constraining magnesium cycling in marine sediments using ...apaytan/290a_winter2014/pdfs... ·...

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Geochimica et Cosmochimica Acta 74 (2010) 5039–5053

Constraining magnesium cycling in marine sedimentsusing magnesium isotopes

J.A. Higgins *, D.P. Schrag

Department of Earth and Planetary Science, Harvard University, 20 Oxford St., Cambridge, MA 02138, USA

Received 22 April 2009; accepted in revised form 14 May 2010; available online 25 May 2010

Abstract

Magnesium concentrations in deep-sea sediment pore-fluids typically decrease down core due to net precipitation of dolo-mite or clay minerals in the sediments or underlying crust. To better characterize and differentiate these processes, we havemeasured magnesium isotopes in pore-fluids and sediment samples from Ocean Drilling Program sites (1082, 1086, 1012,984, 1219, and 925) that span a range of oceanographic settings. At all sites, magnesium concentrations decrease with depth.At sites where diagenetic reactions are dominated by the respiration of organic carbon, pore-fluid d26Mg values increase withdepth by as much as 2&. Because carbonates preferentially incorporate 24Mg (low d26Mg), the increase in pore-fluid d26Mgvalues at these sites is consistent with the removal of magnesium in Mg-carbonate (dolomite). In contrast, at sites where therespiration of organic carbon is not important and/or weatherable minerals are abundant, pore-fluid d26Mg values decreasewith depth by up to 2&. The decline in pore-fluid d26Mg at these sites is consistent with a magnesium sink that is isotopicallyenriched relative to the pore-fluid. The identity of this enriched magnesium sink is likely clay minerals. Using a simple 1Ddiffusion–advection–reaction model of pore-fluid magnesium, we estimate rates of net magnesium uptake/removal and asso-ciated net magnesium isotope fractionation factors for sources and sinks at all sites. Independent estimates of magnesium iso-tope fractionation during dolomite precipitation from measured d26Mg values of dolomite samples from sites 1082 and 1012are very similar to modeled net fractionation factors at these sites, suggesting that local exchange of magnesium between sed-iment and pore-fluid at these sites can be neglected. Our results indicate that the magnesium incorporated in dolomite is 2.0–2.7& depleted in d26Mg relative to the precipitating fluid. Assuming local exchange of magnesium is minor at the rest of thestudied sites, our results suggest that magnesium incorporated into clay minerals is enriched in d26Mg by 0& to +1.25&

relative to the precipitating fluid. This work demonstrates the utility of magnesium isotopes as a tracer for magnesiumsources/sinks in low-temperature aqueous systems.� 2010 Elsevier Ltd. All rights reserved.

1. INTRODUCTION

The chemical composition of pore-fluids in marine sedi-ments records interactions between seawater, sediments,and the oceanic crust. Because the concentrations of mostelements in the pore-fluid are much lower than their con-centrations in the surrounding sediment, reactions in thesediment column and/or underlying basalt often lead tolarge changes in pore-fluid chemistry but will have little

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doi:10.1016/j.gca.2010.05.019

* Corresponding author. Tel.: +1 617 852 9931.E-mail address: [email protected] (J.A. Higgins).

impact on the chemical composition of the solid phases.As a result, the elemental and isotopic composition of sed-imentary pore-fluids can be used to study both the distribu-tion and identity of precipitating/dissolving minerals in thesediment column and underlying basalt, even when thereare no measurable changes in bulk sediment composition.Extensive sampling of sedimentary pore-fluids by the DeepSea Drilling Project (DSDP) and Ocean Drilling Program(ODP) has revealed large gradients in major seawater cat-ions (Ca2+, Mg2+) in pore-fluids associated with a rangeof sediment types (carbonate and siliciclastic) and accumu-lation rates (Gieskes, 1975; Sayles and Manheim, 1975;Baker, 1985). Although there is some variability, in general,

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5040 J.A. Higgins, D.P. Schrag / Geochimica et Cosmochimica Acta 74 (2010) 5039–5053

the concentration of magnesium in pore-fluid decreaseswith depth, whereas the concentration of calcium increases.At many sites, magnesium and calcium profiles are approx-imately linear and associated with decreasing pore-fluidd18O values, consistent with low-temperature alteration ofthe underlying oceanic crust (e.g., Lawrence et al., 1975).However, more complicated profiles are also frequently ob-served, which has been interpreted as representing signifi-cant precipitation and dissolution of carbonate andsilicate minerals within the sediment column (Gieskes andLawrence, 1981; Baker, 1985; Moore et al., 2004), changesin the chemical composition of seawater over time (Fantleand DePaolo, 2006) and/or significant pore-fluid advection(McDuff and Gieskes, 1976).

Measurements of pore-fluid magnesium isotope ratios(25Mg/24Mg and 26Mg/24Mg) are a potentially useful toolfor unraveling the processes responsible for magnesium gra-dients in deep-sea sediments. Recent work on the variabilityof magnesium isotope ratios in carbonate and silicate min-erals suggests that carbonate minerals are depleted in 25Mgand 26Mg, whereas silicate minerals preferentially retain25Mg and 26Mg (Galy et al., 2002; Young and Galy,2004; Tipper et al., 2006b), though there are still uncertain-ties (Pogge von Strandmann et al., 2008). Galy et al. (2002)and Carder et al. (2005) report magnesium isotope fraction-ation factors (e26=24

s�f ; where e26=24s�f ¼ 1000� ðaÞ26=24

s�f � 1 and(a26=24

s�f ¼ ð26Mg=24MgÞs � ð24Mg=26MgÞf ) for the precipita-tion of speleothem calcite and dolomite of �2.7& and�1.4&, respectively, whereas even larger fractionationfactors (e26=24

s�f � �3:5 to � 4:5&) have been observed inforaminiferal calcite (Chang et al., 2004; Pogge von Strand-mann, 2008). Estimates of magnesium isotope fractionationduring the precipitation of silicate minerals are primarilyfrom measured soil samples and shales, both of which areenriched in 26Mg/24Mg (e26=24

s�f � þ0:3 to þ 0:5&) com-pared to primary silicate rocks (Tipper et al., 2006a; Tenget al., 2007). As a result, cycling of magnesium in sedimen-tary pore-fluids may be expected to produce different pro-files in magnesium isotopes depending on whether or notthe primary magnesium sources and sinks are carbonateor silicate minerals. In addition to identifying magnesiumsources and sinks, pore-fluid profiles of magnesium isotopescan also be used to estimate magnesium isotopic fraction-ation associated with the precipitation of Mg-minerals, inparticular dolomite (Ca0.5Mg0.5CO3) and Mg-clays.Because dolomite and clay minerals are thought to be theprimary sinks of magnesium in seawater on geologic time-scales (Hardie, 1996; Holland and Zimmerman, 2000),estimates of magnesium isotope fractionation factors asso-ciated with the precipitation of these minerals under naturalconditions are important for understanding how the mag-nesium isotopic composition of seawater may have changedover geologic time.

In this paper, we present the first magnesium isotopicdata on pore-fluids from six ODP sites (925, 984, 1012,1082, 1086, and 1219). We also present magnesium isotopedata for six dolomite samples from sites 1082 and 1012. Weuse the data to identify sources and sinks of magnesium indeep-sea sediment pore-fluids. We show that similar magne-sium concentration depth profiles are associated with differ-

ent depth profiles in magnesium isotopes. Using a diffusion-reaction model of magnesium isotopes in deep-sea sedimentpore-fluids, we estimate magnesium isotope fractionationfactors associated with net magnesium addition to and/orremoval from the pore-fluid. We discuss implications formagnesium isotope fractionation and magnesium cyclingin deep-sea sediments.

