Dissolved inorganic carbon and alkalinity fluxes from coastal marine sediments

28 Pages·2013·1.52 MB·English

Biogeosciences,10,371–398,2013
Biogeosciences
www.biogeosciences.net/10/371/2013/
doi:10.5194/bg-10-371-2013
©Author(s)2013. CCAttribution3.0License.
Dissolved inorganic carbon and alkalinity fluxes from coastal
marine sediments: model estimates for different shelf environments
and sensitivity to global change
V.Krumins1,*,M.Gehlen2,S.Arndt3,P.VanCappellen4,andP.Regnier1,5
1DepartmentofEarthSciences–Geochemistry,FacultyofGeosciences,UtrechtUniversity,3584CDUtrecht,
TheNetherlands
2LSCE/IPSL,LaboratoiredesSciencesduClimatetdel’Environnement,91191Gif-sur-YvetteCedex,France
3SchoolofGeographicalSciences,UniversityofBristol,BS81SSBristol,UK
4DepartmentofEarthandEnvironmentalSciences,UniversityofWaterloo,Waterloo,ONN2L3G1,Canada
5DepartmentofEarthandEnvironmentalSciences,CP160/02,Universite´ LibredeBruxelles,1050Brussels,Belgium
*currentaddress:DepartmentofEnvironmentalSciences,RutgersUniversity,NewBrunswick,NJ,USA
Correspondenceto:V.Krumins([email protected])
Received:30May2012–PublishedinBiogeosciencesDiscuss.:17July2012
Revised:14December2012–Accepted:17December2012–Published:24January2013
Abstract. We present a one-dimensional reactive transport sitionalfluxfallsinthehighrangeofglobalvaluesgivenin
modeltoestimatebenthicfluxesofdissolvedinorganiccar- the literature, the reported DIC and alkalinity fluxes should
bon (DIC) and alkalinity (A ) from coastal marine sedi- beviewedasupper-boundestimates.Increasingcoastalsea-
T
ments. The model incorporates the transport processes of water DIC to what might be expected in year 2100 due to
sediment accumulation, molecular diffusion, bioturbation the uptake of anthropogenic CO increases PIC dissolution
2
and bioirrigation, while the reactions included are the re- by 2.3Tmolyr−1and alkalinity efflux by 4.8Teq.yr−1. Our
dox pathways of organic carbon oxidation, re-oxidation of reactive transport modeling approach not only yields global
reduced nitrogen, iron and sulfur compounds, pore wa- estimates of benthic DIC, alkalinity and nutrient fluxes un-
ter acid-base equilibria, and dissolution of particulate inor- der variable scenarios of ocean productivity and chemistry,
ganiccarbon(calcite,aragonite,andMg-calcite).Thecoastal butalsoprovidesinsightsintotheunderlyingprocesses.
zone is divided into four environmental units with differ-
ent particulate inorganic carbon (PIC) and particulate or-
ganic carbon (POC) fluxes: reefs, banks and bays, carbon-
ate shelves and non-carbonate shelves. Model results are 1 Introduction
analyzed separately for each environment and then scaled
up to the whole coastal ocean. The model-derived estimate Based on globally averaged carbon box models driven
for the present-day global coastal benthic DIC efflux is by atmospheric CO2, globally averaged land-use changes
126Tmolyr−1, based on a global coastal reactive POC de- and climate, Andersson et al. (2005) recently estimated
positional flux of 117Tmolyr−1. The POC decomposition that the coastal ocean currently receives ∼80Tmol (inor-
leads to a carbonate dissolution from shallow marine sed- ganic and organic) Cyr−1 from terrestrial sources, pumps
iments of 7Tmolyr−1 (on the order of 0.1PgCyr−1). As- ∼20TmolCyr−1 from the atmosphere and sequesters on
suming complete re-oxidation of aqueous sulfide released the order of 40TmolCyr−1 in shallow marine sediments.
fromsediments,theeffectivenetfluxofalkalinitytothewa- Because the deposition flux of carbon could exceed its
ter column is 29Teq.yr−1, primarily from PIC dissolution burial flux by a factor of two or so (e.g., Andersson et al.,
(46%)andammonification(33%).BecauseourPOCdepo- 2005), the dissolved inorganic carbon (DIC) recycling flux
frombothbenthicorganicmatterdegradationandcarbonate
PublishedbyCopernicusPublicationsonbehalfoftheEuropeanGeosciencesUnion.
372 V.Kruminsetal.: Coastalsedimentalkalinityfluxes
Table1.GlobalCfluxestimatesincoastalsediments.
Globalestimate(TmolCyr−1) Reference
POCdepositionflux 15.6 Jahnke(2010)
40(pre-industrial) Anderssonetal.(2005)
52 Muller-Kargeretal.(2005)
97(0–150m,reactiveonly) Thullneretal.(2009)
117(reactiveonly) Thiswork
135 Dunneetal.(2007)
167 Wollast(1991)
183 SarmientoandGruber(2006)
POCburial 5 Muller-Kargeretal.(2005)
9 Anderssonetal.(2005)
10 Berner(1982)
14.2 SarmientoandGruber(2006)
15 Chen(2003)
17 Wollast(1991)
18–20 WollastandMackenzie(1989)
20 Mackenzieetal.(1998)
56 Dunneetal.(2007)
PICdeposition 23 MillimanandDroxler(1996)
23 Thiswork
24.5(notincludinginertriverinePIC) Anderssonetal.(2005)
PICburial 13 MillimanandDroxler(1996)
14.5(notincludinginertriverinePIC) Anderssonetal.(2005)
15 Chen(2003)
16 Thiswork
17 Feelyetal.(2004)
21 Mackenzieetal.(1998)
dissolution may be on the order of 40TmolCyr−1. These denitrification and sulfate reduction in sediments deposited
estimatesaresupportedtosomeextentbyobservationalev- onthecontinentalshelves.
