doi: 10.1038/ngeo2069 carbon isotope ... - media.nature.com1 carbon isotope equilibration during...

1

Carbon isotope equilibration during sulphate-‐limited anaerobic oxidation of

methane

Marcos Y. Yoshinaga1*†, Thomas Holler2†, Tobias Goldhammer1, Gunter Wegener1,2,3*, John W.

Pohlman4, Benjamin Brunner2, Marcel M. M. Kuypers2, Kai-‐Uwe Hinrichs1, Marcus Elvert1

1. MARUM Center for Marine Environmental Sciences and Department of Geosciences, University of Bremen, D-‐28359

Bremen, Germany; 2. Max Planck Institute for Marine Microbiology, D-‐28359 Bremen, Germany; 3. Alfred Wegener

Institute for Marine Ecology and Technology, HGF-‐MPG Research Group for Deep Sea Ecology and Technology, D-‐27515

Bremerhaven, Germany; 4. U.S. Geological Survey, Woods Hole Coastal and Marine Science Center, Woods Hole,

Massachusetts 02543, USA.

* Email: [email protected]; gwegener@mpi-‐bremen.de; † these authors contributed equally to this work.

Supplementary Information Table of Contents

1. Culture Experiments

2. Detailed method description and evaluation of background methanogenesis

3. Compilation of global flux data

4. Isotope mass balance model

5. Supplementary References

Carbon isotope equilibration during sulphate-limited anaerobic oxidation of methane

SUPPLEMENTARY INFORMATIONDOI: 10.1038/NGEO2069

NATURE GEOSCIENCE | www.nature.com/naturegeoscience 1

© 2014 Macmillan Publishers Limited. All rights reserved.

http://www.nature.com/doifinder/10.1038/ngeo2069

2

1. Culture experiments

Sediment-‐free anaerobic oxidation of methane (AOM) enrichment cultures were retrieved by

continuous incubation of methane seep sediments obtained from Hydrate Ridge (offshore Oregon,

USA) and Amon Mud Volcano (Mediterranean Sea) with methane and sulphate as the only energy

sources (see ref. S1). Background methane production was evaluated in these AOM cultures

incubated with sulphate-‐free medium and both N2:CO2 and N2:CO2:H2 headspace. Resulting methane

production rates were compared with sulphide production values (equal to methane oxidation) at

AOM conditions (Eq. 1, Supplementary Fig. S1). Biotic (methane-‐free) and abiotic (incubation

medium without AOM culture) control incubations were performed (Supplementary Fig. S2) to verify

the carbon isotope effects examined in sulphate-‐limited AOM experiments (Fig. 2 and

Supplementary Fig. S3).

2. Detailed method description and evaluation of background methanogenesis

To determine background methane production in relation to AOM, cultures were washed 4

times with sulphate-‐free medium with pH set to 7.2 (via addition of NaOH). Aliquots of 10 mL that

contained similar cell density were transferred into 20 mL Hungate culture vials. Background

methanogenesis was evaluated by filling the headspace with 2 atm of N2:CO2 (90:10; all gases Air

Liquide; Düsseldorf, Germany). We also attempted to stimulate methanogenesis by adding a

hydrogen-‐enriched headspace of 2 atm N2:CO2:H2 (45:15:40). All samples were continuously shaken.

Every 3 to 4 days, 1 mL headspace gas was sampled with a gas tight glass syringe (gas volume was

replaced by N2:CO2) and methane concentrations were determined by gas chromatography coupled

to a flame ionization detector (GC-‐FID, Focus GC, Thermo). The detection limit for methane

concentration using this setup was 1 ppm. To ensure that the culture was metabolically capable of

AOM and to compare rates of methane production to methane oxidation, we performed parallel

incubations under AOM conditions (10 mM sulphate, 2 mM methane). Sulphide production, as proxy

for AOM activity, was determined spectrophotometrically via copper sulphide formationS2.


3

Supplementary Figure S1| Evaluation of background methanogenesis in sulphate-‐free incubations of AOM cultures from Hydrate Ridge (HR) and Amon Mud Volcano (AMV). Panels a and c, development of methane (CH4) concentrations without additional energy substrates (triangles) and with the addition of 0.8 mM H2 headspace (squares) during 27 days. Panels b and d, sulphide (HS–) production (dots) in parallel incubations at AOM conditions (10 mM sulphate; 2 mM CH4) over 29 days. Methane production (MP) and AOM rates were calculated from the gradient of monitored methane and sulphide concentrations. Dilution effects from sampling were considered. Shaded areas represent the magnitude of methane production in comparison with AOM.

To determine the carbon isotope effects of AOM at sulphate limitation, cell material was

transferred to artificial seawater medium with reduced sulphate concentration (1 mM) compared to

seawater (28 mM). Incubation medium pH was set to 7.2 via addition of NaOH. Cell suspensions were

transferred into 256 mL culture vials, sealed with butyl rubber stoppers and the headspace was

flushed with methane (purity 5.5, Air Liquide, Düsseldorf, Germany). Subsequently, vials were filled

completely with saturated methane medium to reach a starting concentration of 2 mM methane.