1.1. Sample locations and description

ODP sites used in this study span a range of deposi-tional settings and sediment types (Figs. 1 and 2). Cl�

content of the pore-fluid is similar at all sites and at alldepths (538–575 mM), indicating that subsurface brinesare not important. Sites 1082, 1086, and 1012 are locatednear upwelling regions on the western continental marginof Africa and North America, respectively (Lyle et al.,1997; Wefer et al., 1998). Sedimentation rates at thesesites are rapid (average 4–12 cm/kyr) and cored intervalsrange in age from late Miocene to the late Pleistocene.Sediments at sites 1082, 1086, and 1012 are characterizedby significant contributions from both biogenic and detri-tal sources. The biogenic component of the sediment isprimarily CaCO3, averaging 41–78 weight% (wt.%); con-tributions from siliceous organisms (diatoms, radiolari-ans, sponge spicules) are variable, but generally minor.The detrital component of the sediment at these sites isdominated by clay minerals (smectite, kaolinite, illite,and micas) with minor amounts of volcanic glass,feldspar, and quartz. Sediments at these sites alsocontain significant amounts of organic carbon (average1.1–3.7 wt.%). Diagenetic reactions at these sites are dom-inated by organic carbon remineralization. Associatedchanges in the chemical composition of pore-fluids in-clude rapid SO4

2� depletion ([SO42�] = 0 at 22, 28, and

182 m at sites 1082, 1012, and 1086, respectively) andhigh concentrations of NH4

þ (maximum = 25.5, 15.2,and 2.8 mM at sites 1082, 1012, and 1086, respectively)and alkalinity (maximum = 75.6, 67.3, and 17.4 mEq atsites 1082, 1012, and 1086, respectively). Sites 1082 and1012 also contain multiple cm-thick beds of dolomite atdepths >120 m at site 1082 and >170 m at site 1012.Discrete dolomite beds are not observed in the coredsequence from site 1086.

Site 984 is a high-sedimentation-rate (average �14 cm/kyr) pelagic sequence located on the Bjorn Drift on theeastern flank of the Reykjanes Ridge. Sediments at this siterange in age from the late Pliocene to the Holocene and arecomposed primarily of fine-grained terrigenous silts andclays transported from Iceland (Jansen et al., 1996). Themineralogy of the clastic component is dominated by pla-gioclase, smectite, and quartz, with minor amounts of illite,kaolinite, and chlorite (Carter and Raymo, 1999). Carbon-ate and organic carbon contents at site 984 are low, averag-ing �8 wt.% and <0.25 wt.%, respectively. However,because sedimentation rates are high, organic carbon accu-mulates rapidly and observed changes in pore-fluid chemis-try are dominated by respiration, as indicated by thecomplete consumption of SO4

2� ([SO42�] = 0 mM at

120 m) and elevated concentrations of NH4þ (maxi-

90°E 120° 150°E 180° 150°W 120° 90° 60° 30°W 0° 30°E 60°

60° S

40°

20°

20°

40°

60°

80° N

0°

925(3042 mbsl)

1219(5063 mbsl)

1012(1773 mbsl)

1082(1279 mbsl)

1086(781 mbsl)

984(1649 mbsl)

ARCTICOCEAN

PACIFICOCEAN

SOUTHERNOCEAN

ATLANTICOCEAN

INDIANOCEAN

Fig. 1. ODP sites measured in this study. See text for further details.

Magnesium isotopes in sediment pore fluids 5041

mum = 3.1 mM at 290 m) and alkalinity (maxi-mum = 11 mEq between 80 and 140 m).

Sites 925 and 1219 are located on the Ceara Rise in thewestern tropical Atlantic, and the central Pacific abyssalplain, respectively (Curry et al., 1995; Lyle et al., 2002). Thesesites are typical of deep-water pelagic sequences in that theyare characterized by slow sedimentation rates (averaging2.4 and <1 cm/kyr, respectively) and very low organic carboncontents (averaging 0.07 and 0.08 wt.%, respectively). Sedi-ments at site 925 consist primarily of CaCO3 (averaging66 wt.%) with variable amounts of clay minerals. At site1219, sediments consist primarily of nanofossil and radiolar-ian oozes with varying clay content. Highest clay contents oc-cur in the upper 30 m (>50 wt.% smectite + illite), whereasthe rest of the sequence is composed principally of CaCO3

(average �75 wt.% between 30 and 170 m) or opal (average�70 wt.% between 170 and 230 m). In contrast to sites1082, 1086, 1012, and 984, changes in the chemical composi-tion of pore-fluids at sites 925 and 1219 are not dominated bythe respiration of organic carbon. SO4

2� concentrations de-crease with depth, but the decline occurs over a much largerdepth scale and non-zero sulfate concentrations are found inthe deepest measured samples at both sites 1219 ([SO4

2�] =26.6 mM at 200 m) and 925 ([SO4

2�] = 7.2 mM at 740 m).Similarly, both NH4

þ concentrations (maximum = 7 lMand 1.4 mM at sites 1219 and 925, respectively) and pore-fluid alkalinity (maximum = 3.1 and 10.6 mEq at sites 1219and 925, respectively) remain relatively low at both sites.

1.2. Magnesium concentration profiles

Magnesium concentrations decrease down core at allsites, though the shape and magnitude of the decline varies(Fig. 2). Although large increases in the concentration of

magnesium in seawater cannot be excluded as the reasonfor lower magnesium concentrations at depth (Fantle andDePaolo, 2006), the observed magnesium profiles have beeninterpreted to reflect net precipitation of Mg-minerals(dolomite and Mg-clays) in the sediment column and/orunderlying basalt (Lawrence et al., 1975; Gieskes and Law-rence, 1981; Baker, 1985; Moore et al., 2004). Similarly, thetransient increase in magnesium concentration observed attwo sites (1082 and 1012; Fig. 2), indicates a local net mag-nesium source. Possible sources include the release of mag-nesium adsorbed to sediment–particle surfaces by ionexchange with pore-fluid NH4

þ (Von Breymann and Suess,1988; Vonbreymann et al., 1990), weathering of igneousminerals (Gieskes and Lawrence, 1981), and dissolutionof carbonates (calcite and aragonite; Fantle and DePaolo,2006). At sites 1082, 1086, and 1012, dolomite precipitationis thought to be the principal magnesium sink, at least inthe upper portion of the sediment column (Lyle et al.,1997; Wefer et al., 1998; Moore et al., 2004). At these sites,the decline in magnesium is accompanied by a similar de-cline in calcium, consistent with dolomite precipitation.However, the ratio of the change in magnesium and cal-cium (DMg/DCa) varies significantly between sites, from+0.71 in the upper 50 m at site 1012 to +1.09 in the upper40 m at site 1082 to +2.30 in the upper 100 m at site 1086(Fig. 2a–c). The chemical composition of pore-fluids atthese sites is also suggestive of dolomite precipitation assimilar chemical conditions (high alkalinity, depleted inSO4

2�) are associated with dolomite precipitation in a rangeof modern sedimentary systems (Baker, 1985; Vasconcelosand McKenzie, 1997; Mazzullo, 2000; Teal et al., 2000). Fi-nally, the presence of multiple cm-thick beds of authigenicdolomite at sites 1082 and 1012 indicate once active dolo-mite precipitation in the sediment column at these sites.

0 10

100

200

300

400

500

0D

epth

(mbs

f)

Mg and Ca (mM)20 30 40 50 60

nann

o. &

fora

m. c

lay

nann

o-ric

hdi

atom

aceo

us c

lay

nann

o.cl

ayna

nno.

ooze

Mg (mM)

Ca (mM)

Site 1082

Mg and Ca (mM)0 10 20 30 40 50 60

50

100

150

200

250

0

Dep

th (m

bsf)

Fora

min

ifer-

nann

ofos

sil o

oze

Mg (mM)

Ca (mM)

Site 1086

d

d

ddddddd

d dolomitebed

0 10

60

120

180

240

300

0

Dep

th (m

bsf)

Mg and Ca (mM)20 30 40 50 60

Nan

nofo

ssil

clay

N

anno

foss

il cl

ay to

ooz

e

Site 1012

d

d

d

Mg (mM)

Ca (mM)

d dolomitebed

0 10

100

200

300

400

500

0

Dep

th (m

bsf)

Mg and Ca (mM)20 30 40 50 60

clay

Site 984

Mg (mM)

Ca (mM)

silty

cla

yna

nnof

ossi

l cla

y

0 10

50

100

150

200

250

0

Dep

th (m

bsf)

Mg and Ca (mM)20 30 40 50 60

Rad

io.

clay

Site 1219

Mg (mM)

Ca (mM)

Nan

nofo

ssil

ooze

Rad

iola

rian

clay

/ooz

eB

asal

tCal

c.ch

alk

0 10

150

300

450

600

750

0

Dep

th (m

bsf)

Mg and Ca (mM)20 30 40 50 60

Site 925

Mg (mM)

Ca (mM)

Nan

no.

clay

Nan

nofo

ssil

chal

k/lim

esto

neN

anno

.oo

ze

(a)

(b)

(c)

(d)

(e)

(f)

Age(Ma)

5

0

1

2

3

4

Age(Ma)

0

1

2

3

Age(Ma)

2

4

6

0

Age(Ma)

010

30

40

50

20

Age(Ma)

0

1

3

4

2

Age(Ma)

0

10

30

20

Fig. 2. Magnesium (�) and calcium (x) concentration profiles, lithologies, and sediment ages for the studied sites. See text for further details.


