dissolution rates. Note, however, that there are two possible explanations for the difference in kn estimates at A and H, in addition to our assertion that the value at A may be overestimated: the difference may be real, and there may be as yet unexplained differences in k,, over small spatial scales; or the 45th order rate law for calcite dissolution may overstate the dependence of k,, on the bottomwater saturation state. Our limited dataset does not allow us to evaluate these possibilities.

FIG. 6. Porewater X0, (pmol/kg) and Alkalinity @q/kg) profiles, collected by in situ whole-core squeezing, at sites A and H. The site A plots include alkalinity and ZCOl profiles generated by the C,,,, oxidationKaC0, dissolution model with k,, = 20%/day and 300%/day (dashed lines). The site H plots include model profiles with kn = lO%/day (solid line), 0.7%/day, and 40%/day (dashed lines).

The genera1 features of the porewater profiles at all sites are consistent with the organic matter oxidation/calcite dissolution mode1 (Figs. 4, 6). It is noticeable, though, that at several sites the model curve that best fits the deeper porewater data ( 2 15 cm) follows the upper concentration limit of the shallower data. This tendency could reflect uncertainties in the assignment of depths to the fixed sample probes (depths > 15 cm). Because of the tendency of the instrument to sink into the sediments, the depth assignments could be as much

Kinetics of CaCOi dissolution as 3 cm shallower than the true depths. Making such a correction would improve the fit in several cases, but we have used the nominal assignments in the absence of concrete evidence for changing them. The effect of such a change on the estimated C:N ratios would be small (< 1 mol/mol) Measured porewater A Alk:ACC02 and ACa*+ :A Alk are also, in all but two instances, consistent with model predictions (Fig. 7). The former ratio is particularly diagnostic of coupled organic carbon oxidation/CaCO, dissolution, as it is slightly less than 0 in the absence of CaC03 dissolution in oxic sediments, and slightly less than 1 when CaCO, is dissolving. The first of the two exceptions to the good agreement between model and data is the Ca’+:Alk ratio at site G: the observed ACa*’ is consistently smaller than that required by the model. The Ca*’ profiles at G are different from those at the other sites in one respect. While, at the other sites, the Ca’+ flux into the sampled layer from depth ( implied by the linear gradient at the base of the sampled layer) is approximately 0.5 times the alkalinity flux, as CaCOj dissolution/ oxic organic matter decomposition requires, the flux ratio is about 0.25 at G. However, since we have taken this into account in the model by specifying the measured Ca” gradient at the base of the modeled layer, 45 cm, we do not believe that this difference can explain the site G result. There is no measurable CaCO, in the sediments at a depth of 30 cm; however, in the upper 20 cm, where our NO.? data indicate that organic matter oxidation is occurring, CaCO, concentrations do not fall below 30%. In short, we do not have an explanation for the departure of the ACa’+ :A Alk ratio from the model prediction at G. The second exception to the general agreement between model and measured ratios is the AAlk:AX02 ratio at site B. The measurements indicate a higher ratio than that predicted by the model. This deviation from the model may be related to the presence of aragonite at B, our shallowest site (3279 m). We will discuss this in detail below. We must specify the rain rate of CaC03 to the seafloor for each model run. Since the Ceara Rise extends over a very limited spatial area, it is reasonable to expect that this rain rate is approximately the same at all the sampled sites. Thus, the combination of the dissolution model with our estimates of bottomwater saturation states, organic carbon oxidation rates, and kD should be able to reproduce measured CaCO? accumulation rates with an approximately constant rain rate of CaCO?, despite the variation in CaCO, dissolution rates. Curry and Lohmann ( 1990) and Franqois et al. ( 1990) have estimated CaCOX accumulation rates on the Ceara Rise, and we have made a series of model runs to see if we can reproduce their accumulation rates without having to invoke large variations in rain rates between sites. The depths of our sampling sites are not always the same as those of the sites at which accumulation rates were estimated (depths are very close for B, E, A, and J). and there are differences in accumulation rates estimated by the different authors (7.1 pmol CaCO,/cm’ly by Curry and Lohmann (1990) at 4341 m vs. 9.4 by FranGois et al. (1990); 8.4 at 4556 m by Curry and Lohmann ( 1990) vs. 5.3 by Franqois et al. ( 1990)). The most important difference in depths is for our site G, which is 120 m deeper than the deepest site at which accumulation rates were estimated, and has a lower CaC07 concentration ( -35% in