Samples were continuously shaken during the incubations. For geochemical measurements, 5 mL of

the medium were sampled via syringe (every 2−3 days) and the medium was concurrently replaced

with sulphate-‐, methane-‐ and carbonate-‐free medium to avoid gas loss into the headspace. The

dilution effect by medium replacement was taken into account in all calculations and displayed data.

All AOM-‐performing experiments were conducted in duplicates (Fig. 2 and Supplementary Fig. S3).

For methane and DIC measurements, sampled medium was transferred into two gas tight 2.2

mL exetainer® vials (Labco, Lampeter, United Kingdom) filled with zinc chloride solution (50 µL, 50%),

HR

time (d)

AOM = 0.11 mmol L–1 d–1

AOM = 0.30 mmol L–1 d–1

10 15 20 25 3050time (d)

0.000.010.02 MP = 0.0005 mmol L–1 d–1

MP = 0.00008 mmol L–1 d–1

MP = 0.0008 mmol L–1 d–1

MP = 0.00006 mmol L–1 d–1

CH4 (m

M)

10 15 20 25 3050

time (d)10 15 20 25 3050

0.03

time (d)

0.000.010.02

CH4 (m

M)

10 15 20 25 3050

0.03

2.04.06.08.010.0

H S– (m

M)

H S– (m

M)

0.0

2.04.06.08.010.0

0.0

a b

c dAMV

range of CH4 concentrations


4

and 1 mL of medium was replaced by helium. From this set of samples, methane concentrations

were determined using gas chromatography coupled to a flame ionization detector (GC 14; Shimadzu,

Kyoto, Japan) equipped with a Supel-‐Q Plot column, 30 m × 0.53 mm (Supelco, St. Louis, MO, USA).

The detection limit for methane concentration using this setup was 25 ppm. The elevated initial

methane concentrations might have been underestimated (e.g. Fig. 2a, d) due gas phase loss and/or

high room temperatures above 20 °C during preparation of headspace-‐free incubations. Methane

carbon composition was determined by coupled gas chromatography (G 1530A; Agilent Santa Clara,

CA, USA) to combustion (CuO-‐Ni-‐Pt catalyst; 940 °C) isotope ratio mass spectrometry (Thermo

Finnigan, Bremen, Germany), respectively. Analytical error for carbon (< 0.1‰) was checked against

an external methane standard of known isotopic composition. To determine DIC isotopic

composition of the incubation medium, medium was transferred into butyl rubber sealed glass vials,

gas phase was exchanged with helium and samples were acidified with phosphoric acid. The isotopic

composition of the released CO2 was measured via gas-‐bench coupled to isotope ratio mass

spectrometry (MAT 252, Finnigan, Bremen, Germany) and values were measured against external

standards (Solnhofener Plattenkalk), which was calibrated against NBS19 (IAEA, Vienna, Austria).

To determine the sulphate concentrations of the medium, sulphide was removed (via

precipitation with zinc chloride) and sulphate was measured from 50 times diluted medium using ion

chromatography (761 Compact IC, Ω Metrohm, Filderstadt, Germany). Sulphide concentrations were

determined from ZnCl2 fixed medium by colorimetric determination using the methylene blue

formation reactionS3 in a small assayS4.


5

Supplementary Figure S2| Biotic (methane-‐free) and abiotic (without AOM culture) controls for the sulphate-‐limited AOM experiments during 27 days. From left to right: Panels a and c, development of sulphate (SO4

2–) and sulphide (HS–) concentrations. Methane (CH4) concentrations remained below 5 µM throughout incubation. Panels b and d, development of δ13C values of dissolved inorganic carbon (DIC) in methane-‐free biotic controls for AOM cultures from Hydrate Ridge (HR) and Amon Mud Volcano (AMV). Panel e, development of CH4, SO4

2– and HS– concentrations in the abiotic control incubation. Error bars represent analytical imprecision of 10% for CH4 concentrations. Panel f, development of δ13C values of DIC and CH4 in the abiotic control.

a b

c d

(mM

)

0.0

0.5

1.0

1.5

2.0

2.5

-60

-58

-56

-24

-22

-20Abiotic control

(mM

)

0.0

0.5

1.0

1.5

2.0

-60

-58

-56

-24

-22

-20

(mM

)

0.0

0.5

1.0

1.5

2.0

-60

-58

-56

-24

-22

-20

e f

time (d)10 15 20 25 3050

time (d)10 15 20 25 3050

time (d)10 15 20 25 3050

time (d)10 15 20 25 3050

time (d)10 15 20 25 3050

time (d)10 15 20 25 3050

AMV, Biotic control

HR, Biotic control

AMV, Biotic control

HR, Biotic control

Abiotic control

!13C

(‰)