At sites 984, 925, and 1219, magnesium removal hasbeen attributed to the precipitation of Mg-clay minerals(e.g., smectites) in the sediment column and/or underlyingbasalt (Curry et al., 1995; Jansen et al., 1996; Lyle et al.,2002). At site 1219, a nearly one to one relationship existsbetween changes in magnesium and calcium concentrationswith depth (DMg/DCa = �0.95; Fig. 2e). At site 925, a sim-

ilar relationship is observed; however, the decline in magne-sium is �50% larger than the increase in calcium (DMg/DCa = �1.54; Fig. 2f). Profiles of this type are thought toreflect submarine weathering of basalt (calcium source)and associated precipitation of clay minerals (magnesiumsink) at low temperatures, an interpretation that is consis-tent with the observed decline in pore-water d18O values


(Lawrence et al., 1975; Schrag et al., 1996). Magnesium andcalcium profiles at site 984 are similar in many respects tothose associated with dolomite precipitation in the upper70 m of the sediment column, where the decline in magne-sium is accompanied by a decline in calcium (DMg/DCa = +1.87; Fig. 2d). Changes in pore-fluid chemistry(declining SO42�, increasing alkalinity) at site 984 are alsoconsistent with dolomite precipitation. However, the abun-dance of detrital basalt from Iceland, together with mea-sured rates of mineral dissolution and precipitation frompore-fluid U isotopes (Maher et al., 2006) and lack of evi-dence for authigenic dolomite, suggest that the decline inmagnesium at site 984 could also be due to low-temperatureweathering of basalt and associated formation of Mg-clayswithin the sediment column.

2. METHODS

Precise and accurate magnesium isotope measurementsusing inductively coupled mass spectrometry (ICP-MS) re-quire that contributions from molecular interferences, (e.g.,CN+), doubly-charged ions (e.g., 48Ca2+), and species otherthan Mg that affect instrument fractionation (matrix effects(e.g., Ca, Na; Galy et al., 2001) be reduced. To reducemolecular interferences associated with species containingO, N, C, and H, samples are introduced to the mass spec-trometer using an Elemental Scientific, Inc. Apex Q nebu-lizer fitted with an ACM membrane desolvator. Tominimize contributions from doubly-charged ions and vari-ations in the composition of the sample matrix, samplescontaining 2–4 lg of Mg were loaded in 1 N HNO3 acidand passed twice through ion exchange columns (Bio-RadX12 resin). Ultra-pure acids (VWR Omnitrace) and 18-Xde-ionized water were used to assure small proceduralblanks (<10 ng Mg or <0.33% of the sample size). Our col-umn procedure efficiently separates magnesium from othercations and achieves yields that are >99%. Purity of sepa-rated samples was verified using a Thermo-Finnigan X-Ser-ies quadrupole ICP-MS with a H2/He collision cell in theLaboratory of Geochemical Petrology at Harvard Univer-sity. A dilution series of a multi-element standard was usedto construct elemental calibration curves. Machine driftwas corrected by interpolation from repeated analyses ofthe multi-element standards during a sample run. Externalerrors (1rSD) from repeat analysis of a 10 ppb multi-ele-ment standard are <15% for all elements. Final cation(Na, Ca, Mn, Sr, Fe, Al, K)/Mg ratios for all samples mea-sured for magnesium isotopes are <0.05 (g/g).

Pore-fluid samples were obtained from the Ocean Dril-ling Program and taken from 5- to 15-cm whole-round sec-tions of the core while at sea (Sayles and Manheim, 1975).Collected pore-fluid samples were then passed through a0.45-lm Gelman polysulfone filter and stored in glass amp-ules for later analyses. Major element compositions, salin-ity, pH, and alkalinity of the pore-fluid were all measuredwhile at sea with a precision of 2–3%. Samples were centri-fuged and diluted prior to ion exchange chromatography,but additional filtering was not required as the magnesiumcontents of the studied pore-fluids are high (tens of mM),making significant contamination from nanophases unli-

kely. The presence of dolomite in lithified samples fromsites 1082 and 1012 was inferred from ODP reports (Lyleet al., 1997; Wefer et al., 1998; Pufahl and Wefer, 2001)and verified by XRD on a Scintag XDS 2000 diffractome-ter, using CuKa1 radiation at �40 kV and 30 mA at Har-vard University. The non-carbonate fractions of thesamples were high (>60 wt.%). Approximately 1 mg ofpowdered sample was dissolved in 12 mL of 0.1 N ace-tate-buffered acetic acid for 4 h at 25 �C to minimize thecontribution of Mg from the non-carbonate fraction. Thesamples were then centrifuged for 20 min and the superna-tant removed for Mg separation. Mg/Ca ratios of the dis-solved sample range from 0.95 to 1.12 and are consistentwith dolomite as the principle source of Mg. However,there is a small excess of magnesium of �10% comparedto stoichiometric dolomite in two samples, suggesting somecontamination of non-dolomite Mg from ion exchangeand/or dissolution of soluble clays. We consider the conse-quences of this contamination for the measured d26Mg val-ues. Solid samples were also measured for carbon andoxygen isotopes on a VG Optima dual inlet mass spectrom-eter attached to a VG Isocarb preparation device in theLaboratory of Geochemical Oceanography at HarvardUniversity. d13C and d18O values were calibrated to VPDBusing the Cararra Marble standard. External errors (1rSD)of the measurements are typically < 0.1&.

Magnesium isotopic analyses were carried out on a GVIsoprobe P multi-collector inductively coupled plasma massspectrometer (MC-ICP-MS). Samples were diluted to150 ppb Mg in 2% HNO3 to give 1.5–3 V on mass 24. Datawere acquired in blocks of 15–20 ratios of 10 s integrationtimes. Internal reproducibility of 25Mg/24Mg and26Mg/24Mg ratios was typically <0.1& and 0.2& (1rSD),respectively. Sample-standard-sample bracketing was usedto correct for instrumental mass fractionation (e.g., Galyet al., 2001). Magnesium isotope ratios are reported usingdelta notation:

d25;26Mg ¼ ð25;26Mg=24MgÞSAM

ð25;26Mg=24MgÞSTD

� 1

� �� 1000

where (25;26Mg=24Mg) is the measured 26Mg/24Mg ratio ofthe DSM-3 Mg standard (Galy et al., 2003). Repeat mea-surements of the Cambridge-1 Mg standard yieldedd25Mg and d26Mg values of �1.32 ± 0.13& and �2.62 ±0.25& (2rSD external; n > 200) relative to DSM-3, indis-tinguishable from published values (d25Mg = �1.33 ±0.07&; d26Mg = �2.58 ± 0.14&, Galy et al., 2003). To ver-ify that our sample procedure did not fractionate magne-sium isotopes we compared measured d26Mg values ofpure Mg standards to both pure and cation-spiked Mgstandards that had been run through our ion exchange pro-cedure (Table 1). In addition, repeat separation and isoto-pic analyses (N = 5) of seawater from Bermuda yieldedd25Mg and d26Mg values of �0.41 ± 0.09& and �0.79 ±0.18&, respectively, indistinguishable from published val-ues (d25Mg = �0.40 ± 0.07&; d26Mg = �0.81 ± 0.14&

(Galy et al., 2003). Plotted in three-isotope space (d25Mgvs. d26Mg; Fig. 3) the samples measured in this study lieon a line with a slope of 0.527 ± 0.009 (R2 = 0.996).This slope is very similar to that expected for equilibrium

Table 1Measured and expected d26Mg values for magnesium isotopestandards. HU-1 is a pure Mg solution spiked with equal amountsof Na, Ca, K, Al, Mn, Fe, and Sr. Analytical uncertainties arereported as the 2r of repeat measurements on the MC-ICP-MS. N,number of total sample replicates; n, number of repeat measure-ments on the MC-ICP-MS.