in marine

255

sediment

the upper 10 cm at our site vs. -48% at the Curry and Lohmann 4556 m site). Our data and model do reproduce the measured CaC03 accumulation rates with a constant rain rate of CaCOJ to the seafloor of 16.5 + 1 ~mol/cm’/y, with kn allowed to range from 0.7- lO%/day. There is some systematic variation in the kn required to match model to measured accumulation rates, however, as the two shallowest sites, B and E, appear to require k, near our upper limit of lO%/day, while the two deepest sites, J and G, appear to require kD nearer the lower limit of 0.7%/day. Although the data are insufficient to allow us to draw firm conclusions, this apparent systematic variation in k,) may imply the 4.5th order rate law that we use in our model overestimates the dependence of dissolution rate on bottomwater saturation state. 7. DISCUSSION

7.1. CaCO, Dissolution on the Ceara Rise The evaluation of the rate of dissolution of CaCO? in sediments above the lysocline and between the lysocline and calcite compensation depth is important to the quantification of oceanic CaCO? cycling. Milliman ( 1993), for instance, has estimated that 45% of marine CaC07 production and 34% of CaC07 accumulation occur in the deep sea. In addition, he concludes that, since 45% of the Atlantic seafloor lies above the lysocline and 94% above the CCD, the Atlantic accounts for over 40% of deep-sea CaC03 accumulation. These conclusions were reached indirectly, from estimates of CaCO? production and the assumptions that, on average, 80% of CaCOl reaching the seafloor above the lysocline and 40% of that falling within the lysocline are preserved. A direct evaluation of these assumptions is essential to accurate quantification of the CaCOj cycle in the modem ocean. Knowledge of the dissolution rate of CaCOJ in above-lysocline sediments is also important to the interpretation of the past history of CaCO, accumulation (e.g., Curry and Lohmann, 1990; Francois et al., 1990). In particular, if significant above-lysocline dissolution occurs, then inferences of CaCO, production rates from above-lysocline accumulation rates are inaccurate. Further, variations in these accumulation rates depend on all the factors that influence the coupling between organic matter oxidation and CaCO? dissolution, and may not reflect variations in surface water CaCOz production in a simple way. We have estimated the CaC07 dissolution rate in three different ways. The first (Sayles, 1981) uses fluxes calculated from Ca” profiles from in situ-collected porewater samples. The profiles are fit with an exponential function, and the flux is calculated from the slope of the function at x = 0 (stars in Fig. 9; WHIMP Ca *+ in Table 7). The assumption implicit in this procedure is that the profile has smooth curvature between the shallowest sample depth (2 or 3 cm) and the interface. The assumption would obviously break down if CaC03 were precipitating at the sediment-water interface. Although we cannot disprove such precipitation, since the calculated Ca” fluxes are similar in magnitude to measured CaCO? accumulation rates, it would require that a significant fraction of the accumulating CaC07 was formed at the sediment-water interface. Such an occurrence is unlikely. Benthic foraminifera make up only a very small fraction of total sedimentary CaCO, . In addition, the occurrence of such a large proportion

W. R. Martin and F. L. Sayles

256

V

ta) 3200

2400 2200

2600 x0,

3000

2600

3000

2400

2800

3200

Alkalinity

,x

:g 2800 7 2 4 2400 2200

2400

2800 Alkalinity

3200 10.8

x .z 2 2800 B u

10.6

2400 2200

2600 =m2

3000

2400

2800

3200

Alkalinity

2400 2200

2600 x0,

3000

2400

2800

3200

Alkalinity

FIG.7. Alkalinity: CCOz and Ca’+: Alkalinity relationships for in situ collected porewaters at each sample site. Open symbols represent samples collected by the WHIMP’s harpoon-type samplers.Closed symbols (sites A and H) represent samples collected by in situ whole-core squeezing. Stars represent bottomwater samples. The solid lines show the relationships predicted by the Corg oxidationKaC0, dissolution model. Alkalinity data are in peq/kg, CCOZ in pmolkg, and Ca’+ in mmol/kg.