!13C

(‰)

!13C

(‰)

DIC

CH4

SO42–

HS–DIC

CH4

DICSO42–

HS–

SO42–

HS–


6

Supplementary Figure S3 | Experimental data (symbols) and model output (lines) for replicate AOM culture incubations Hydrate Ridge 2 (left column) and Amon Mud Volcano 2 (right column). From top to bottom: Panels a and d, development of measured concentrations of methane (CH4, circles), hydrogen sulphide (HS

–, closed triangles), sulphate (SO4

2–, open triangles), along with a first-‐order model approach (lines); DIC concentrations were derived from the modelled AOM reaction; methane concentrations are plotted with 10% error bars due to analytical imprecision. Panels b and e, development of δ13C values (in ‰) for measured DIC (black dots) and CH4 (red dots), along with modelled values by mass balancing (lines). Panels c and f, ratio of back flux (ƒ–) relative to AOM net reaction (ƒnet) as the main variable to achieve best fit for measured δ13CCH4 and calculated Gibbs free energies for AOM (kJ mol−1).

3. Compilation of global sulphate flux data

Location, diffusive flux values, stable carbon isotopic composition of methane and DIC and

references are given in Supplementary Table S1. Sedimentary diffusive fluxes (Jsed) of dissolved

sulphate and methane into the sulphate-‐methane transition zone (SMTZ) were calculated using a

modification of Fick’s first law of diffusionS5 (Eq. S1):

Equation S1

time (d)

(mM

)

Hydrate Ridge 2

time (d)10 15 20 25 30

0

0.5

1.0

1.5

2.0

5.0

6.0a

50

CH4

HS–

SO42–

DIC DIC

CH4

HS–

SO42–

Amon Mud Volcano 2

0

0.5

1.0

1.5

2.0

5.0

6.0

(mM

)

d

10 15 20 25 3050

time (d)

!40

!20

0

0

50

100

10 15 20 25 30

c

50

!Gr

ƒ – / ƒ

net (%

)

!40

!20

0

0

50

100!G

rf

time (d)10 15 20 25 3050

ƒ – / ƒ

net (%

)

b!20!22!24!26"13

C DIC

time (d)10 15 20 25 3050

!28!62!60!58!56

"13C CH

4

!64

e!20!22!24!26"13

C DIC

time (d)10 15 20 25 3050

!28!62!60!58!56

"13C C

H 4

!64


7

with Φ the sediment porosity and Dsed the sedimentary diffusion coefficient. The concentration

gradient ∂c/∂x is based on a linear regression of concentration data versus sediment depth directly

above the SMTZ. Dsed is calculated by correcting the seawater diffusion coefficient (Dsw) using

Boudreau’s lawS5 (Eq. S2):

Equation S2

We used a sulphate Dsw value of 5.72*10–6 cm2 s–1 (ref. S6) and estimated Φ to 0.66 when no porosity

value was available. Figure S4 provides additional examples for the prevalence of carbon isotope

equilibration during AOM in sediments from different oceanic provinces.

Supplementary Table S1. Data compilation of sulphate and methane fluxes (J in µmol cm–2 yr–1) into the SMTZ from various oceanic sediments. δ13C values (in ‰ VPDB) of methane and DIC within the SMTZ are reported and the ∆δ13CCH4 values are calculated as the offset in δ13C values of methane in the methanogenic zone (MGZ) and at the SMTZ (For details see main text and Fig. 1). na = not available and Refs. = references.

Site/Core information J SO4

2-‐ J

CH4 δ13C SMTZ

δ13C MGZ CH4

∆δ13CCH4 SMTZ Depth Refs.

Cascadia Margin DIC CH4 (cmbsf) Bullseye Vent 2002 (C-‐1) 3.7 -‐2.7 -‐46 -‐112 -‐80 32 478 S7*