Sample d26Mg expected(&)

d26Mg measured(&)

2r N(n)

Cambridge-1 �2.58 �2.62 0.25 0(>200)DSM-3 0.0 �0.02 0.04 1(4)seawater �0.82 �0.79 0.18 5(20)spiked HU-1 �2.76 �2.81 0.27 4(16)


fractionation (0.521; Young and Galy, 2004) as well as pre-viously published terrestrial fractionation slopes (0.513,Chang et al., 2004; 0.519, Young and Galy, 2004; Poggevon Strandmann, 2008; Huang et al., 2009), and demon-strates that the variability in magnesium isotopes that weobserve is due to mass-dependent fractionation. For clarity,we present and discuss only d26Mg values.

3. NUMERICAL MODEL OF MAGNESIUM IN

SEDIMENT PORE-FLUID

We developed a numerical model to quantify the cyclingof magnesium in deep-sea pore-fluids. We use the model,together with measured magnesium concentration andd26Mg profiles to quantify rates of net magnesium cyclingand try and identify net magnesium sources and sinks inthe sediment column and underlying crust. We model thedeposition of a sediment column and the chemical cyclingof its pore-fluid using boxes of unconsolidated sediment,deposited at rates determined by the age and thickness ofthe sediment column, and compacted according to the ob-served porosity/depth relation (Richter and DePaolo,1987; Richter and Depaolo, 1988; Schrag et al., 1992). Wecalculate down-core trends in magnesium concentrations

2σext

δ25Mg = δ26Mg x 0.527 (± 0.009) - 0.008 (± 0.008)

δ25Mg (‰)

R2 = 0.996

δ26Mg (‰)

-3.0 -1.0-2.0 0.0 +1.0 +2.0-1.5

-1.0

-0.5

0.0

+0.5

+1.0

+1.5

pore fluid

bulk solid

Fig. 3. Three isotope plot (d25Mg vs. d26Mg) for all samplesmeasured in this study. The slope of 0.5271 (uncertainties define the95% confidence envelope) is similar to published values andindicates mass-dependent behavior of magnesium isotope fraction-ation in this study.

and isotopes that result from reaction with the sedimentand/or the underlying crust and transport by diffusionand advection. Initial conditions for the pore-fluid are ac-quired from the concentration and isotopic compositionof magnesium in seawater at the time of sediment deposi-tion. Boundary conditions for the model are the concentra-tion and isotopic composition of magnesium in seawaterthrough time (upper) and either zero-flux or finite-flux (low-er), depending on the observed magnesium concentrationprofile at that site.

Our model is based on a 1D diffusion–advection–reactionequation used by many authors (Berner, 1980; Richter andDePaolo, 1987; Schrag et al., 1992; Fantle and DePaolo,2006):

/ðzÞ � @ðnMgÞ@t

¼ @

@z/ðzÞ � DnMg �

@nMgðzÞ@z

� �� @

@zð/ðzÞ

� xðzÞ � nMgðzÞÞ � /ðzÞ � RnMgðzÞ ð1Þ

where nMg is the concentration of 24Mg, 25Mg, or 26Mg inmass per unit volume of pore-fluid, z is depth below the sed-iment–water interface, DnMg

is the depth-dependent bulk

sediment diffusion coefficient (m2 yr�1) for 24Mg, 25Mg, or26Mg, / is the depth-dependent but time-invariant porosity,x is the advective velocity (m yr�1), and RnMg

is a source-

sink term for 24Mg, 25Mg, or 26Mg (mol Mg m�3 yr�1).Porosity (/) is obtained from fits to observed porosity pro-files at each site. Bulk diffusion coefficients for magnesium(DnMg

) depend on temperature and tortuosity. The effect

of temperature on DnMgis modeled after Boudreau (1997)

using observed temperature profiles or geothermal gradi-ents. Tortuosity is taken into account following Boudreau

(1997) where DnMg¼

DonMg

h2 and DonMg

is the diffusion coefficient

for magnesium in solution (Li and Gregory, 1974). Tortu-osity (h) is related to porosity (/) by the function:

h2 ¼ 1� lnð/2Þ (Boudreau, 1997). Richter et al. (2006)showed that magnesium isotopes are not significantly frac-

tionated during diffusion (a26=24diffusion = 1.00006). As a result,

we use a single diffusion coefficient for 24Mg, 25Mg, and26Mg. Sediment compaction decreases sediment porosityand drives pore-fluid toward the sediment–water interface.These advective velocities (x) are calculated for each siteusing observed porosity profiles and sedimentation rates.At each time step where a unit of sediment is added, com-paction of the other sediment layers is calculated and thereduction in pore space is converted into a flux of pore-fluidthat is oriented towards the sediment–water interface. At allother time steps, advection of the pore-fluid is zero. Othersources of advection (e.g., hydrothermal convection) are as-sumed to be minor at sites 1082, 1086, 1012, 925, and 984due to the thickness and/or clay content of the sedimentcover (Anderson and Langseth, 1977; McDuff, 1981).External fluid flow is considered at site 1219, where the sed-iment cover is relatively thin (240 m).

The reaction term, RnMg, is a tunable model parameter

that reflects the net addition/removal of magnesium fromthe pore-fluid. Each unit of sediment is associated with avalue of RnMg

for each isotope of magnesium. We assumethat the magnitude of RnMg

is a function of depth below


the sediment–water interface (and hence sediment age) andthat values of RnMg

are time-invariant for each depth inter-val. Our treatment of RnMg

differs from previous studies(Richter and DePaolo, 1987, 1988; Fantle and DePaolo,2006, 2007) in that we do not assume a particular form ofthe reaction rate expression with depth/age. We do thisfor three reasons: (1) the precipitation of diagenetic miner-als can be localized at discrete boundaries in the sedimentcolumn that may migrate in response to changes in sedi-ment composition (Meister et al., 2007, 2008), (2) sedimentsat the studied sites are heterogeneous in composition somultiple sources/sinks may be expected, and (3) we havefew constraints on gross rates of magnesium cycling. A con-sequence of our approach is that RnMg

does not reflect theuptake/loss of magnesium from an individual mineralphase but rather the net change in magnesium resultingfrom all precipitation, dissolution, and exchange reactions.Using a simple iterative routine, we tune RnMg

at each depthinterval for each isotope (e.g., R24Mg

, R25Mg, and R26Mg

) to fitboth the magnesium concentration and isotope depth pro-files. The quality of the model fit is evaluated by visualinspection. Using the tuned reaction rates of 25Mg, 26Mg,and 24Mg as a function of depth, we calculate the ratio ofthe isotopic composition of the flux to that of the fluid aswell as the average isotopic composition of net magnesiumsources (d26Mgsource). The ratio of the isotopic compositionof the flux to the fluid can be understood as an empiricalisotopic fractionation factor associated with net magnesiumuptake/addition:

e26=24s�f ¼ 1000 � a26=24

s�f � 1� �

; ð2Þ

where a26=24s�f ¼

R26Mg

R24Mg

�24Mg26Mg

� �pore-fluid

It is important to emphasize that the fractionation fac-tors calculated by the model reflect the summed isotopic ef-fect of all local precipitation, dissolution, and exchangereactions. This is because the isotopic composition of mag-nesium in the pore-fluid depends on the gross fluxes of mag-nesium through the pore-fluid. In contrast, changes inmagnesium concentrations with depth (and hence the mod-eled reaction rates) reflect only the net uptake/addition ofmagnesium. This difference complicates our interpretationof the modeled magnesium isotope fractionation factorsand is explored in detail in the discussion.