idence (e.g., Mackenzie et al., 2005) and clearly highlight Because of the limited availability of comprehensive ob-
theprominentroleofshallowsedimentsintheglobalcoastal servationaldatasetsthatcanresolvetheheterogeneityofthe
ocean carbon budget. A number of recent studies have also global coastal ocean, it is nonetheless important to recog-
discussed the potential importance of alkalinity generation nize that the quantitative global significance of coastal sed-
in shallow sediments to the global ocean alkalinity budget, iments to the carbon cycle, DIC and alkalinity generation
suggesting that the benthic alkalinity source could act as fluxes remains poorly known. For instance, only a few esti-
negativefeedbacktorisingatmosphericCO conditionsand matesofPIC(particulateinorganiccarbon)depositionfluxes
2
oceanacidification(Chen,2002;Chenetal.,2003;Thomas in the coastal zone are available (23–24.5Tmolyr−1, Ta-
et al., 2009; Hu and Cai, 2011). This benthic source of al- ble 1), while global estimates of particulate organic carbon
kalinity to the overlying water column has mainly been at- (POC) deposition fluxes for seafloor depths between 0 and
tributed to anaerobic degradation pathways of deposited or- 200mvarybyanorderofmagnitude(16–183TmolCyr−1,
ganicmatter.Forinstance,Thomasetal.(2009)proposethat Table1).Thelargeuncertaintiesassociatedwiththedeposi-
the 0.073Teq.yr−1 of alkalinity generated in the Wadden tion fluxes and post-depositional fate of POC and PIC lead
Sea are produced by denitrification and sulfate reduction. to poorly constrained inorganic carbon budgets and ben-
This release could balance 20–25% of the CO uptake of thic fluxes for shallow marine sediments (see, for example,
2
the entire North Sea. By extrapolating budgets for the East Schneideretal.,2006).
China Sea, Chen (2002) estimates that continental margin Most commonly, benthic alkalinity fluxes are derived in-
sediments contribute 16–31Teq.yr−1 of alkalinity globally, directly by balancing water column budgets; that is, they
a value that is comparable to the global bicarbonate alka- are not based on an explicit representation of the benthic
linity delivered by rivers (Suchet et al., 2003). Hu and Cai compartment and the alkalinity-generating early diagenetic
(2011), however, obtained a much smaller benthic alkalin- processes. However, over the past three decades, early di-
ityfluxof4–5Teq.yr−1 basedonwhole-oceanestimatesof agenetic modeling has greatly improved our mechanistic
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V.Kruminsetal.: Coastalsedimentalkalinityfluxes 373
andquantitativeunderstandingofbiogeochemicalcyclingin is drawn from simple mass balance calculations, which do
marine sediments and should thus be considered a method notresolvethecomplexinterplaybetweenthelargenumber
of choice to quantify benthic alkalinity and DIC fluxes. A ofbenthicprocesses.
few studies have fully addressed the dynamic interplay be- Thepurposeofthiscontributionistodeterminethequan-
tween redox processes, driven by organic matter decompo- titative significance of major early diagenetic processes for
sition,andinorganiccarbonatedissolutionandprecipitation the production and consumption of DIC and alkalinity, as
(Boudreau and Canfield, 1993; Jourabchi et al., 2005; Bur- well as their benthic fluxes in coastal ocean sediments, for
digeetal.,2010).Mostearlydiageneticmodelingstudiesthat bothpresent-dayconditionsandinresponsetocoastalocean
explorebiogeochemicaldynamicsinshallowwaterenviron- eutrophicationandacidificationscenarios.Tothisend,aone-
mentstendneverthelesstobesite-specific,withlittleattempt dimensional (1-D) reactive transport model (RTM) of ben-
to extrapolate the results to larger scales. (e.g., Blair and thic organic carbon and inorganic carbon dynamics is pre-
Aller,1995;Martensetal.,1998;Eppingetal.,2002;Berget sented.Wesimulateorganiccarbonoxidationpathways,re-
al.,2003;LuffandMoll,2004;RojasandSilva,2005;Thull- oxidation of reduced nitrogen, iron and sulfur compounds,
neretal.,2005;MorseandEldridge,2007;Daleetal.,2008, pore water acid-base equilibria, and dissolution of particu-
2009; Anggara Kasih et al., 2008; Mogollon et al., 2009). lateinorganiccarbon(calcite,aragonite,andMg-calcite).In
Soetaert et al. (1996) developed an early diagenetic model addition,wedistinguishbetweenfourdifferentcarbonateen-
of sediments from the shelf (200m) to the abyss (3000m), vironmentsandanumberofdistinctenvironmentalsub-units
whichquantifiedcarbonoxidationviaoxicandanoxicpath- that are characterized by different biogeochemical drivers.
ways.TherecentstudybyThullneretal.(2009)alsoquanti- TheRTMisthenappliedtoquantifytheimpactofearlydia-
fiesorganiccarbonoxidationpathwaysalonganoceanhyp- geneticreactionsonthedissolutionofthreecarbonatephases
sometryfrom100to5000m,anddeterminestheassociated (calcite, aragonite, and a 15% Mg-calcite) and the benthic
dissolved inorganic carbon (DIC) fluxes. A number of cou- DIC and alkalinity fluxes for each of the identified carbon-
pled global biogeochemical models such as GENIE (Ridg- ate environments. We also provide a first-order estimate of
well,2007;RidgwellandHargreaves,2007),CLIMBERand the global alkalinity and DIC fluxes from marine sediments
UVic(Archeretal.,2009),andHAMMOC(Palastangaetal., totheshallowcoastalzonebyextrapolatingourregionalre-
2011)alsoapplysimplifiedmodelsofearlydiagenesis(e.g., sults to the scale of the global coastal ocean. Furthermore,
Muds (Archer et al., 2002) or the DCESS sediment module weassessthesensitivityofbenthicDICandalkalinityfluxes
of Shaffer et al., 2008). However, these global models can- ineachenvironmenttouncertaintiesassociatedwiththeben-
notresolvetheshallowoceandynamicsandtheirdiagenetic thicironandsulfurcycles.Finally,weexploretheresponse
componentsaretypicallybasedonsimplifyingassumptions of DIC and alkalinity fluxes to changes in organic and in-
suchasconsideringasinglePICphase(calcite)orneglecting organic carbon deposition fluxes due to eutrophication and
Fecycling/burial. oceanacidification,respectively.