Bullseye Vent 2002 (C-‐2) 4.5 -‐2.9 -‐46 -‐111 -‐73 38 330 S7

Bullseye Vent 2002 (C-‐3) 28.7 -‐15.4 -‐48 -‐73 -‐69 4 35 S7

Bullseye Vent 2002 (C-‐4) 4.8 -‐3.6 -‐45 -‐103 -‐75 29 386 S7

Bullseye Vent2002 (C-‐6) 7.9 -‐4.3 -‐47 -‐109 -‐74 35 250 S7

Bullseye Vent 2002 (C-‐7) 2.9 -‐3.0 -‐43 -‐106 -‐75 31 586 S7

Bullseye Vent 2008 (C-‐12) 22.9 -‐16.4 -‐42 -‐72 -‐62 9 170 S8

Bullseye Vent 2008 (C-‐5) 4.8 -‐3.0 -‐42 -‐101 -‐87 14 420 S8

Bullseye Vent 2008 (C-‐6) 2.9 -‐2.6 -‐40 -‐104 -‐76 28 520 S8

Amnesiac Vent (C-‐19) 13.0 -‐8.9 -‐51 -‐94 -‐78 16 230 S8

Barkley Canion (C-‐23) 18.1 -‐12.8 -‐21 -‐56 -‐43 13 110 S8

IODP 311 (U1329) 1.0 -‐0.5 -‐25 -‐105 -‐64 41 935 S9,S10

IODP 311 (U1327) 1.7 -‐1.2 -‐30 -‐99 -‐63 36 590 S10,S11

IODP 311 (U1325) 6.1 -‐1.5 -‐21 -‐102 -‐71 31 290 S10,S11

IODP 311 (U1326) 6.3 -‐2.2 -‐37 -‐94 -‐72 22 215 S10,S11

Porangahau Ridge, NZ

PC17 21.7 -‐18.9 -‐39 -‐98 -‐79 19 180 S12

PC7 11.5 -‐8.2 -‐21 -‐91 -‐79 12 210 S12

PC4 7.4 -‐5.5 -‐39 -‐108 -‐79 29 440 S12

Peru and Chile margins


8

Chilean Margin (C-‐10) 9.8 -‐7.1 -‐28 -‐101 -‐79 22 190 S13





ODP Leg 201 (1227A) 0.5 -‐0.2 -‐25 -‐84 -‐55 28 2,000 S14,S15

ODP Leg 201 (1229A) 0.2 0.0 -‐11 -‐89 -‐50 39 3,500 S14,S15

Gulf of Mexico

KC03-‐05 9.7 -‐9.7 -‐62 -‐107 -‐85 22 450 S16

KC03-‐07 10.5 -‐11.4 -‐56 -‐93 -‐78 15 450 S16

KC03-‐19 5.9 -‐6.2 -‐58 -‐104 -‐72 32 350 S16

MD02-‐2751 13.4 -‐9.2 -‐63 -‐90 -‐73 16 300 S17

GeoB10610 3.7 -‐7.0 -‐25 -‐73 -‐50 22 600 S18

Black Sea

E Black Sea (P824GC) 11.3 -‐4.6 -‐35 -‐97 -‐83 15 160 S19

NW Black Sea (core 214) 6.6 -‐3.2 na -‐96 -‐77 19 160 S20

Mud Volcanoes

(Bonjardim GeoB9051) 28.8 -‐30.6 na -‐64 -‐52 12 60 S21

(CAMV MSM1-‐205) 44.7 -‐8.4 -‐12 -‐53 -‐47 6 140 S21

GC38 NW 51.7 -‐11.3 na -‐60 -‐52 8 75 S22

GC46 S 34.8 -‐16.3 na -‐66 -‐58 8 90 S22

Blake Ridge

Kr-‐140-‐1 TC-‐3 35.4 -‐4.9 na -‐73 -‐68 5 53 S23

ODP Leg 164 (995) 0.9 -‐0.1 -‐36 -‐101 -‐74 27 2,190 S24,S25

ODP Leg 164 (997) 0.8 na -‐31 -‐102 -‐67 35 2,130 S24,S25

Shallow water Skagerrak, Denmark (Site 821GC) 36.1 -‐12.0 na -‐82 -‐76 6 70 S26 Eckernförde Bay (1993 Core) 39.9 -‐10.7 -‐2 -‐71 -‐63 8 30 S27 * Methane δ13C values are firstly reported in this study.


9

Supplementary Figure S4 | Sulphate flux as a proxy for sulphate availability at SMTZs, and its relationship with the prevalence of carbon isotope equilibration during AOM in sediments from different oceanic provinces. a, Scheme from exemplary sites under distinct regimes of sulphate fluxes into the SMTZ (Cascadia margin Bullseye Vent 2002 C-‐1 and C-‐3, respectively left and right panels). Within the limits of SMTZs, sulphate availability in the AOM zone (grey area) is dependent on the steepness of sulphate profiles (i.e. concentration gradients) from which diffusive fluxes are derived/calculated. In low flux, deep SMTZs, methane overlaps low concentrations of sulphate. Conversely, high flux, shallow SMTZs higher methane concentrations overlap relatively higher sulphate concentrations. b, Bullseye vent 2002 (C-‐6)S7. c, Core 19 from Keathley Canyon, Gulf of MexicoS16. d, Core 10 from the mid-‐Chilean marginS13. Depth profiles of sulphate and methane concentrations are shown in the left panels, with values for sulphate fluxes into the SMTZ (J SO4

2− in µmol cm−2 yr−1). Stable carbon isotope profiles of methane and DIC are shown in the right panels with corresponding Δδ13CCH4 (for details see Main Text and Supplementary Table S1).