of inorganic CaCO, would be readily apparent and has not been observed. The smooth curvature assumption may also be violated less drastically, within the context of our CaCO, dissolution model. In this case, errors in estimating Ca2+ fluxes would result from the relatively coarse resolution of the WHIMP porewater data. We evaluated this possibility by estimating

fluxes from model-derived Ca*+ profiles at each of our sampling sites. These “synthetic” profiles were created by choosing points from model-generated Ca” vs. depth distributions, such that the synthetic profile had the same sample depths as the measured Ca2+ profile at the sample site. Then, the Ca2’ flux was estimated from the synthetic profile using the same fitting procedure as was applied to measured Ca2’ profiles.

Kinetics of CaCO, dissolution in marine sediment

257

3200 10.8

h .s 2 4

2800

10.6

2400 2400

2800

3200

2400 2300

2700

3100

2400

x0,

2800

3200

Alkalinity FIG.

7. (Continued)

The exponential fit to points taken from the synthetic profile at our shallowest site, B, yielded the same flux as that catculated from the complete model profile (which has a resolution of 0.01 cm at the interface). At all other sites, exponential fits to coarse-resolution, synthetic Ca*+ profiles underestimate the Ca2+ flux; although the degree of the underestimate depends on several factors, it generally increases as the degree of undersaturation of the bottomwater increases. The effect of the possible underestimation of CaZf fluxes due to the coarse resolution of the porewater data is shown in Fig. 9 by the one-sided error bars attached to the stars denoting the fluxes calculated from measured CaZf profiles. The second method we used to estimate CaC03 dissolution rates is model-based. It uses the depth distribution of organic matter oxidation and the bottomwater saturation state at each site, the range of kD values estimated from the in situ WCS alkalinity profile/dissolution model for site H (0.7-lo%/ day), and the organic matter oxidation/calcite dissolution model to infer the Ca*’ flux at each site. The results of these calculations are shown by the heavy horizontal bars in Fig. 9.

Finally, the most direct estimates of dissolution rates are inferred from the high-resolution alkalinity profiles at sites A and H. At A, which is essentially at the calcite saturation horizon, the inferred Ca*+ flux is 11 ,umol/cm2/y; at H, the Ca*’ flux is 7.6. These dissolution rate estimates, as well as the others, described above, are shown in Fig. 9. The modelbased estimates (thick horizontal bars in Fig. 9) have minimum values (fork,, = 0.7) of 5.3-7.2 at the sites at and above the calcite saturation horizon (B, E, A), and increase fairly regularly below the saturation horizon to a value of 12 at site G. The Ca’+ profile-based estimates are lower in most cases, but, with the exception of site G, show the same general trend. As we noted above, the Cal+ data at G are anomalous, in that the ACa2+ :A Alk ratio is lower than can be explained by generally accepted reaction stoichiometry. The calculation of the fraction of the CaCOj rain dissolving in the sediments is shown in Table 7. At each site, the CaCO? accumulation rate was derived from the measurements of Curry and Lohmann ( 1990) and/or Fratqois et al. ( 1990). The fractional dissolution was estimated as dissolutionl(dissolution + accumulation). Values derived from coarse-reso-

W. R. Martin and F. L. Sayles

258

A. Site A

1

B. Site H

10 kd (% day-‘)

100

1000

0.01

0.1

10 1 b (% day-‘)

loo

FIG. 8. The estimation of k, values using alkalinity fluxes calculated as described in the text from in situ WCS porewater samples. The range of alkalinity fluxes shown on the left-hand axis (in peqlcm’ly) is derived from the porewater profiles. The solid line and its flanking dotted lines show the relationship between k,, and alkalinity flux derived from the C,, oxidationKaC0, dissolution model. The range of k,, values, determined from the range of alkalinity fluxes using this relationship, is shown by the dashedlines. The value of k. that correspondsto the mean of the alkalinity flux estimates is indicated by an arrow, and the range in kD allowed by the uncertainty in C:N is shown by the dotted lines on either side of this arrow. (A) Site A. (B) Site H