We solve Eq. (1) for each isotope of magnesium (24Mg,25Mg, and 26Mg) using an explicit finite difference formulafor an uneven grid (Pearson, 1968). Grid spacing reflects thechange in sediment thickness with depth due to compac-tion. As a result, the vertical resolution of the model in-creases with depth as sediments with uniform initialthickness are compacted according to the observed theporosity–depth relation. The thickness of units of un-com-pacted sediment at the studied sites ranges from 6.9 to26.6 m. Increasing the vertical resolution of the grid (i.e.,depositing smaller units of sediment more frequently) bya factor of three does not significantly change the resultspresented here. Time steps (Dt) are chosen to avoid numer-ical instability (Boudreau, 1997) and range from 210 to2000 yrs. The total number of time steps that the model is

run is determined by the age of the sediment column at eachsite (see Fig. 2). Boundary conditions are prescribed at thesediment–water interface and the bottom of the cored sed-iment column. The magnesium concentration and isotopiccomposition at the sediment–water interface are set con-stant at the values for modern seawater ([Mg] = 53 mM,d25Mg = �0.41& and d26Mg = �0.79&). Boundary condi-tions at the bottom of the sediment column are either zero

dMgdz jzend

¼ 0� �

or constant dMgdz jzend

¼ const:� �

flux, depend-

ing on the magnesium concentration profile at each site.

4. RESULTS

Pore-fluid magnesium isotope data measured in thisstudy are presented in Table 2. At all sites, measuredd26Mg profiles show systematic trends with depth. At sites1012, 1082, and 1086, d26Mg values increase with depthby 1.25–2&. For example, at site 1082, d26Mg values in-crease from �0.74& at 1.4 m to +1.03& at 461.6 m. Simi-lar increases in d26Mg values are observed at sites 1086(�0.74& at 1.4 m to +0.59& at 201.1 m) and 1012(�0.77& at 3 m to +1.13& at 267.1 m). In contrast, at sites925, 1219, and 984 pore-fluid d26Mg values decrease withdepth. The decrease in d26Mg values at sites 1219 and 925is small (0.1–0.2&), from �0.71& at 3 m to �0.91& at200.5 m and �0.59& at 3 m to �0.72& at 424 m, respec-tively. A larger decline in d26Mg values (1.9&) is observedat site 984, from �0.75& at 9.6 m to �2.62& at 271 m. Themagnesium isotopic composition of measured dolomitesamples from sites 1082 and 1012 are also presented inTable 1. At site 1082, d26Mg values for dolomite range from�2.02& at 128.6 m below seafloor (mbsf) and �2.06& at199.7 mbsf to �1.72& at 327.5 mbsf. At site 1012, dolomitesamples are slightly more depleted, with d26Mg values rang-ing from �2.52& at 169–170 mbsf to �2.38& at 244.9mbsf.

Modeled magnesium concentration and d26Mg profilesare shown in Figs. 4 and 5. At all sites, declining magnesiumconcentrations with depth can be explained by net precipita-tion of Mg-minerals in the sediment column and/or underly-ing basalt. Excursions towards higher magnesiumconcentrations at sites 1082 and 1012 suggest a local netmagnesium source within the sediment column. Modeledrates of net magnesium uptake/addition from/to the pore-fluid range from <10�6 to 10�4 mol Mg m�3 yr�1. Theserates are small and suggest that the net contribution ofnewly precipitated Mg-minerals will be small compared tothe mass of sediment. For example, dolomite precipitationat a rate of 10�4 mol Mg m�3 yr�1 would have to be main-tained for �19 million years to fully cement sediment withan initial porosity of 70%. As these rates are likely main-tained for only a fraction of this time (i.e., the residence timeof sediment in the depth interval associated with those highrates), authigenic minerals are expected to make only a smallnet contribution (<5 wt.%) to the bulk sediment. In light ofthese slow precipitation rates, the formation of the thick(up to 10’s of centimeters) layers of dolomite at sites 1082and 1012 seems problematic. One possibility is that a largefraction of the magnesium precipitated in the dolomite is

Table 2Measured d26Mg values for pore-fluids and dolomite samples from studied sites. Analytical uncertainties are reported as the 2r of repeatmeasurements on the MC-ICP-MS. N, number of total sample replicates; n, number of repeat measurements on the MC-ICP-MS.

Sample Depth (mbsf) [Mg] (mM) d26Mg (&) 2r N(n) Sample Depth (mbsf) [Mg] (mM) d26Mg (&) 2r N(n)

Seawater 0 53 �0.81 0.13 5(20)Pore-fluid Pore-fluid

Site 1082 1.4 53.43 �0.74 0.17 2(7) Site 984 9.6 47.9 �0.75 0.27 1(5)4.4 52.19 �0.74 0.15 1(3) 30.1 46.3 �0.92 0.03 1(2)7.5 52.52 �0.85 0.25 1(3) 49.1 43 �0.95 0.17 1(4)

20 46.13 �0.29 0.07 1(3) 68.1 37.8 �0.94 0.22 1(5)40.4 44.41 0.07 0.01 1(2) 96.6 32.7 �1.04 0.17 1(5)54.1 46.86 0.19 0.15 1(4) 153.6 28.2 �1.20 0.20 1(6)97.5 44.23 0.38 0.24 1(3) 185.5 24.9 �1.46 0.24 1(5)

160 39.06 0.86 0.25 1(4) 271.0 17.7 �1.57 0.06 1(6)365 26.89 0.97 0.30 1(3) 346.2 10.8 �2.07 0.25 1(5)461.6 23.51 1.03 0.05 1(2) 498.5 5.3 �2.56 0.18 1(3)

Site 1086 1.4 52.4 �0.74 0.18 1(4) Site 1219 3.0 48.6 �0.71 0.09 1(6)20.6 50.71 �0.57 0.11 1(5) 39.0 30.4 �0.77 0.18 1(6)39.6 47.97 �0.51 0.07 1(5) 48.5 27.5 �0.73 0.08 1(3)58.6 45.84 �0.41 0.16 1(4) 105.5 22.9 �0.84 0.25 1(6)77.5 41.87 �0.37 0.03 1(4) 200.5 20.1 �0.91 0.14 1(3)

153.6 33.7 0.15 0.17 1(4)Site 925 3.0 51.75 �0.59 0.17 1(6)182.1 31.75 0.32 0.17 1(5)

46.5 48.91 �0.53 0.13 1(4)201.1 30.5 0.59 0.18 1(4)94.5 47.75 �0.51 0.19 1(6)

Site 1012 3.0 51.3 �0.77 0.26 1(3)293.9 41.38 �0.58 0.03 1(3)

28.2 47.1 �0.41 0.09 1(6)336.0 40.51 �0.70 0.22 1(6)

47.2 45.7 �0.04 0.17 1(6)424.0 35.36 �0.72 0.07 1(3)

75.7 49.3 �0.16 0.18 1(6)104.3 50.2 0.12 0.13 2(11) Dolomite

131.4 50.4 0.00 0.19 1(6) Site 1082 128.6 �2.02 0.17 1(4)161.7 49.2 0.18 0.07 1(5) 199.7 �2.06 0.21 1(5)211.8 40.7 0.56 0.18 1(5) 327.5 �1.72 0.21 1(4)

240.5 37.9 0.98 0.08 1(6)Site 1012 169.9 �2.52 0.26 1(5)267.1 36.6 1.13 0.14 1(4)

170 �2.52 0.23 1(3)244.9 �2.38 0.26 1(6)


sourced locally (i.e., from the surrounding sediment andnot from seawater). However, comparisons between themodeled fractionation factor and estimates of the fraction-ation factor from the d26Mg value of the dolomite itselfsuggest that the vast majority of the magnesium in thedolomite is derived from seawater (i.e., local exchange reac-tions are minor – see Section 5). A more likely explanationfor the massive layers of diagenetic dolomite is that they re-flect periods when the sulfate–methane interface shoaled dra-matically and/or remained fixed at a particular depth due tochanges in the composition of deposited sediments underconditions of non-steady state diagenesis (Meister et al.,2007, 2008). Reaction rates explicitly linked to the composi-tion of deposited sediments are not explored here, though itmay be appropriate to investigate this in future models.