Thelimitedquantitativeunderstandingofbenthicalkalin-
ity and DIC fluxes from shallow marine environments con-
trastswiththepotentialsignificanceofcoastalsedimentsfor 2 Methods
the global carbon cycle and climate. In particular, our abil-
ity to better constrain global coastal ocean C budgets will 2.1 Modelenvironments
strongly depend on a better quantification of biogeochemi-
caltransformationsandfluxesintheseshallowsediments.In The coastal ocean is one of the most complex and hetero-
addition,suchquantificationwillalsoimproveourabilityto geneousaquaticenvironmentsintheEarthsystem.Oneway
predict the response of the seafloor to perturbations of the to deal with this heterogeneity is to propose a classification
carbon cycle, including productivity changes and acidifica- ofthecoastaloceanintodifferentenvironmentswithbroadly
tionofcoastalwaters.Forinstance,Anderssonetal.(2005) similarcharacteristics,organizedaccordingtoseriesofhier-
predictthatincreasingcoastalproductivityoverthenextcen- archically related system levels (Table 2). Each hierarchical
turywillsignificantlyaltertheCO pumpingefficiencyofthe leveldescribesaleveloforganizationthatcontrolsDICand
2
globalcoastalocean.ThiseutrophicationwillincreasePOC alkalinityfluxesataspecificspatialscale.Processesathigher
depositionfluxes,and,becauseorganiccarbonmetabolismis levels have a direct influence on lower level characteristics,
the major driving force for PIC dissolution in shallow envi- andsmallerscalesarenestedwithinlargerscales.
ronments (Morse and Mackenzie, 1990), benthic carbonate At the highest hierarchical level (Level I), the coastal
dissolution will increase. In contrast, model simulations by ocean can be divided into distinct environments, each char-
Anderssonetal.(2005)indicatethatoceanacidification,and acterized by a distinct mode of carbonate production and
theresultingpHdropofcirca0.3unitspredictedbytheend accumulation (input control). Here, we follow the well-
ofthecentury(e.g.,CaldeiraandWickett,2003),shouldhave established classification proposed by Milliman and Drox-
a limited impact on benthic CaCO dissolution. Neverthe- ler(1996)intofourdifferenttypesofshallow-watercarbon-
3
less,thisconclusionrequiresfurtherinvestigation,becauseit ate environments: reefs, banks and bays, carbonate shelves,
www.biogeosciences.net/10/371/2013/ Biogeosciences,10,371–398,2013
374 V.Kruminsetal.: Coastalsedimentalkalinityfluxes
Table2.Hierarchicalclassificationscheme(seetextfordetails).
ControllingFactor Characteristics Scenarios
LevelI Modeofcarbonate JPIC 1 Reefs
Production/accumulation 2 Banks&embayments
3 Carbonateshelve
4 Non-carbonateshelves
LevelII Carbonatedissolution JPOC,O2,NO3,SO4 1 Globallyaveragedconditionsforreefs
2 Globallyaveragedconditionsfor
banks&embayments
3 Globallyaveragedconditionsfor
carbonateshelves
4 Globallyaveragedconditionsfornon-
carbonateshelves
LevelIII Localvariability waterdepthmastervariable 1 Averagewaterdepth25m
forvariabilityin 2 Averagewaterdepth25m
JPOC,JPIC,O2,NO3, 3 25,75,150m
SO4,w,α,Db 4 25,75,150m
and non-carbonate shelves. Reefs comprise coral reefs and ceptors O , NO and SO are additional control factors on
2 3 4
lagoons and show the highest carbonate production on an carbonatedissolution,sincetheydeterminetherelativecon-
areabasisinthepresent-dayocean.Accumulatedcarbonates tributionofaerobicandanaerobicdecompositionpathways.
mainlyoriginatefromcoralsandalgaeand,locally,alsofrom However,thespatialvariabilityinthemagnitudeandquality
foraminifera(e.g.,Milliman,1993).Banksandembayments oftheorganicmatterdepositionfluxremainsextremelydif-
encompasstropicalandsub-tropical,littoralandsub-littoral ficulttoassess,sincecoastalfluxesgenerallyrepresentady-
environmentsthatarecharacterizedbyhighratesofcarbon- namicmixtureoforganicmatterfromterrestrial,lateraland
ateproductionandaccumulation,sustainedbybenthicalgae, insitusources(e.g.,MollenhauerandEglinton,2007;Mol-
mollusks, benthic foraminifera and locally also bryozoans lenhauer et al., 2005). This lack of quantitative understand-
and serpulids (e.g., Stockman et al., 1967; Boscence et al., ing about the fate of coastal organic matter is also reflected
1985). In contrast to the reef and bank environments, the inthelargevariabilityintheglobalestimatesoforganicmat-
production and accumulation of carbonates on continental terdepositionfluxesforseafloordepthsbetween0and200m
shelveshavecertainlynotbeensurveyedtothesameextent. (16–183TmolCyr−1, Table 1) and in model-derived degra-
WeusethesamedivisionsasMillimanandDroxler(1996)in dationrateconstants(Arndtetal.,2011).