lution Ca*+ profiles at the three sites at and above the saturation horizon suggest that about 40% of the CaCOz rain is dissolved in this region; model-based estimates for those sites suggest 40% at B (700 m above the saturation horizon) and 55-60% at E and A; the in situ WCS alkalinity profile suggests a value of 68% at A. These calculations cover essentially the same range of values as Hales et al. ( 1994) estimated, using in situ OZ and pH data with a dissolution model, for the continental slope and rise of the western north Atlantic. The shallowest of the Hales et al. (1994) stations was 2000 m above the calcite saturation horizon. Together, these results, although still limited in scope, suggest that about 40% of CaCO, raining to the seafloor above the lysocline dissolves. This dissolution rate would result in a decrease of 25% in the Milliman ( 1993 ) estimate of CaCOX accumulation rate in this depth interval. Our data from undersaturated sites (F, H, J, G) suggest that a minimum of 50% of the CaCO? rain dissolves (the Ca2+ profile-based estimates); model-based estimates suggest that 60-70% of the CaC03 rain dissolves over the 4000-4675 m depth interval (compared to the regional CCD at about 5200 m (Biscaye et al., 1976)). Applying the reasoning, that the CaC03 preservation rate between the calcite saturation horizon and CCD is about half the rate at the

COmpariSOnS

References BERELSON ET AL., 1994 JAHNKE ET AL., 1994 HALES ET AL.. 1994 ARCHER ET AL., 1989 This work

saturation horizon (Milliman, 1993). these results suggest a preservation rate of 20-25% over that depth interval. If these preservation estimates are applied to the Milliman ( 1993) estimates of CaCO? production, they suggest that 47% of CaCO, production is preserved in Atlantic ocean sediments. If the results are extended further, to the entire ocean, then the estimated preservation rate of CaC07 in the deep sea is 35%, and the estimate of the accumulation rate of CaCOj in the deep sea is 27% smaller than that of Milliman ( 1993) (0.8 vs. 1.1 x lo9 tons/y). Since the existing data are limited in scope and since direct benthic flux and porewater measurements do not always yield similar dissolution flux estimates (Hales and Emerson, 1994; Jahnke, 1994), these conclusions would be strengthened by CaCO? dissolution rate measurements in a broader set of locations. 7.2. Aragonite We have estimated the saturation state of bottomwater with respect to aragonite at our shallowest site (B: 3279 m) using the relationship between Kbp (aragonite) and KS, (calcite) of Mucci (1983), K,, (calcite) of Ingle (1975), and the AV for aragonite dissolution from Miller0 ( 1982). The calculation

of ko VaheS from in situ measurements Method

Location

BFC (I) BFC MEP (2) MEP WCS

E. Equa. Pacific W. Equa. Pacific N.W. Atlantic E. Equa. Atlantic W. Equa.Atlantic

ko %/day (3) 20flO

0.05-0.5 3-30 10-100 0.7-40

(I) BFC = In situ Benthic Flux Chamber (2) MEP = Microelectrode Profiler (3) All based on the rate law, Diss rate = kD [CaC03] (1 - Q)4.5

Kinetics of CaCO, dissolution in marine sediment Ca2+ Flux (pmol/cm’/yr)

3500 -

3 G 3 4000 n

4500 -

The probable dissolution of aragonite at B implies that some of the Ca”+ flux is from this source, rather than from calcite dissolution driven by organic matter oxidation. We estimate that the Cal’ flux is about 36-42% of the rain of CaCO, to the seafloor at this site, a value which is well above estimates of the proportion of aragonite in the CaCOJ rain. Although no data are available from the Ceara Rise region, Fabry ( 1990) and Fabry and Deuser ( 1991) have estimated that aragonite is 8-15% of the CaCO, flux at locations in both the N. Atlantic and Pacific oceans. The proportion of aragonite may be smaller at our sites, since the sediments lie well below the aragonite saturation horizon. In addition, some of the aragonite at B appears to be accumulating in the sediments, further decreasing its possible contribution to the dissolution flux. Thus, it seems likely that less than ‘& of the Cal+ flux at B is derived from undersaturation-driven aragonite dissolution, and that the majority must be due to organic matter oxidation-driven calcite dissolution. 7.3. Organic Matter Oxidation