Modeled magnesium isotope fractionation factors(e26=24

s�f ) range from �2.7& to +1.25& and vary systemati-cally between sites (Table 3 and Figs. 4, 5). At sites wherethe decline in magnesium concentrations in the pore-fluidis thought to reflect dolomite precipitation (1082, 1086,and 1012), the modeled net magnesium sinks in the upperpart of the sediment columns are significantly depleted in

26Mg relative to the pore-fluid (e26=24s�f = �2.0& to �2.7&;

Fig. 4). In contrast, at sites where magnesium cycling isthought to reflect alteration of the underlying crust (1219and 925), the modeled net magnesium sinks are slightly en-riched in 26Mg (e26=24

s�f = 0.0& to +1.25&; Fig. 5). Similarly,the observed increase in e26=24

s�f with depth at site 1082 mayalso indicate increased precipitation of clay minerals atgreater depths. At site 984, where the identity of the magne-sium sink is more ambiguous, the observed increase in thed26Mg values of the pore-fluid is consistent with a net mag-nesium sink both within and below the cored sediment col-umn that is 0.95& to 1.25& enriched in 26Mg. Althoughonly sites 1082 and 1012 are characterized by a net magne-sium source within the sediment column, modeled d26Mgvalues of these sources are similar, at 0.0& to �0.2&.

We examine the sensitivity of modeled d26Mg profiles toestimates of e26=24

s�f by varying the modeled fractionation fac-tors at each site by ±0.5&. As shown in Figs. 4 and 5, theanalytical uncertainties of our measured d26Mg profilestranslate into uncertainties of �±0.5& for modeled frac-tionation factors and values of d26Mgsource (Figs. 4 and5). Reactions near the seafloor have somewhat larger

-1.5 -0.5 +0.5 +1.5

20 30 50

[Mg]model

δ26Mgmodel

10-5 10-4 10-3

ε ± 0.50‰

δ26Mg (‰)

Mg (mM)site 1082

ε = -2.4 to -0.32‰ (sink; inc. w/depth)δ26Mgsource = 0 to -0.1‰ (source)

ε1 = -2.0 to -2.1‰ (sink)

0

100

200

300

400

500

Dep

th (m

bsf)

(a)10-6

net reaction rate(mol Mg m-3 yr-1)

40

26/24

26/24

-1.5 -0.5 +0.5 +1.5

20 40 60

[Mg]model

δ26Mg (‰)

Mg (mM)site 1086

ε1 = -2.6‰ (sink)

0

50

100

150

200

250

δ26Mgmodel

ε1 ± 0.50‰

Dep

th (m

bsf)

(b)10-5


10-4 10-3

0

26/24

(c)-1.5 -0.5 +0.5 +1.5

35 45 55

[Mg]model

10-5

δ26Mg (‰) net reaction rate(mol Mg m-3 yr-1)

Mg (mM)site 1012

0

60

120

180

240

300

δ26Mgmodelε1 ± 0.50‰

Dep

th (m

bsf)

10-4

δ26Mgsource = -0.1 to -0.2‰ (source)ε1 = -2.4 to -2.7‰ (sink)

ε = -2.4 to -0.2‰ (sink)

26/24

26/24

Fig. 4. (a–c) Measured magnesium isotope (d26Mg) profiles and modeled [Mg] and d26Mg profiles for sites 1082, 1086, and 1012. d26Mg datafor each site can be found in Table 1. Model fits are indicated by the solid lines (—). At each site, net reaction rates (Rmodel) for 24Mg, 25Mg,and 26Mg are tuned to fit both concentration and isotope profiles. Each source/sink of magnesium is associated with an isotopic fractionation(e26/24) or composition (d26Mg), indicated by the color of the reaction rate bar. Uncertainties in modeled fractionation factors (e26=24

1 ) arecalculated by varying all fractionation factors by ±0.5& and is indicated by the dotted lines (- - -). (Rmodel) at the bottom of the modeledsediment column is the lower boundary condition and reflects the summed net reaction rate for all of the sediment/crust underlying themodeled sediment column.


-3 -2 -1 0

10 30 50

[Mg]model

δ26Mgmodel

10-5 10-4 10-3

ε1 ± 0.5‰

δ26Mg (‰)

Mg (mM)site 984

0

100

200

300

400

500

Dep

th (m

bsf)

(a)

ε1 = +0.95 to 1.25‰ (sink)26/24


(b)-1.5 -0.5 0 +0.5

20 30 50

[Mg]model

10-7 10-5

δ26Mg (‰)

Mg (mM)site 1219

ε1 = <0.1‰ (sink)

0

50

100

150

200

250

δ26Mgmodelε ± 0.50‰

ε2 = +0.1 to +0.39‰ (sink)

Dep

th (m

bsf)

-1.0

40


10-3

10-7 10-5 10-3

26/24

26/24

(c)-2.0 -1.0 0

60

[Mg]model

δ26Mg (‰)

Mg (mM)site 925

0

150

300

450

600

750

ε2 = +0.6‰ (sink)ε1 = -2.4 to -2.6‰ (sink)

Dep

th (m

bsf)

4020


26/24

26/24

δ26Mgmodel

ε ± 0.50‰

Fig. 5. (a–c) Measured magnesium isotope (d26Mg) profiles and modeled [Mg] and d26Mg profiles for sites 984, 1219, and 925. d26Mg data foreach site can be found in Table 1. Model fits are indicated by the solid lines (—). At each site, net reaction rates (Rmodel) for 24Mg, 25Mg, and26Mg are tuned to fit both concentration and isotope profiles. Each source/sink of magnesium is associated with an isotopic fractionation(e26/24) or composition (d26Mg), indicated by the color of the reaction rate bar. Uncertainties in modeled fractionation factors (e26/24) arecalculated by varying all fractionation factors by ±0.5& and is indicated by the dotted lines (- - -). (Rmodel) at the bottom of the modeledsediment column is the lower boundary condition and reflects the summed net reaction rate for all of the sediment/crust underlying themodeled sediment column.


Table 3Measured d26Mg values, Mg/Ca ratios, d13C values, and d18Ovalues for dolomite samples from sites 1082 and 1012.

Site Depth(mbsf)

Mg/Ca(mol/mol)

d13C(&)

d18O(&)

d26Mg(&)

1082 128.6 1.05 11.50 4.13 �2.021082 199.7 1.04 13.40 5.68 �2.061082 327.5 1.00 5.65 5.38 �1.72

1012 169.9 1.11 5.59 5.20 �2.521012 170 1.12 3.62 4.17 �2.521012 244.9 0.95 12.58 6.07 �2.38


uncertainties in estimated fractionation factors due tobuffering of pore-fluid d26Mg values by the overlying sea-water. Uncertainties in model parameters that change thedistribution and/or relative magnitude of modeled magne-sium fluxes ðRnMg

Þ, such as rates of external fluid flow andthe effect of tortuosity on DnMg

, may also contribute touncertainties in e26=24

s�f and d26Mgsource. However, model cal-culations using alternative parameterizations of tortuosityand diffusion in sedimentary systems (see Boudreau, 1997)suggest that the effect on estimated values of e26=24

s�f andd26Mgsource is small. In addition, rates of external fluid flowat sites 1082, 1086, 1012, 984, and 925 are likely low due tothe hydraulic resistance provided by a thick (>300 m) sedi-ment column (Anderson and Langseth, 1977).

At site 1219, a thin sediment column (240 m), 50% ofwhich is permeable carbonate ooze, may support high ratesof fluid flow. Although the abundance of smectite (bulksediment = 2–3 wt.% Mg; (Lyle et al., 2002)) and presenceof volcanic glass (trace to 3 wt.%) suggest that the modeledmagnesium sink in the upper �25 m of the sediment

Table 4Summary of estimated fractionation factors and d26Mg values of sourcesamples is related to the uncertainty in the d26Mg value of the precipitaconsidered.