dividingthecontinentalshelfaccordingtoitscarbonatecon- Therefore, further regionalization based on organic mat-
tent into carbonate and non-carbonate shelves . In these en- ter and redox criteria is not achievable at this stage
vironments, carbonate production is generally supported by and we choose to extract average environmental condi-
benthicandpelagicproduction,aswellasrelictcomponents, tions for each of the four environmental units selected
such as oysters, ooids and dead coral reefs (e.g., Milliman, at Level I only. Values are either extracted from the lit-
1993andreferencestherein). erature or derived from the 1◦ dataset of the Levitus94
Atthenexthierarchicallevel(LevelII),theenvironmental database(NOAAWorldOceanAtlas1994,availableonline
units identified above should be further subdivided into re- at http://iridl.ldeo.columbia.edu/). This simplified approach
gionsthatarecharacterizedbysimilarcontrolsoncarbonate canbeseenasafirststeptowardsabetterregionalclassifica-
dissolution. Previous research has shown that organic mat- tionofthecoastaloceanintermsofthedominantcontrolson
terdegradationexertsadominantcontrolonbenthiccarbon- carbonatechemistry.Italsoimpliesthatourresultsonlypro-
ate dissolution through the production of metabolic carbon videafirst-orderglobalestimateofDICandalkalinityfluxes
dioxide, the associated lowering of the pore water pH and, forthecoastalocean.
thus, the saturation state (e.g., Jahnke et al., 1994). In this Onalocalscale,depositionalfluxes,bottomwaterconcen-
context,themagnitudeandqualityoftheorganicmatterde- trations and transport parameters often show a strong cor-
position flux is a fundamental control factor of the carbon- relation with water depth (e.g., Soetaert et al., 2002; Ep-
ate dissolution fluxes. In addition, anoxic degradation path- ping et al., 2002). Pressure, and thus water depth, also di-
ways generally produce alkalinity and, therefore, favor car- rectly affects the solubility products of the various carbon-
bonate preservation or even precipitation. Thus, the bottom atephases.WaterdepthisthususedatLevelIIIofourclas-
water concentrations of the dominant terminal electron ac- sification, and we assume an average value of 25m for the
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V.Kruminsetal.: Coastalsedimentalkalinityfluxes 375
reefs,andthebanksandbays.Theshelvesarefurthersubdi- transientsimulationsthatarerununtilsteadystate(generally
vided into three depth ranges: 0–50m, 50–100m, and 100– 200–400simulationyearsarerequiredfortheconcentration
200m.Thesedepthintervalscontribute41,31,and28%of profiles to adjust to their final configurations from arbitrary
the global coastal seafloor area shallower than 200m, re- initial profiles. This is approximately the time required to
spectively, based on the data of Smith and Sandwell (1997) advect the solid phase through the modeling domain.). The
extracted from a ∼1min regridding from 72◦N to 72◦S, modeled sediment thickness is 50cm, discretized in 116 ir-
2
accessed online at http://www.grdl.noaa.gov/cgi-bin/bathy/ regularlyspacednodes.Nodespacingvariesfrom0.1cmat
bathD.pl. thesediment–waterinterfaceto2cmatthebaseofthemod-
eledsedimentcolumn.
2.2 Modelapproach
2.3 Transport
A 1-D reactive transport model for coastal sediments is
developed using the Biogeochemical Reaction Network Thechoiceoftransportparametersisguidedbyempirically
Simulator (Regnier et al., 2002; Aguilera et al., 2005), determined global values or global relationships proposed
which has previously been used to simulate mixed kinetic- in the literature. The sediment advection (burial) rate (ω,
equilibrium reaction networks describing early diagenesis cmyr−1 at the sediment–water interface (SWI)) and biotur-
(e.g.,Jourabchietal.,2005;Thullneretal.,2005).Thever- bation coefficient (D ) are obtained from the global depth
b0
ticalconcentrationfieldsin thesedimentaremodeledusing relationshipsproposedbyMiddelburgetal.(1997):
the1-Dadvection-diffusion-reactionequations:
ω=3.3×10(−0.87478367−0.00043512·SFD) (3)
ϕ∂∂Cti =−∂(∂ωxϕCi)+∂∂x(cid:26)(Db+Di)∂(∂ϕxCi)(cid:27)+αi(Ci0−Ci)+XRi(1)
and
D =5.2×10(0.76241122−0.00039724·SFD), (4)
b0
(1−ϕ)∂∂Cti =−∂(ω(1∂x−ϕ)Ci)+∂∂x(cid:26)Db∂((1−∂xϕ)Ci)(cid:27)+XRi(2) whereSFDistheseafloordepthinmeters.
A global average bioturbation depth of 10cm in marine
for solutes and solid-bound species, respectively. In the
sediments(Boudreau,1994)isassumedforallenvironments
aboveequations,x isdepthbelowthesediment–waterinter-
and water depths. A complementary error function is used
face(cm),and
to attenuate the bioturbation coefficient with depth. Thus at
ϕ=porosity[](assumedtobeconstantintime), 10cm,D is50%ofthevalueattheSWI.At15cm,itis4%
b
ofthevalueatthesurface.Anexponentialdecayfunctionis
ω=sedimentburialvelocity[cmyr−1],
imposed for the porosity (ϕ) depth profile (Wang and Van
C = concentration of species i at depth x and time t Cappellen,1996):
i
[µmolcm−3],
ϕ(x)=0.75+0.1e−0.15x. (5)
C =concentration of i at sediment–water interface
i0
[µmolcm−3], Diffusion coefficients for the dissolved species are cor-
rected for temperature and sediment porosity according to
D =bioturbationcoefficient[cm2yr−1], Boudreau (1996). Bioirrigation intensity is assumed to de-
b
creaseexponentiallywithdepthinthesediment:
D =moleculardiffusioncoefficientofsolutespeciesi
i
[cm2yr−1], α(x)=α ·e(−x/xirr) (6)
0
α =non-localbioirrigationcoefficient[yr−1],and
i where α is the bioirrigation coefficient at the sediment–
0
6R =net rate of all reactions involving species i water interface (yr−1), and xirr is a depth attenuation coef-
i
[µmolcm−3yr−1]. ficient (cm). The value of xirr is set to 3.5cm, while α0 is
estimatedfollowingThullneretal.(2009).