Estimatedfrom WCS Alkalinity FIG. 9. Estimates of the Ca” flux, resulting from the dissolution of sedimentary CaCO,, at each sampling site. The stars are the fluxes estimated directly from porewater Ca” profiles; the one-sided error bars show the model-based estimate of the amount by which these calculated Ca’+ fluxes would underestimate true Ca’+ fluxes, assuming the model is correct. These possible underestimateswere calculated using k[, = O.‘l%lday; at two sites (F and J), the result for kD = IO%/day is shown as well. The heavy horizontal bars are the range of fluxes, with k,, = 0.7- lO%/day, calculated from the C,, oxidation/ CaCOz dissolution model at each site. The open circles (sites A and H) are the dissolution fluxes determined from high-resolution alkalinity profiles at sites A and H.

indicates that the bottomwater at the site is well below saturation, with ACO:- (aragonite) = -30 pmol/kg. Under these conditions, models would predict that there would be no aragonite present in the sediments (Hales et al., 1994). However, measurements of aragonite at site B do not support this prediction. Six samples, from O-O.5 to 12 cm depth, have aragonite concentrations of 7-l 1% (?2%). Although the quantification is somewhat uncertain, the presence of aragonite is clearly indicated. A sample at 30 cm at this site had no detectable aragonite; further, surficial (O-O.4 cm) sediments from all the remaining sites, as well as profiles from sites A and G, showed no detectable aragonite. We noted above that the AAlk:AX02 data at site B deviated from the prediction of the organic matter oxidation/ calcite dissolution model: the model curve indicates proportionately less alkalinity production than the data show (Fig. 7 ) The disagreement between model and data is due largely to the buildup of metabolic acids near the interface that the model requires to overcome the supersaturation of the bottomwater with respect to calcite before dissolution (and alkalinity production) can occur. A possible reason for the disagreement would be the dissolution of aragonite, which would begin at the sediment-water interface.

259

and CaCOj Dissolution

Our results, based on fluxes calculated directly from porewater Ca’+ profiles as well as on model-dependent calculations, indicate that there is significant dissolution of CaC03 at and above the calcite lysocline. The dissolution flux at the shallowest site, B, probably includes a contribution from aragonite dissolution; at E. where no aragonite was detected in the sediments, the dissolution must be driven by organic matter oxidation. The coupled organic matter oxidation/calcite dissolution model indicates that organic matter oxidation is important to dissolution at all of our sites, including those with undersaturated bottomwater (Table 8); we estimate that 4060% of calcite dissolution at our most undersaturated site is linked to organic matter oxidation. These results complicate the interpretation of sedimentary records of CaCO, accumulation. The occurrence of sedimentary CaCO? dissolution above the lysocline means that the estimation of surface ocean CaCO? production from CaCO.l accumulation rates must include a correction factor to account for the dissolution. The interpretation of temporal and spatial variability in accumulation rates must take into account the factors that affect the dissolution flux: the degree of saturation of bottomwaters, the ratio of organic carbon to CaCOl in the particle rain to the seafloor, the kinetics and stoichiometry of organic matter oxidation and calcite dissolution. If the model-based inference, that dissolution linked to organic matter oxidation is important to the dissolution flux below the saturation horizon as well as above it, is correct, then changes in dissolution rates below the saturation horizon may not be simply linked to bottomwater chemistry. In this section, we examine the factors determining the importance of organic matter oxidation to the dissolution flux on the Ceara Rise. Using our model-based estimate of the rain rate of CaC03 to the seafloor on the Ceara Rise ( 16.5 pmol/cm’/y), we can estimate the C,,,:CaC03 ratio in the particle rain (Table 8). The average ratio is 1.1 + 0.3 molImo1. Because our estimated organic carbon oxidation rates tend to decrease as water column depth increases, the ratio decreases from values greater than 1 at the shallower sites to values less than 1 at the deeper sites. This ratio gives the maximum dissolution

W. R. Martin and F. L. Sayres

260

Fraction of CaCO3 Rain that Dissolves

Site

MEASURED In situ Alkalinity WHIMP Ca2+ F&l 90 dissolved3 Fcac4) 70 dissolved 36 5.1