Site Depth (mbsf) Type Mg-carbonates

est � e26=24 (&)

1082 0–50 Pore-fluid �2.0 to �2.2>100 Pore-fluid �2.4128.6 Solid �1.7 to �2.2199.7 Solid �1.8 to �2.3327.5 Solid �1.4 to �1.960 to 100 Pore-fluid

1012 0–70 Pore-fluid �2.4 to �2.5>180 Pore-fluid �2.4 to �2.7169.9 Solid �2.2 to �2.7170 Solid �2.2 to �2.7244.9 Solid �2.1 to �2.690–180 Pore-fluid

1086 0 to >200 Pore-fluid �2.6

984 0 to >500 Pore-fluid

1219 0–80 Pore-fluid80 to >240 Pore-fluid

925 0–150 Pore-fluid �2.4 to �2.6>750 Pore-fluid

column (assuming no external flow; Fig. 5b) may be real,we explore the possibility that the magnesium profile at site1219 can be explained by a single magnesium sink in theunderlying basalt. Model calculations indicate that a singlemagnesium sink in the underlying basalt is consistent withthe magnesium profiles at site 1219 if rates of upward (ba-salt to seafloor) fluid flow are �0.04 cm/yr. This rate is fas-ter than rates of advection due to sediment compaction(�0.01–0.001 cm/yr), but significantly lower than estimatesof convective fluid flow (10–100 cm/yr) based on measuredcurvilinear temperature-depth profiles at sites with thin(<100 m) sediment cover (Anderson and Langseth, 1977).Including external fluid flow at site 1219 has two effectson estimated values es�f. First, it eliminates the fraction-ation factor associated with the magnesium sink withinthe sediment column. Second, the value of es�f associatedwith the magnesium sink in the underlying basalt decreasesfrom +0.39& to +0.21&.

4.1. Magnesium isotopic composition of dolomites from site

1082 and 1012

The d26Mg values of dolomite samples from sites 1082and 1012 are all isotopically depleted relative to thepore-fluid, with d26Mg values ranging from �2.06& to�2.52& (Table 4). d13C values for the samples are all signif-icantly enriched relative to seawater (+3.62& to +13.40&),consistent with formation in the zone of methanogenesiswhere the d13C of dissolved inorganic carbon (DIC) is ele-vated due to the production of isotopically depleted CH4.d18O values for the samples range from 4.17& to 6.07&

and are similar to diagenetic dolomites measured elsewhere(Meister et al., 2008). The small Mg excess observed in some

s. Range of estimated fractionation factors for measured dolomiteting solution. � Indicates model results when external fluid flow is

Mg-clays Mg-source

d26Mg (&) est � e26=24 (&) est � d26Mg (&)

�2.02�2.06�1.72

0.0 to �0.1

�2.52�2.52�2.38

�0.1 to �0.2

+0.92 to +1.25

<0.1 none*

+0.1 to +0.4 +0.21*

+0.6


of the samples may be due to contamination from associatedclay minerals. Assuming a d26Mg value of clays of �0&,correcting for this contribution would yield dolomites thatare slightly more depleted in d26Mg (0.1–0.25&).

5. DISCUSSION

Our results show that the isotopic composition of magne-sium in deep-sea pore-fluids exhibits systematic and repro-ducible behavior characterized by trends of both increasingand decreasing d26Mg values with depth. Because, at all sites,magnesium concentrations decline with depth, the simplestexplanation for our data is that the differences between thed26Mg profiles at the studied sites reflect differences in theidentity of the Mg-minerals that are precipitating in the sed-iment column and/or underlying crust. In this context the in-crease in the d26Mg values of the pore-fluid at sites 1082,1086, and 1012 can be explained by the precipitation of aMg-mineral, presumably dolomite, that is depleted ind26Mg relative to the pore-fluid by �2.0& to �2.7&.Although the magnesium isotope fractionation factor associ-ated with the precipitation of authigenic dolomite is not wellknown, modeled fractionation factors are broadly similar tofractionation factors reported in previous work on magne-sium isotopes in dolomite as well as carbonate minerals moregenerally (e.g.,�0.49& to�4.5&; (Galy et al., 2002; Changet al., 2004; Carder et al., 2005). In addition, the modeledd26Mg values of magnesium sources at sites 1082 and 1012are consistent with the d26Mg values of silicate and clay min-erals (e.g., 0& to �0.3&; Young and Galy, 2004; Tipperet al., 2006a; Teng et al., 2007; Pogge von Strandmann,2008). Similarly, the decrease in d26Mg values observed atsites 1219, 925, and 984 could be explained by the precipita-tion of Mg-clays that are enriched in d26Mg relative to thepore-fluid by 0& to +1.25&. This is consistent with theobservation that most published d26Mg values of clay miner-als and shales are slightly enriched in d26Mg relative toprimary silicates (d26Mgclay � d26Mgsilicate � þ0:3 to 0:5&;(Young and Galy, 2004; Tipper et al., 2006a; Teng et al.,2007), though some have argued the opposite based on stud-ies of Mg isotopes in weathered basaltic terrains (Pogge vonStrandmann et al., 2008).

Although this straightforward interpretation of our datais consistent with previously published studies of magne-sium isotopes in Mg-carbonates and Mg-clays, it makestwo critical assumptions about the modeled net fraction-ation factors. First, it assumes that changes in the d26Mgvalues of the pore-fluid can be directly related to the net fluxof magnesium to/from the pore-fluid calculated from theconcentration profile. If there is significant local exchangeof magnesium between sediment and pore-fluid, for exam-ple, due to recrystallization of dolomite, our modeled mag-nesium fluxes will underestimate the actual fluxes ofmagnesium through the pore-fluid. Second, it assumes thateach net source/sink reflects a single Mg-mineral. We ex-plore these assumptions for the studied sites below.

The interpretation of modeled magnesium isotope frac-tionation factors in deep-sea pore-fluids depends on two fac-tors: (1) how much more magnesium moves through thepore-fluid than is indicated by our modeled net reaction rates,

and (2) how many isotopically-distinct Mg-minerals are in-volved. Potential sources of this additional magnesium cy-cling are local exchange between sediment and pore-fluid,either by the recrystallization of a single mineral (e.g., calcite;Richter and DePaolo, 1987, 1988; Fantle and DePaolo, 2006,2007) or the coupled dissolution/precipitation of differentminerals (e.g., weathering of Mg-silicates to Mg-clays andcarbonates). If the flux of magnesium through the pore-fluidis dominated by these local exchange reactions, changes inthe magnesium isotopic composition of the pore-fluid withdepth will depend on the combined effect of the isotopic com-position of the magnesium source and any isotopic fraction-ation associated with magnesium removal. In this case,additional information on gross rates of magnesium cyclingis needed to define a unique solution for the system and deter-mine the real magnesium isotopic fractionation factors asso-ciated with mineral precipitation. In contrast, if the flux ofmagnesium through the pore-fluid is equal to or only slightlyhigher than the modeled net reaction rates, changes in themagnesium isotopic composition of the pore-fluid with depthwill depend primarily on the isotopic composition of the dis-solving solid or any isotopic fractionation associated withprecipitation. In this case, the modeled isotopic fractionationfactor will accurately reflect the d26Mg value of the dissolvingsolid or the real fractionation factor for mineral precipitation(assuming the system can be described by the dissolution orprecipitation of a single mineral).

For most of the studied sites, constraints on gross rates ofmagnesium cycling are lacking and the best we can do is to as-sume that the contribution from local exchange reactions issmall and examine how contributions from potential local ex-change reactions would affect our interpretation of the mod-eled magnesium isotopic fractionation factors. At sites 1082and 1012, we can derive an independent estimate of the frac-tionation factor associated with dolomite precipitation fromthe isotopic composition of the dolomite itself (d26Mg, d13C).By comparing this estimate with our model-derived fraction-ation factors we can place some constraint of the importanceof local exchange reactions at these sites.

5.1. Comparing modeled fractionation factors at sites 1082

and 1012 to estimates from the d26Mg of authigenic

dolomites

The d26Mg values of authigenic dolomites from sites1082 and 1012 provide an independent measure of magne-sium isotope fractionation during dolomite precipitation,though there are uncertainties because the d26Mg value ofthe pore-fluid from which the dolomite precipitated is un-known. As the d26Mg values of the pore-fluid increase withdepth at both sites, the d26Mg value of modern seawater(d26Mg = �0.79&) represents a conservative minimumestimate for the d26Mg value of the precipitating fluid.Additional information on the likely d26Mg value of theprecipitating fluid can be gleaned from the d13C values ofthe dolomites, which range from +3.62& to +13.40&

(Table 3). These d13C values require that precipitationoccurred in the zone of methanogenesis (Meister et al.,2007, 2008). At present at site 1082, the zone of enrichedd13C of DIC occurs below 50 mbsf (Moore et al., 2004).