4. Isotope mass balance model for the anaerobic oxidation of methane (AOM)

The simplified isotope mass-‐balance model describes the incremental development of δ13C

of dissolved inorganic carbon (DIC) and methane (CH4) for discrete time steps during the incubation

experiment. The model consists of two parts: an approximation of the experimental concentration

time series of CH4, SO42–, DIC and sulphide (HS–) together with a thermodynamic analysis of AOM,

and the actual isotope mass balance based on this concentration time series. The calculations are

stepwise, so that δ13C values for a given time increment t are calculated from the preceding time

0 5 10 2015 0 3 6 129 15Methane (mM)

100

200

300

400

500

600

700

0

Dep

th (c

m)

50

100

150

200

250

300

350

0

Dep

th (c

m)

c

J SO42- = 5.9 J SO4

2- = 9.8

!"C = 32‰

[CH4]

0 4 8 12

[SO42-]

0 10 20 30

Sedi

men

t dep

th

[CH4]

Methane (mM)0 4 8 12

[SO42-]

Sulphate (mM)0 10 20 30

Low !ux (deep SMTZ) High !ux (shallow SMTZ)

Methane (mM)

Sulphate (mM)

0 2 4 86 10Methane (mM)

3020100Sulphate (mM)

0-120 -90 -60 -30"13C (‰)

CH4

DIC

100

200

300

400

500

600

700

0

Dep

th (c

m)

b

d

J SO42- = 7.9

Methane (mM)

Sulphate (mM) "13C (‰)

a

SMTZ

SMTZSMTZ

3020100 0-120 -90 -60 -30Sulphate (mM) "13C (‰)

3020100 0-120 -90 -60 -30

CH4 !"C = 22‰ CH4

!"C = 35‰ CH4


10

increment t-‐1. The model was implemented in a spreadsheet that can be handled with an

appropriate computer program (LibO Calc®, MS Excel® or similar). This material is available upon

request with the corresponding authors.

Concentration time series and thermodynamics

To approximate the experimental concentration time series and to create an incremental

temporal domain for the isotope calculations, we used a first-‐order reaction model for CH4, SO42–,

bicarbonate (HCO3–; here equal to DIC) and HS–. The k value of the fit was adjusted to the

experimental SO42– data (Eq. S3), as we considered the SO4

2– data the most reliable.

Equation S3

with c(SO42–) the sulphate concentration (mM), k the first-‐order rate constant (s–1), and Δt the time

increment (s).

According to the AOM net reaction equation where substrates and products are equimolar (Eq. S4)

Equation S4

we obtain CH4, HCO3– and HS– concentration changes Δc for each time increment from the SO4

2– time

series (Eq. S5):

Equation S5

The Gibbs free energy for the net AOM reaction (Eq. S4) was calculated for each time step according

to Eq. S6

Equation S6

with the standard energy of reaction

Equation S7

and the activity of substrates and products (Eq. S8) in seawater (example CH4)

Equation S8

with R = 0.00831441 kJ mol−1 K−1, T = 285.15 K (HR) and 293.15 (AMV), respectively, and γ and ΔG0f

values listed in Supplementary Table S2.

[ ] [ ][ ] [ ]−

−−

⋅⋅

⋅⋅+Δ=Δ 244

30,, ln

SOCHHSHCORTGG AOMrAOMr


11

Supplementary Table S2. Activity coefficients γ and Gibbs energy of formation ΔG0

f for AOM reactants and products. Data from refs. S28 and S29. Species γ ΔG0

f kJ mol−1 CH4 (aq) 1 –34.39 SO4

2– 0.11 –744.6 HCO3

– 0.55 –586.8 HS– 0.55 12.05 H2O 1 –237.2

Isotope mass balance

The realisation of the isotope mass balance is technically a simplifying term of combined gain

and loss in the two relevant carbon pools, CH4 and HCO3–. The C transfers depend on the relative

proportion of the AOM reaction components, which we define as ƒ+ the forward and ƒ– the catalytic

backward flux (Supplementary Fig. S5). Here, ƒ– is the main variable to achieve the best fit of the

experimental data.

Supplementary Figure S5 | Carbon back flux in AOM illustrated by reaction vectors ƒ+ (forward flux), ƒ– (backward flux) and ƒnet (sum or net reaction) modified from ref. S30. Arrows lengths indicate rates and for convenience are not drawn with exact proportionality.

The observable net reaction ƒnet is the balance of forward and backward flux (Eq. S9) and by definition

set to 100 %.

!"#$ = ! !! ! ! Equation S9

The proportion r of backward and forward reaction is (Eq. S10)

! = " !" !

Equation S10

With the above derivations, we set up the δ13C mass balance for HCO3– (here abbreviated as DIC) and

CH4 (Eq. S11, S12)

CH4 HCO3–

ƒ–

ƒ+

ƒnet


12

!!"!"#! #$

!%"#! #$

=!!"!"#! #$"!