Est Act Rate(‘) pnol cm-2 yil

9.1

MODE1 tiTEs ko = 150(G) kD = 0.7-10 %/day(s) Fca

70 dissolved

5.3-6.5

37-42

FP,

% dissolved

(1) CaCO3 accumulation rates. From CURRY AND LOHMANN (1990) and FRANCOISET AL. (1990) (2) Caz+ flux calculated from Ca2+ profiles from pore waters collected in situ. In pmol/cm2/yr. (3) % of the CaC03 rain that dissolves; calculated from Fca / (Fca + Est. Act. Rate) (4) CaZ+ flux determined by the organic matter oxidation I CaC03 dissolution model with parameters chosen to match the alkalinity flux calculated by a linear fit to the first two points of the in situ WCS alkalinity profile. (5) The range of Ca2+ flux calculated by the C,,r oxidation / CaCO3 dissolution model at each site with the optimal kD range estimated for site fi. (6) The Ca2+ flux estimated for site A with kD=l50 %/day

tion of C,, oxidation decreases as undersaturation increases, at a rate that depends on the calcite dissolution rate constant: with /Q, = 0.7%/day, the decrease is from 97% at site F to 60% at G; with kD = 10, the decrease is more dramatic, from 92-37% (Table 8). The efficiency with which organic matter oxidation drives dissolution shows two general features when plotted vs. bottomwater ACO:- (Fig. 10). First, dissolution models predict that, if reaction kinetics do not vary too much between sites, there will be a general trend toward increasing efficiency, from about 25 to about 50%, as undersaturation increases (the marked trend in Fig. 10). The general trend can be seen in our data. This trend can be explained by the way in which porewater gradients of the solutes that buffer seawater, thereby limiting the extent to which metabolic acids drive dissolution, vary with bottomwater saturation state. At both the most supersaturated site (B) and the most undersaturated site (G), the dissolution model predicts gradients of CO:and B (OH ); that lead to fluxes into the region of most rapid

that can be driven by organic matter oxidation. The actual dissolution can be estimated by running the coupled Corgoxidation/CaC03 dissolution model twice for each site: once with optimized parameters, once with the organic carbon oxidation rate set to 0. The difference gives the dissolution flux that is driven by organic matter oxidation (Hales et al., 1994). In addition, the efficiency with which organic carbon oxidation drives dissolution can be estimated as the ratio of the above result to the organic carbon oxidation rate. We have carried out these calculations at our study sites (Table 8, Fig. 10). Our model is a calcite dissolution model. To the extent that sedimentary aragonite dissolution is driven by undersaturation of bottomwaters with respect to that mineral, we will overestimate the contribution of dissolution driven by organic matter ox ,dation. This may be significant at site B; we do not believe it is a factor at the other sites. In the context of our model, all the dissolution at and above the calcite saturation horizon is driven by organic matter oxidation. The contribu-

E *A

1.6

100

1.1

99

F H J G

1.2 0.65 0.93 0.85

97-92 92-16 84-62 60-(37)

27-37 60 44-56 63-73 50-54 Sl-(45)

(1) The fraction of CaC03 dissolution that is attributable to neutralization of

metabolic acids producedby the oxidation of organic matter. Calculatedusing ko = 0.7-10 %/day, except for site A, for which kD= lSO%/day was used. (2) The rate of dissolution of CaCO3 due to the neutralization of metabolic acids, divided by the organic C oxidation rate (as %). The range is for kD = 0.7-10 %/day. The values for site G, especially with kD=lO%/day, are low because the sedimentary %CaCO3 becomes low enough to affect the rate of CaCO3 dissolution. * kD = 150 %/day

261

Kinetics of CaCO, dissolution in marine sediment

l

/

w- 20-J 20

I 10

I I I 0 -10 -20 ACOi- (umobkg)

I

-30

FIG. IO. The ratio of the rate of CaCOz dissolution, attributable to the neutralization of metabolic acids, to the rate of C,, oxidation at each sampling site, plotted vs. ACO:. The closed circles are points calculated from optimized model parameters.The open circles (sites E and A) are calculated by adjusting the parameters describing the organic C oxidation rate (site E) or the CaCO, dissolution rate (site A) so that they are similar to the optimal values for the other sites. The dashed lines mark the trend predicted by the Corgoxidation/ CaC09 dissolution model if C,, oxidation and CaC4 dissolution kinetics do not vary greatly from site to site.