AdiscretizedversionofEq.(1)issolvedwithanoperator Someofthedissolvedoxygenthatistransportedintothe
splitting approach for the transport and reaction terms. So- sediment via bioirrigation will oxidize reduced iron in the
lutetransportisdiscretizedusingtheCrank–Nicholsonalgo- vicinity of burrows. The iron (oxy)hydroxide precipitates
rithm,whileforsolidsandsorbedspecies(Eq.2)afluxlim- sorbaqueousphosphate(e.g.,Spiterietal.,2008),hencere-
iting total variation diminishing scheme is used (e.g., Reg- ducingtheapparentirrigationaltransportofphosphate.Berg
nieretal.,1998).Thereactiontermiscastasasetofdiffer- et al. (2003) and Meile et al. (2005) showed that effective
ential algebraic equations (DAEs) and solved using a New- bioirrigationratesdifferwidelyamongchemicalspeciesde-
ton algorithm (for further details, see Aguilera et al., 2005; pendingontheirbehavioratburrowwalls.Forinstance,ac-
Centleretal.,2010).A5×10−6yrtimestepisimposedfor cording to Berg et al. (2003), α ≈0 for Fe2+. Rather than
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376 V.Kruminsetal.: Coastalsedimentalkalinityfluxes
Table3.Kineticandequilibriumreactions,andgenerationofDICandAT.
Reaction DIC AT
Hydrolysis
(R1,R2) 1106(CH2O)106(NH3)16(H3PO4)+11056H+→CH2O+11066NH+4 +1106H2PO−4 0 +0.14
Primaryredoxreactions
(R3) CH2O+O2→CO2+H2O +1 0
(R4) CH2O+0.8NO−3 +0.8H+→CO2+0.4N2+1.4H2O +1 +0.8
(R5) CH2O+4Fe(OH)3+8H+→CO2+4Fe2++11H2O +1 +8
(R6) CH2O+12SO24−+12H+→CO2+12HS−+H2O +1 +1
Secondaryredoxreactions
(R7) 12NH+4 +O2→12NO−3 +12H2O+H+ 0 −1
(R8) 4Fe2++O2→4Fe(OH)3+8H+ 0 −8
(R9) 12HS−+O2→12SO−4 +12H+ 0 −1
(R10) 21FeS+O2→12Fe2++12SO42− 0 0
PICdissolution
(R11) CaCO3(calcite)→Ca2++CO23− +1 +2
(R12) CaCO3(aragonite)→Ca2++CO23− +1 +2
(R13) Ca0.85Mg0.15CO3→0.85Ca2++0.15Mg2++CO23− +1 +2
Otherkineticreactions
(R14) Fe2++HS−→FeS+H+ 0 −2
(R15) FeS+H2S→H2+FeS2 0 0
(R16) Fe2++CO23−→FeCO3 −1 −2
(R17) H2S+inert-POC→organic-S 0 0
Equilibria
H2CO3∗↔H++HCO−3
HCO−↔H++CO2−
3 3
H2O↔H++OH−
H2S↔H++HS−
B(OH)3+H2O↔H++B(OH)−4
Fe(OH)3+H2PO−4 ↔iron-sorbed-P
H2PO−4 ↔sorbed-P
Fe2+↔sorbed-Fe
NH+4 ↔sorbed-NH4
ReactionnumbersforreferencetoTables2and3.
DIC–molofDICgeneratedorconsumedforthereactionaswritten.
AT–eq.ofalkalinitygeneratedorconsumedforthereactionaswritten.
EquilibriumreactionsdonotleadtoanynetgenerationofDICoralkalinity.
modifyingα (inEq.1),Katsevetal.(2007)reducedthelo- 2.4 Reactions
calconcentrationsofFe2+ andphosphatebyafactorof0.1
toaccountforreducedbio-irrigation.Here,wechosetomul- 2.4.1 POChydrolysis(R1–2)
tiplyα calculatedaccordingtoEq.(6)by0.01forFe2+ and
−
H PO .Thiscorrespondsroughlytotheratioofthedepth-
2 4 Organic matter in the sediments is assumed to have Red-
integrated α value for Fe2+ to that of solutes such as O2 or field stoichiometry: (CH2O)106(NH3)16(H3PO4) (Table 3,
+
NH inMeileetal.(2005). R1 and R2). Organic matter degradation is modeled as a
4
two-step process. It is first hydrolyzed to yield dissolved
organic carbon (DOC), ammonia and phosphate. Hydroly-
sis rates are assumed to exhibit a first-order dependence on
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V.Kruminsetal.: Coastalsedimentalkalinityfluxes 377
the concentrations of the reactive POC pools. Ammonia re- Table4.Kineticratelaws.