!%"#! #$"!

"$%"#! #$"!

"%"#! #$

%! !!" &

!!"!"#! #$"!

"!+ " ! ! !!"!!"# #$"!

"!"( )#$

%&

Equation S11

and

( )[ ]−−

+−−

−−

−⋅−−−

⋅−−

⋅=⋅

εδεδ

δδ

1,13

1,13

,1,

1,113

,13

11)( tDICttt

tDICttt

CrCr

cc

cCcC

CH4CH4CH4

CH4CH4CH4 ,,

Equation S12

with δ13C the stable carbon isotope value (‰PDB), c the concentration (mM), and ε the isotope

enrichment of forward (ε+) and backward flux (ε–). By convention, ε is calculated from experimental

isotope fractionation factors to

! = !" !"#""$$$%&' Equation S13

We adopted empirical isotope fractionations (α+) in Hydrate Ridge and Amon Mud Volcano cultures

(ref. S1; Supplementary Table S3) for the forward flux. The α– values that we employ in the mass

balance are intrinsic to the isotope effect of the AOM back flux and represent the sum of the isotope

fractionation of the forward flux and the pure physicochemical isotope effects of the CO2-‐HCO3– and

CO2-‐CH4 isotopic equilibria (refs. S31,32). For the calculation of the isotope signature at each time

point t, we solved Eq. S11 and S12 for δ13CDIC,t and δ13CCH4,t, respectively. Finally, the best fit was

achieved by adjusting ƒ– until δ13CCH4 data in the model and the experiment were congruent with a

deviation of ±0.1‰. We emphasize that the step-‐ or kink-‐like nature of the f– adjustment solely

results from the linear approximation of f–, and does not represent an analytical derivation from the

δ13C time series.


13

Supplementary Table S3. Parameterisation for isotope mass balance calculations. Starting conditions for concentration time series, isotope values and isotope fractionation factors employed in the isotope mass balance calculations. Values are given for the Hydrate Ridge (HR1, HR2) and Amon Mud Volcano (AMV1, AMV2) experiments. Parameter unit HR1 HR2 AMV1 AMV2 Reference

α+ 1.012 1.012 1.019 1.019 S1 α– 1.094 1.094 1.099 1.099 S31,S32 T °C 12 12 20 20 measured δ13CCH4 ‰ PDB –58.12 –58.02 –57.96 –58.07 measured δ13CDIC ‰ PDB –20.95 –20.53 –21.51 –21.23 measured c(SO4

2–) mM 1.03 1.02 1.02 1.02 measured c(CH4) mM 2 2 2 2 estimated c(HS–) mM 0.15 0.19 0.20 0.22 measured c(HCO3

–) mM 5 5 5 5 estimated kAOM s–1 1.15*10–6 1.15*10–6 1.25*10–6 1.25*10–6 estimated

Sensitivity analysis

We tested the sensitivity of our mass balance model against variations in the initialized CH4

concentrations and the carbon isotopic composition of DIC. First, we constructed a confidence belt

for the underlying CH4 concentration time series by calculating the AOM first order reaction with

minimum and maximum concentrations using all measured values plus their analytical error of 10%

(HR1: 1.5-‐2.2 mM CH4, AMV1: 1.6-‐2.4 mM CH4; Supporting Fig. S6a, c). Subsequently, the back flux

proportion f–/fnet was fitted considering four scenarios: I) minimum CH4 concentration, II) maximum

CH4 concentration, III) CH4 as measured at time zero, and IV) 2mM CH4 as used in the applied model

(Main Text Fig. 2a, d). All four scenarios started with identical f– values and continued with diverging

values during the later stages of the experiments (Supplementary Fig. S6b, d). While the first-‐order

CH4 consumption rate remained similar, the variable residual CH4 concentrations obtained from each

scenario produced different results towards the end of the experiments. In scenarios with CH4

concentrations below the model value of 2 mM, generally lower back flux adjustments were required,

whereas the opposite was necessary for concentrations above 2 mM. Similar results have been

obtained for the replicate experiments HR2 and AMV2. We emphasize that the marked shift from

initial f– values (HR1: 0.23, AMV1: 0.45) to their terminal range (HR1: 0.49–0.71, AMV1: 0.66–0.96) is

the most critical finding, rather than the range of the terminal values. This and the fact that


14

increasing f– values coincide with sulphate depletion below 0.5 mM strongly support that the

observed carbon isotope time series can only be explained by increasing carbon back flux during

sulphate-‐limited AOM.

Supplementary Figure S6 | Model sensitivity towards initialized methane concentrations in experiments HR1 (a, b) and AMV1 (c, d). Panels a and c: Confidence belt for CH4 concentration time series according to the first-‐order AOM reaction employed in the mass balance model for four scenarios (grey area defined by lower and upper limit of all measured values, red line: cCH4,0=2mM; dashed line: cCH4,0 as measured). Circles represent measured CH4 concentrations with a standard analytical error (± 10 %). Panels b and d: Corresponding best fit model results showing the percentage of back flux (f–) vs. net flux (fnet). The best fit model achieved identical values for f– at the beginning of each experiment, and diverging values in the later course of the experiments. The grey area represents the CH4 confidence belt examined in the above four scenarios.