CaCO? dissolution (Fig. 11). The gradients (hence, the fluxes) are largest at the supersaturated site; in this case, the dominant source of the acid-neutralizing solutes is the supersaturated bottomwater. The gradients at the undersaturated site are smaller. The contribution of B (OH); to neutralization of metabolic acids at this site is insignificant; that of CO:- is still significant, and the dominant source of the CO:- is no longer bottomwater, but deeper porewaters, since the CO:concentration increases as saturation with respect to calcite is approached. Note that, at strong undersaturation (e.g., site G), the efficiency of dissolution by organic matter oxidation stops

C&O, Diss. rate 0.0 1.0 3.0x10d

CaCO, Diss. rate 0 20x10-6

70

co:-110x10-‘

increasing (Fig. 10) _ This occurs because, at these undersaturation levels, the supply of CaCO? to the sediments begins to limit the dissolution rate. Three out of seven of our study sites depart significantly from the shaded trend of Fig. 10. In each case, the deviation can be explained in terms of reaction kinetics. Sites E and H differ from the others in their depth distributions of organic matter oxidation. At site E, fits to O2 and NO.; profiles required that a larger proportion of the oxidation occur closer to the sediment-water interface (j, = 3 cm-‘, compared to 2 cm-’ at the other sites). At this site, the dissolution efficiency fell below the general trend: the closer organic matter oxidation is to the interface, the greater the flux of bottomwater solutes that can neutralize metabolic acids, and the less efficient Corg oxidation is at driving dissolution. At site H, the opposite was found: NO; data required Carg oxidation to occur deeper in the sediments. At this site, the dissolution efficiency fell well above the general trend. Site A differed from the others in that its WCS alkalinity profile required a more rapid rate constant for calcite dissolution (lSO%lday vs. 0.7-lO%lday, based on site H and used at all other sites); dissolution efficiency at this site also fell well above the general trend. We could verify that the departures from the trend at A and E are due to the hypothesized factors: by running the model using a fit to NO? data with j, = 2 for site E, and using kn = 0.7 for A. With these parameter modifications, the dissolution efficiencies for these sites fell within the general trend (the open symbols in Fig. 10). The existence of a general trend in the efficiency with which organic matter oxidation drives sedimentary calcite dissolution would be encouraging for the interpretation of the sedimentary record, since it may be predictable. However, three out of seven sites deviated significantly from the trend

45

45

B(OH)b

5.5 65~10~

B(OWh

55 65~10‘~

7.80

PH 7.90 8.00

7.80

PH 7.90 8.00

FIG. 11. Model profiles of the calcite dissolution rate and various porewater solutes vs. depth in the sedimentsat sites B (supersaturated)and G (undersaturated).The dissolution rate is in mol/cm~d/y;solute concentrationsare in molfL,,

W. R. Martin and F. L. Sayles

262

in ways that could be related to site-to-site variations in organic matter oxidation and calcite dissolution kinetics. This result indicates that a deeper understanding of the relationship between these processes is needed for accurate interpretations of sedimentary CaC03 accumulation rates. Acknowledgmenrs-Many people contributed to the work presented here. S. Smith was instrumental in the design and operation of the in situ whole-core squeezer. J. Goudreau carried out the bulk of the analyses whose results are presented. The work of the captain and crew of R/V Knorr was essential to our success, as was the assistance of C. Sweeney, K. Coluzzi, and D. Sigman with sample processing at sea. R. Jahnke generously allowed us to use his multi-corer for sample collection. D. McCorkle provided NO; results from shipboard sectioninglcentrifugation. The manuscript benefitted from reviews by D. McCorkle, R. Fran$ois, R. Jahnke, J. Milliman, and an anonymous reviewer. The work was funded by the National Science Foundation under OCE-9201344. This is Woods Hole Oceanographic Institution Contribution #8976. Editorial

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