leased during POC hydrolysis is protonated to ammonium
(NH+),whilephosphateisreleasedasdihydrogenphosphate Reaction
4
−
(H2PO4),thepredominantformundertypicalporewaterpH R1=kPOC1·[POC1]
conditions. The hydrolysis products (represented as DOC) R2=kPOC2·[POC2]
aoornfeePtOorefCatthetedopathsriemresaaercdytiimvreeedniontxteisrpmaatsehsdwuimaatyeessd,(twSoehccitco.hn2sa.i3sret.2oo).fxiTtdwhizoeeidrnepvaucia-t RR43==kk33··(cid:16)(cid:16)YkYONO2O[2O+32[(cid:17)O]·2(cid:18)](cid:17)kN·[OC[3N+HO[23NO]O]3](cid:19)·[CH2O]·(cid:18)kiNkOi3N+O[3O2](cid:19)
tgahilayvlndyeir)cpoiolcsyoazarlesblldooancnwadfltieutohdxnae(to6rr7etah%ftereaochcftoiogtnrhhsyeltyaptnoortteoaaollc:fot8ir1vg0eya%rnP−iOc1o,cfCawtrhpbheooiolnrele,flatwhcutexhivigocelthohobeirsr-- RRR567===kkk33N··it(cid:16)(cid:16)riYYfYYOS·FOOe2[24N(cid:17)(cid:17)H··(cid:16)+4(cid:16)k]kF·eS[[+OFO[4e[S(F+2OOe][(H4SO]O)H34])]3(cid:17)](cid:17)·[·C[CHH2O2O]·](cid:16)·(cid:16)kikSiOkFi4ekS++iOF[4[eOO22]](cid:17)(cid:17)
20% of the reactive organic carbon flux (17% of the total R8=kFeOx·[O2]·[Fe2+]/[H+]2
organiccarbonflux)ismoreslowlyhydrolyzed,withk=0.1 R9=kSulfOx·[O2]·([HS−]+[H2S])
yr−1. The reactivity, k1, of the most reactive POC fraction R10=kFeSOx·[FeS]·[O2]
is representative for phytoplankton-derived POC and falls R11=kCal(1−(cid:127)Cal)nCal (cid:127)Cal<1
within the typical range of 1–10yr−1 reported in the liter- R12=kAra(1−(cid:127)Ara)nAra (cid:127)Ara<1
ature(e.g.,WestrichandBerner,1984;HenrichsandDoyle, R13=kMg−cal(1−(cid:127)Mg−cal)nMg−cal (cid:127)nMg−cal<1
1is9u8s6u;aBlluyrldeisgse,re1a9c9ti1v)e.,Aangdedtyopricrea-lsvuaslpueensdreedpoorrtgeadniinctmheatltietr- R14=kFeS·(cid:16)[F[He2++]]·K·[HsFSe−S]−1(cid:17) R14≥0
eraturerangebetween0.1–1yr−1(e.g.,WestrichandBerner, R15=kPyr·[Fe(cid:18)S][F·e[H2+2]·S[C]O2−] (cid:19)
1984).ThethirdPOCpoolofrefractoryorganicmatterrep- R16=kFeCO3· KsFeCO33 −1 R16≥0
resents 16% of the total carbon pool and is assumed to be R17=kOrgS·[inert-POC]·[H2S]
unreactive,k=0.0yr−1onthetimescalesconsideredhere.It
is not explicitly modeled, but it is included in sulfidization
reactionsandinCbudgetestimates. 2.4.3 PICdissolution(R11–13)
2.4.2 Primaryredoxreactions(R3–6) TheratesofdissolutionofthePICphasesarecomputedusing
The primary redox reactions of aerobic oxidation, denitrifi-
cation, iron reduction, and sulfate reduction are coupled to R =k·(1−(cid:127))n·C (7)
dissolution
DOC oxidation (Table 3). For the depositional conditions
considered,sulfateisneverentirelyexhaustedwhenreaching where C is the concentration of the PIC phase and (cid:127), the
the bottom of the early diagenetic zone (x=50cm); hence, saturationstate,isgivenby
methanogenesisisassumedtobeofminorimportanceinor- (cid:127)=[Ca2+]·[CO2−]/K ∗forcalciteandaragonite,and
3 SP
ganic matter degradation. This is consistent with observa- (cid:127)=[Ca2+]0.85·[Mg2+]0.15·[CO2−]/K ∗ for 15% Mg-
3 SP
tions and model results for shelf sediments receiving high calcite.
organicmatterinputswhichshowthatthedepthscaleforthe The stoichiometric solubility products (K ∗) of calcite
SP
onset of methanogenesis is typically on the order of a cou- and aragonite are calculated as a function of temperature,
pleofmetersormore(e.g.,Daleetal.,2008;Regnieretal., pressure and salinity according to Millero (1995), and the
2011).Theratesoftheprimaryredoxreactionsareassumed apparent solubility product of 15% Mg-calcite in seawa-
to exhibit a first-order dependence on the concentration of ter is assumed to be 21% greater than that of aragonite,
DOCandaMichaelis–Menten-typedependenceonthecon- based on the findings of Morse et al. (2006). The kinetic
centrationoftheterminalelectronacceptor(Table4).Inad- parameters k and n for calcite, aragonite and 15% Mg-
dition,theratesofthedifferentelectronacceptingpathways calcite are obtained from Walter and Morse (1985) using
are scaled by their relative cell yields (Table 4). Yield val- Balanus, Halimeda, and Peneroplis data, respectively. The
uesandhalf-saturationconstantsaretakenfromThullneret experimentally-determinedrateconstantskaremultipliedby
al. (2005) (Table 5). The anaerobic processes are also non- a factor 0.1 to account for the effects of inhibitors, surface
competitively inhibited by oxygen. The inhibition constant poisoningandcoatingsthatarelikelytoreducethedissolu-
for denitrification is set to result in 90% inhibition when tionkineticsofPICinnaturalsediments.
[O2]=63µM(2mgL−1),whilethoseforironandsulfatere- Pore water calcium concentrations are explicitly com-
ductionleadto90%inhibitionat[O2]=6µM(0.2mgL−1). puted, while the concentration of pore water magnesium is
TheoverallDOCoxidationrateconstantk3iscommontoall expected to vary little, and is kept constant at the average
primaryredoxpathways,aerobicandanaerobic. seawater value of 53mM. Although pore waters may be
oversaturatedwithrespecttooneormoreCaCO phases,in
3
www.biogeosciences.net/10/371/2013/ Biogeosciences,10,371–398,2013
378 V.Kruminsetal.: Coastalsedimentalkalinityfluxes
Table5.Selectedparametersforkineticequations. Soluble ferrous iron, ammonium, and phosphate are as-
sumed to adsorb non-specifically to solids. Specific phos-
Parameter Value Units Note phate sorption onto ferric iron phases is also modeled. The
+
sorbed-P, sorbed-NH , sorbed-Fe(II) and iron(III)-sorbed-P
kkkPPPOOOCCC123 0.101 yyyrrr−−−111 athree tpreoaretedwaastesrolpihda-bs4eo.unFdincaollnys,tirtuefernatcstoinryeqourgilaibnriciummawttietrh,
whichisnototherwisemodeled,isassumedtoreactwithsul-
kO2 20 µM a fideduringtheprocessofsulfurization(R17).