Second, we tested our isotope mass balance model against changes in δ13CDIC that were

observed in the biotic controls (Fig. S2b, d). Applying an initial variability of ± 2 ‰ in the measured

δ13CDIC on the isotope mass balance, we found that such variability triggers a maximum deviation in

δ13CCH4 of ± 0.7 ‰ at the end of the experiments HR1 and AMV1. This effect could be buffered by an

CH4 (m

M)

Hydrate Ridge 1 Amon Mud Volcano 1

a

time (d)

ƒ– /

ƒ net (%

)

00 5 10 15 20 25 30

00 5 10 15 20 25 30

0

0.5

1

1.5

2

2.5

3

0 5 10 15 20 25 30

0

25

50

75

100

0 5 10 15 20 25 30

0.5

1

1.5

2

2.5

3

25

50

75

100

time (d) time (d)

time (d)

ƒ – / ƒ

net (%

)

c

b d

CH4 (m

M)


15

average increase or decrease in f– values of 1%, which is small in comparison to the overall change

from initial to terminal f– values (Supplementary Figure S6b, d). There is no experimental evidence

that these changes may result from the formation of solid carbonates and C-‐uptake into AOM

biomass was presumably low during the 30 days of incubationS33. Hence, we consider the impact of

δ13CDIC changes as of minor importance to our conclusion that increased carbon back flux explains the

observed carbon isotope reversal in δ13CCH4 values during the experiments.

5. References S1. Holler, T. et al. Substantial 13C/12C and D/H fractionation during anaerobic oxidation of methane

by marine consortia enriched in vitro. Environ. Microbiol. Rep. 1, 370–376 (2009).

S2. Cord-‐Ruwisch, R. A quick method for determination of dissolved and precipitated sulphides in cultures of sulfate-‐reducing bacteria. J. Microbiol. Methods 4, 33–36 (1985).

S3. Cline, J. D. Spectrophotometric determination of hydrogen sulfide in natural waters. Limnol. Oceanogr. 14, 454–458 (1969).

S4. Aeckersberg, F., Bak, F. & Widdel, F. Anaerobic oxidation of saturated hydrocarbons to CO2 by a new type of sulfate-‐reducing bacterium. Arch. Microbiol. 156: 5–14 (1991).

S5. Boudreau, B. P. Diagenetic models and their implementation: Modelling transport and reactions in aquatic sediments. (Springer, 1997).

S6. Schulz, H. D. & Zabel, M. Marine geochemistry. (Springer, 2nd edn, Heidelberg, 2006).

S7. Pohlman, J. W. et al. Anaerobic methane oxidation in low-‐organic content methane seep sediments. Geochim. Cosmochim. Acta 108, 184–201 (2013).

S8. Yoshinaga, M. Y. et al. Methane fluxes modulating the molecular and carbon isotopic composition of microbial lipids in gas hydrate bearing sediments from the northern Cascadia margin. Proceedings of the 25th International Meeting on Organic Geochemistry (Interlaken, Switzerland, 2011).

S9. Heuer, V. B., Pohlman, J. W., Torres, M. E., Elvert, M. & Hinrichs, K-‐U. The stable carbon isotope biogeochemistry of acetate and other dissolved carbon species in deep subseafloor sediments of the northern Cascadia Margin. Geochim. Cosmochim. Acta 73, 3323–3336 (2009).

S10. Torres, M. E. & Kastner, M. Data report: clues about carbon cycling in methane-‐bearing sediments using stable isotopes of the dissolved inorganic carbon, IODP Expedition 311. Proc. ODP, 311, doi:10.2204/iodp.proc.311.206.2009 (2009).

S11. Pohlman, J. W., Kaneko, M., Heuer, V. B., Coffin, R. B. & Whiticar, M. Methane sources and production in the northern Cascadia margin gas hydrate system. Earth Planet. Sci. Lett. 287, 504–512 (2009).

S12. Hamdan L. J. et al. Diversity and biogeochemical structuring of bacterial communities across the Porangahau ridge accretionary prism, New Zealand. FEMS Microbiol. Ecol. 77, 518–532 (2011).


16

S13. Hamdan L. J. et al. Geomicrobial characterization of gas hydrate-‐bearing sediments along the mid-‐Chilean margin. FEMS Microbiol. Ecol. 65, 15–30 (2008).

S14. Biddle, J. F. et al. Heterotrophic Archaea dominate sedimentary subsurface ecosystems off Peru. Proc. Natl Acad. Sci. USA 103, 3846–3851 (2006).