A similar or even shallower depth range is likely for site1012 based on the pore-fluid SO4

2� and CH4 profiles,though the d13C of pore-fluid DIC at this site has not beenmeasured. At these depths, the d26Mg of the pore-fluid is0.5–1.0& heavier than seawater. Although the mass ofdolomite in these layers suggests that the depth where thedolomite layers formed was likely much closer to thesediment–water interface, assuming that the relationshipbetween the d13C value of the DIC and the d26Mg valueof the pore-fluid was similar when these units precipitated,we calculate fractionation factors for dolomite precipitationthat range from �1.4& to �2.3& at site 1082 and �2.1&

to �2.7& at site 1012. These values compare favorablywith our modeled fractionation factors (Table 2), suggest-ing that, at least at these sites and likely also site 1086,the effect of local exchange reactions on the d26Mg valueof the pore-fluid is minor. However, the match is not exact;modeled fractionation factors at sites 1082 and 1012 aresomewhat larger (e.g., �2.0& to –2.7&) than fractionationfactors estimated from measured dolomite samples. Thisdiscrepancy could be due to a number of factors includinglocal exchange reactions, the precipitation of multipleMg-minerals, more closed-system behavior, and/or kineticisotope effects related to the rate of dolomite precipitation.

0

-6

-5

-4

-3

-2

-1

20 40 60 80 100RnetP (%)

-7

-8

-9

0

ε26/24 (‰)1

ε26/24 model = -2.5 ‰

δ26Mgsource = +1.7‰

δ26Mgsource = 0‰

δ26 Mg source

= -4.5‰

δ26Mgpore-fluid = -0.8‰

Fig. 6. The effect of local magnesium exchange on our interpretationof modeled net fractionation factors. (Rmodel) is the net reaction rate,

P is the rate of gross precipitation, e26=241 is the real fractionation

factor for the precipitation of mineral 1, and e26=24model is the modeled net

fractionation factor. Lines denote different d26Mg values for the localmagnesium source. When Rmodel /P is high, modeled fractionationfactors will accurately reflect real fractionation factors. When Rmodel

/P is low, the d26Mg of the local magnesium source becomesimportant and real fractionation factors may deviate significantlyfrom modeled net fractionation factors.

5.2. Quantifying the effect of local exchange reactions and

multiple Mg-minerals on modeled fractionation factors

At sites where the d26Mg value of the pore-fluid stays thesame or decreases with depth (1219, 925, and 984) we do nothave an independent estimate of either the magnitude of iso-topic fractionation associated with mineral precipitation orthe gross fluxes of magnesium through the pore-fluid. As aresult, at these sites, we cannot exclude the possibility thatlocal exchange reactions and/or multiple Mg-minerals con-tribute to the observed d26Mg profile. However, using iso-tope mass balance and some simple assumptions aboutlikely d26Mg values for dissolving/precipitating minerals,we can at least quantify how these processes would affectour interpretation of modeled fractionation factors. Withregard to the precipitation/dissolution of multiple Mg-min-erals, consider the simple case where two minerals precipi-tate. The modeled fractionation factor can be separatedinto contributions from the two minerals:

e26=24model ¼ e26=24

1 � X 1 þ e26=242 � ð1� X 1Þ ð3Þ

where e26=24model is the modeled fractionation factor, e26=24

1 ande26=24

2 are the actual fractionation factors associated withthe precipitation of mineral 1 and 2, respectively, and X1

is the fraction of the net magnesium sink that occurs as min-eral 1. When e26=24

1 and e26=242 are very different, as is likely the

case for dolomite and Mg-clays (Table 2), minor magnesiumsinks (e.g., (1� X 1) = 10–20%) can create large differencesbetween estimated and real fractionation factors. For exam-ple, considering e26=24

model = �2.1& and assuming a fraction-ation associated with Mg-clays (e26=24

2 ) of +1.0&, the realfractionation factor associated with the precipitation ofdolomite, e26=24

1 , increases from �2.10& to �2.88& as thesize of the Mg-clay sink increases from 0% to 20% of thetotal. Thus, minor clay precipitation will lead to underesti-

mates of the fractionation factor associated with dolomiteprecipitation. Similarly, small amounts of dolomite precipi-tation will lead to underestimates of the fractionation factorassociated with the precipitation of Mg-clays.

If local exchange reactions move magnesium throughthe pore-fluid at a rate that is significantly greater thanthe modeled net reaction rates, modeled fractionation fac-tors will reflect the integrated isotopic effects of sourcesand sinks. In the simple case where one mineral precipitateswhile another dissolves, e26=24

model is a function of the fraction-ation factor associated with precipitation (e26=24

1 ), the d26Mgvalue of the magnesium source (d26Mgsource), the d26Mg va-lue of the pore-fluid (d26Mgpore-fluid), the net source/sinkterm from the model (Rmodel), and the gross sink term (P):

e26=24model ¼ 103 �

d26Mgpore-fluid

103 þ 1d26Mgsource

103 þ 1� P � Rmodel

2P � Rmodel

0@

þ e26=241

103þ 1

!� P2P � Rmodel

� 1

!ð4Þ

Depending on the difference between the isotopic com-position of the magnesium source and the pore-fluid, localexchange reactions can lead to large differences betweenmodeled and real magnesium isotope fractionation factors.For example, assuming a modeled fractionation factor(e26=24

model) of �2.5&, a d26Mg value of the pore-fluid of�0.8&, and gross precipitation rates that are two timesgreater than modeled net reaction rates (P > 2 � Rmodel),the real fractionation factor associated with precipitation(e26=24

1 ) increases from �2.5& to �3.3& to �5.6& as thed26Mg of the local source decreases from +1.7& to 0&

to �4.5& (Fig. 6).


6. CONCLUSIONS

The widespread observation that magnesium concen-trations almost always decrease down core implies Mg-mineral precipitation in the sediment column and/orunderlying basalt. The results presented here show thateven though concentrations of magnesium in the pore-fluid look similar in many deep-sea sediments, profilesof the d26Mg values of the pore-fluid may be very differ-ent. This observation is best explained by precipitation ofdifferent Mg-minerals in the sediment column and/orunderlying crust. In half of the studied sites (sites 1082,1086, and 1012), d26Mg values increase with depth byup to �2&, consistent with the precipitation of a Mg-mineral (dolomite) that is lower in d26Mg by �2.0& to�2.7&. Measurements of the d26Mg values of dolomitesamples from sites 1082 and 1012 yield fractionation fac-tors for dolomite precipitation that are almost indistin-guishable from those determined by the pore-fluidmodel. The correspondence between modeled fraction-ation factors and those determined from the dolomitesamples indicates that, at least at these sites, local ex-change of magnesium between sediment and pore-fluidcan be neglected and modeled magnesium isotope frac-tionation factors can be interpreted as reflecting fraction-ation factors for dolomite precipitation. At sites 1219,984, and 925, d26Mg values of the pore-fluid increasewith depth by variable amounts, from <0.2& at site925 to >2& at site 984. The increase in d26Mg valuesis consistent with a magnesium sink that is isotopicallyenriched relative to the pore-fluid by between 0& and+1.25&. Although the identity of this magnesium sinkis not precisely known, modeled fractionation factorscompare favorably to previous work on Mg-clays. Morework is needed to constrain the importance of local ex-change of magnesium between pore-fluid and sedimentsat these sites. Our results for all sites are consistent withprevious work on magnesium profiles in deep-sea sedi-ment pore-fluids that has focused largely on profile shapeand correlative changes in pore-fluid chemistry (e.g., Mgand Ca, alkalinity, K, etc.), mineralogical, and some iso-topic (e.g., d18O of H2O) observations, to identify themineral phases involved in magnesium cycling (Gieskes,1975; Lawrence et al., 1975; Gieskes and Lawrence,1981; Kelts and McKenzie, 1982). Our work demon-strates the utility of using measurements of magnesiumisotopes as a tracer of magnesium sources and sinks inlow-temperature aqueous environments and providesnew estimates of the magnitude of magnesium isotopicfractionation associated with the slow precipitation ofdolomite and Mg-clays in deep-sea sediments.

ACKNOWLEDGMENTS

J.H. received funding from the NSF graduate student fellow-ship. D.P.S. thanks support from H. Breck and W. Breck. Wewould like to thank C. Langmuir and S. Jacobsen for use of theiranalytical facilities. We would like to thank the ODP Program forsamples. This manuscript was substantially improved by reviewsfrom F. Richter, M. Fantle, and A. Galy.

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