kNO3 5 µM a
kFe 1 µmol/g a
2.5 Boundaryconditions
kSO4 1600 µM a
YNO3/YO2 0.75 – a 2.5.1 Baselinescenario
YFe/YO2 0.25 – a
YSO4/YO2 0.2 – a The published estimates of global carbon (POC and PIC)
kkiiFNeO3 0.77 µµMMooffOO22 bb deposition and burial fluxes are summarized in Table 1. In
kiSO4 0.7 µMofO2 b theabsenceofarobustregionalizationofPOCfluxesforthe
kNitrif 10000 mM−1yr−1 c globalcoastalocean,thedegradablePOCfluxesforthecar-
kFeOx 3.6·10−6 mMyr−1 d bonate shelves, non-carbonate shelves, and banks and bays
kSulfOx 320–320000 mM−1yr−1 e are estimated using the depth relationship for the benthic
kCal 40 mM−1yr−1 f oxygen consumption of coastal sediments by Soetaert et
nCal 2.74 – f al.(2002):
kAra 110 mM−1yr−1 f
nAra 2.43 – f JPOC=40.61·SFD−0.571 (8)
kMg-cal 58 mM−1yr−1 f where J is the degradable POC flux in molCm−2yr−1.
nMg-cal 3.61 – f POC
KkFseFSeS 0.060.53 µµmmoollccmm−−33yr−1 cc Tanhdus1,5t0hemresahcetlifvdeePpOthCsadreep6o.s5i,ti3o.n5,flaunxdes2f.3ormtohleC25mm−,27y5r−m1,,
kPyr 60 cm3µmol−1yr−1 g respectively(Table6a)Forthebanksandbaysenvironment,
kOrgS 0.0002 cm3µmol−1yr−1 g thereactivePOCfluxcalculatedfromEq.(8)assuming25m
SFD (=6.5molCm−2yr−1) is also used. The imposed re-
aThullneretal.(2005). active POC deposition flux for reefs, 5.2molCm−2yr−1,
bWangandVanCappellen(1996).
cJourabchietal.(2005). corresponds to the estimated average net primary produc-
dAdjustedforT andseawaterfromStummandMorgan(1996). tivity of coral reef ecosystems (Hatcher, 1990). When in-
eSeeSect.2.7.1.
fRecalculatedfromWalterandMorse(1985). tegrated over the total surface area of the coastal ocean
gDaleetal.(2009). (SFD≤200m), the resulting global reactive POC flux esti-
mateis117.5Tmolyr−1.Thisvaluefallswithintherangeof
publishedvaluesforglobalcoastalPOCdepositionflux(16
situprecipitationofnon-biogenicCaCO3insedimentsisbe- to 183Tmolyr−1), but is somewhat higher than the average
lievedtobeoflimitedimportance,relativetodissolution,due value(101Tmolyr−1)ofallthestudiesreportedinTable1.
toinhibitionbyorganiccompoundsandmagnesium(Morse In addition, estimates of the average global POC burial in
andMackenzie,1990;MorseandMucci,1984).Asaresult, coastalsediments(Table1)indicatethatglobalcoastalsedi-
carbonate precipitation is not included in the reaction net- mentsreceiveanadditional18.4Tmolyr−1ofPOCthatisre-
work. fractoryontheconsideredtimescale.Thetotalglobalcoastal
POCdepositionfluxthusamountsto136Tmolyr−1.Ama-
2.4.4 Otherreactions jorsourceofuncertaintyfortheDICandalkalinityfluxesis
theverylimitednumberofstudiesthatprovideestimatesof
−
DissociationsofH O,H BO ,H S,H CO andHCO are global coastal PIC deposition fluxes (Table 1). The PIC de-
2 3 3 2 2 3 3
modeledasequilibriumreactions.Theapparentequilibrium positionfluxesaretakenfromMillimanandDroxler(1996),
constantsareadjustedforsalinity,temperatureandpressure andrangefrom0.27molm−2yr−1fornon-carbonateshelves
accordingtoMillero(1995).Secondaryredoxreactions(i.e., to 15molm−2yr−1 for coral reefs. Published estimates for
reoxidationreactionsofreducedby-productsoforganicmat- thearagonitefractionofthecarbonateexportfluxrangefrom
terdegradation)arealsoconsidered.Thereactionsincluded 10% to 50%, and may be as high as 75% (Munhoven,
aretheoxidationofammonia,hydrogensulfide,ironsulfides 2007).Ontheshelves,wedividethePICdepositionfluxinto
and ferrous iron by oxygen (R7–10). Iron sulfide precipita- 35% aragonite and 65% calcite, based on the estimates of
tionanditsconversiontopyriteareincludedinthereaction pelagicPICproductionbyGangstøetal.(2008).Forthereefs
network,asisironcarbonateprecipitation(R14–16). and banks and bays, PIC is divided into aragonite, calcite
Biogeosciences,10,371–398,2013 www.biogeosciences.net/10/371/2013/
V.Kruminsetal.: Coastalsedimentalkalinityfluxes 379
Table6a.Boundaryconditions,baselinecase.