S15. Lipp, J. S. data from the ODP Leg 201 (personal communication).

S16. Pohlman, J. W., Ruppel, C., Hutchinson, D. R. & Coffin, R. B. Assessing sulfate reduction and methane cycling in a high salinity pore water system in the northern Gulf of Mexico. Mar. Petrol. Geol. 25, 942–951 (2008).

S17. Ussler III, W. & Paull, C. K. Rates of anaerobic oxidation of methane and authigenic carbonate mineralization in methane-‐rich deep-‐sea sediments inferred from models and geochemical profiles. Earth Planet. Sci. Lett. 266, 271–287 (2008).

S18. Schubotz, F., Lipp, J. S., Elvert, M. & Hinrichs, K-‐U. Stable isotopic compositions of intact polar lipids reveal complex carbon flow patterns among hydrocarbon degrading microbial communities at the Chapopote asphalt volcano. Geochim. Cosmochim. Acta 75, 4399–4415 (2011).

S19. Knab, N. J. et al. Regulation of anaerobic methane oxidation in sediments of the Black Sea. Biogeosciences 6, 1505–1518 (2009).

S20. Henkel, S. et al. Diagenetic barium cycling in Black Sea sediments – A case study for anoxic marine environments. Geochim. Cosmochim. Acta 88, 88–105 (2012).

S21. Nuzzo, M. et al. Shallow microbial recycling of deep-‐sourced carbon in Gulf of Cadiz mud volcanoes. Geomicrobiol. J. 25, 283–295 (2008).

S22. Nuzzo, M. et al. Impact of hot fluid advection on hydrocarbon gas production and seepage in mud volcano sediments of thick Cenozoic Deltas. Earth Planet. Sci. Lett. 341-‐344, 139-‐157 (2012).

S23. Borowski, W. S., Paull, C. K. & Ussler III, W. Carbon cycling within the upper methanogenic zone of continental rise sediments: an example from the methane-‐rich sediments overlying the Blake Ridge gas hydrate deposits. Mar. Chem. 57, 299–311 (1997).

S24. Borowski, W. S., Hoehler, T. M., Alperin, M. J., Rodriguez, N. M. & Paull, C. K. Significance of anaerobic methane oxidation in methane-‐rich sediments overlying the Blake Ridge gas hydrates. In: Paull, C. K., Matsumoto, R., Wallace, P. J. & Dillon, W. P. (Eds.), Gas Hydrate Sampling on the Blake Ridge and Carolina Rise. Proc. ODP, Sci. Res. (College Station, TX) 164, 87–99 (2000).

S25. Paull, C. K. et al. Isotopic composition of CH4, CO2 species, and sedimentary organic matter within samples from the Blake Ridge: gas source implications. In: Paull, C. K., Matsumoto, R., Wallace, P. J. & Dillon, W. P. (Eds.), Gas Hydrate Sampling on the Blake Ridge and Carolina Rise. Proc. ODP, Sci. Res. (College Station, TX) 164, 67–78 (2000).

S26. Parkes, R. J. et al. Biogeochemistry and biodiversity of methane cycling in subsurface marine sediments (Skagerrak, Denmark). Environ. Microbiol. 9, 1146–1161 (2007).

S27. Martens, C. S., Albert, D. B. & Alperin, M. J. Stable isotope tracing of anaerobic methane oxidation in the gassy sediments of Eckernförde Bay, German Baltic Sea. Am. J. Sci. 299, 589–610 (1999).

S28. Boetius A. et al. A marine microbial consortium apparently mediating anaerobic oxidation of methane. Nature 407, 623–626 (2000).


17

S29. Stumm, W. & Morgan, J. J. Aquatic chemistry: Chemical equilibria and rates in natural waters. (Wiley Interscience, 3rd edn, 1996).

S30. Holler, T. et al. Carbon and sulfur back flux during anaerobic microbial oxidation of methane and coupled sulfate reduction. Proc. Natl. Acad. Sci. USA 108, E1484–E1490 (2011).

S31. Mook, W. G., Bommerson, J. C. & Staverman, W. H. Carbon isotope fractionation between dissolved bicarbonate and gaseous carbon dioxide. Earth Planet. Sci. Lett. 22, 169–176 (1974).

S32. Ohmoto, H. & Rye, R. O. Isotopes of sulfur and carbon. In: Barnes, H. L. (Ed.) Geochemistry of hydrothermal deposits (Wiley, 2nd edn, New York, 1979). pp. 509–567.

S33. Wegener, G. Niemann, H., Elvert, M., Hinrichs, K.-‐U. Assimilation of methane and inorganic carbon by microbial communities mediating the anaerobic oxidation of methane. Environ. Microbiol. 10, 2287–2298 (2008).


doi: 10.1038/ngeo2069 carbon isotope ... - media.nature.com1 carbon isotope equilibration during...

Documents