Environment
Parameter Carbonateshelf Non-carbonateshelf Banksandbays Reefs
Fluxes(molm−2yr−1)
POC 25m 6.46 6.46 6.46 5.20
75m 3.45 3.45
150m 2.32 2.32
Fe(OH)3 25m 0.038 0.038 0.038 0.038
75m 0.036 0.036
150m 0.033 0.033
Calcite 0.39 0.17 0.65 1.95
Aragonite 0.21 0.09 3.15 9.45
15%Mg-calcite 0 0 1.20 3.60
Concentrations(mM)
O2 0.195 0.195 0.195 0.195
NO−a 25m 0.005 0.005 0.005 0.005
3
75m 0.008 0.008
150m 0.012 0.012
SO2− 28 28 28 28
4
−
H2PO4 0.001 0.001 0.001 0.001
Ca2+ 10 10 10 10
AT 2.306 2.306 2.306 2.306
DICa 25m 2.022 2.022 2.022 2.022
75m 2.068 2.068
150m 2.138 2.138
CO2− 25m 0.201 0.201 0.201 0.215
3
75m 0.170 0.170
150m 0.124 0.124
◦
Temperature( C) 25m 14.34 14.34 9.91 27.6
75m 11.63 11.63
150m 7.57 7.57
pH(freescale) 25m 8.27 8.27 8.27 7.95
75m 8.19 8.19
150m 8.04 8.04
(cid:127)b 25m 2.98 2.98 2.98 3.19
Aragonite
75m 2.55 2.55
150m 2.01 2.01
aNitrateandDICconcentrationsdeterminedfromSFDcorrelationsfoundinGLODAPdatabase:
[NO3](mM)=0.003734+0.0000584·SFD (r2=0.250,p<10−195)
[DIC](mM)=1.9992+0.0009220·SFD (r2=0.262,p<10−5)
b(cid:127)AragonitecalculatedfromMillero(1995)usingpressure,salinity,andthefollowingtemperaturecorrelationextractedfromthe
Levitus94database:
Temp(◦C)=15.69−0.054·SFD (r2=0.098,p<10−42)
and 15% Mg-calcite in the ratios 63:13:24 (Land, 1967). sults in a nearly depth-independent relationship for the PIC
The PIC flux in a given coastal environment is assumed to flux (=0.63·SFD−0.011molm−2yr−1), which is very close
be independent of SFD (Table 6a). The Koeve (2002) rela- to Milliman and Droxler’s (1996) estimate for PIC deposi-
tionship of the POC:PIC ratio with depth in the Atlantic tiononcarbonateshelves(0.60molm−2yr−1).Notethatour
Basin (POC:PIC=64.3·SFD−0.56), when combined with globally integrated PIC deposition flux for the coastal zone
the Soetaert POC deposition flux relationship (Eq. 8), re-
www.biogeosciences.net/10/371/2013/ Biogeosciences,10,371–398,2013
380 V.Kruminsetal.: Coastalsedimentalkalinityfluxes
Table6b.BoundarypHand(cid:127)AragoniteforDICsensitivitytests.
Environment
Parameter Carbonateshelf Non-carbonateshelf Banksandbays Reefs
+2.5%DIC
pH 25m 8.18 8.18 8.18 7.86
75m 8.08 8.08
150m 7.90 7.90
(cid:127)Aragonite 25m 2.55 2.55 2.55 2.85
75m 2.10 2.10
150m 1.47 1.47
+5%DIC
pH 25m 8.07 8.07 8.07 7.76
75m 7.95 7.95
150m 7.74 7.74
(cid:127)Aragonitea 25m 2.07 2.07 2.07 2.34
75m 1.63 1.63
150m 1.05 1.05
a(cid:127)AragonitecalculatedfromMillero(1995)usingpressure,salinity,andthefollowingtemperaturecorrelationextractedfrom
theLevitus94database:
Temp(◦C)=15.69−0.054·SFD (r2=0.098,p<10−42)
is also in agreement with the value used by Andersson et indicates that dissolved inorganic carbon (DIC) also varies
al.(2005;Table1). with SFD. In contrast, there is no significant correlation
Raiswell(2006)estimatedtheglobaldeliveryofhighlyre- betweensalinityoralkalinityandseafloordepth(p>0.05),
active iron to the shelves to be 53Tgyr−1, primarily from so the average values of 34.64 (psu) and 2.306meqL−1 are
rivers, atmospheric dust, and diagenetic recycling. Dividing used. The bottom water alkalinity and DIC concentrations
by the shelf area, this gives an average deposition flux of are used to calculate pH and partitioning of the aqueous
0.036molFem−2yr−1, which we scale with the sediment carbonatespecies(H CO ∗/HCO−/CO2−),usingtheappar-
2 3 3 3
burialrateinordertoestimatethereactiveirondepositionin ent equilibrium dissociation constants adjusted for depth,
thevariousenvironmentsandwaterdepthranges(Table6a). salinity and temperature (Millero, 1995). Table 6a presents
Bottomwatertemperature,oxygenconcentration,andni- the alkalinity, DIC, pH, [CO2−], and aragonite saturation
3
trate concentration are extracted from the 1◦ dataset of the imposedasbottomwaterboundaryconditions.
Levitus94 database. For each latitude and longitude in the Thetotaldissolvedboronconcentrationisestimatedfrom
dataset,thedeepestwaterdepthwithdataavailableistaken salinity according to Millero (1995). The concentrations of
as representative of bottom water, and the corresponding SO2−, Ca2+ and Mg2+ at the sediment–water interface are
4
(temperature,oxygenornitrate)valuesareinferredtobethe set to 28, 10, and 53mM, respectively, while bottom water
−
bottom water values. Over the 0–200m SFD range, no sig- [H PO ]isassumedtobe0.001mM.Theconcentrationsof
2 4
nificant relationship exists between bottom water [O2] and H2S,HS−,NH+4,Fe2+attheseafloorareassumedtobezero.
water depth (p>0.05); therefore, the average oxygen con-
centration (0.195mM) is used. Nitrate and temperature are 2.6 BenthicDICandalkalinityflux
bothwater-depth–dependent(Table6a).Theaveragebottom
water temperature is used along with pressure to calculate Because the DIC buried with pore waters is very small
K∗ valuesforaragoniteandcalcite(Millero,1995)andthe (<1%),andinsituCaCO3 precipitationandFeCO3 precip-
SP
apparent equilibrium constants for the dissociation of weak itation are predicted to be minimal, the benthic DIC efflux
acids.Theglobalaveragetemperatureforcoralreefs,27.6◦C is essentially equal to the sum of the depth-integrated rates
(Kleypasetal.,2005),isusedasbottomwatervalueforthe ofPOCoxidationandPICdissolution.Forthealkalinityflux
reefs. thesituationismorecomplicatedandcanbeanalyzedasfol-
Analysis of the GLODAP database for sta- lows.
tions with seafloor depths ≤200m (accessed from
http://cdiac.ornl.gov/ftp/oceans/GLODAP bottle files/)
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