Measurements of the 17O Excess in Water with the Equilibration

Anal. Chem. 2008, 80, 3244-3253

Measurements of the Equilibration Method

17

O Excess in Water with the

Joachim Elsig*,† and Markus Leuenberger†,‡

Climate and Environmental Physics, Physics Institute, University of Bern, Sidlerstrasse 5, CH-3012 Bern, Switzerland, and Oeschger Centre for Climate Change Research, University of Bern

The equilibration method is the present-day standard method for measuring δ18O in water samples. The massto-charge ratio of 45 is measured at the same time but generally not used for further analysis. We show that an improved equilibration method can be used for precise determination of δ17O in addition to that of δ18O, and therefore can estimate 17O excess values to a precision of better than 0.1‰. To control the masking effect of the 14 times more abundant 13C on mass 45, we propose to use a chemical buffer in the water samples to keep the pH value and therefore the fractionation during the equilibration process of the 13C constant. With this improved method, the precision for the δ18O value could also be slightly improved from 0.05 to 0.03‰. Furthermore, we discuss the influences of the amount of water, the temperature, the CO2 gas pressures, and changes in the pH during the measuring procedure on oxygen and carbon isotopes. We noticed that measured δ45 values are a good control for δ18O measurements. This study tries to fathom the possibilities and limitations of the equilibration method for measuring 17O excess values of water samples. The isotopic ratio 18R ) 18O/16O of H2O in water samples (s), commonly reported relative to an international reference (ref) standard in the delta notation as δ18O (eq 1), has become a standard tracer for environmental studies.1 18

δ18O ) 18

Rs

Rref

-1

(1)

Note that we report all delta values in per mill (‰) units, but omit the factor 103 in all equations. Thus a delta value of 0.002, as inserted in all equations, is reported as 2‰. The 17O isotopes were, for a long time, supposed to bear no additional information, because of the understanding that isotope effects depend on chemical and physical behavior and therefore on the mass differences among isotopic molecules. This relationship between the 17O and the 18O isotopes is expressed in an exponential form (eq 2) * To whom correspondence should be addressed. Phone: + +41 31 631 44 70. Fax: + +41 31 631 87 42. E-mail: [email protected]. † Physics Institute, University of Bern. ‡ Oeschger Centre for Climate Change Research, University of Bern. (1) Dansgaard, W. Geochim. Cosmochim. Acta 1954, 6, 241-260.

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17

R/17Rref ) (18R/18Rref)λ

(2)

where ref indicates the isotope ratio in a reference material and λ describes the mass-dependent fractionation. In a historical approach, λ was supposed to be 0.5.2 Further investigations have shown that this value differs from 0.5 on earth. In 1998, Meijer and Li found λ ) 0.5281 ( 0.0015 for water samples, and they concluded that this is also the most accurate value for other terrestrial materials.3 But different mass-dependent fractionations could have slightly different values for λ. Kinetic processes such as vapor-phase diffusion, for example, have a different λ value than equilibrium processes such as evaporation and condensation.4 Additionally, λ for a single process does not need to stay constant but can vary with environmental parameters. For example, λ for transpiration depends on the relative humidity.5 Another effect that plays a role and has to be considered in the isotopic budget is that of mass-independent fractionation processes, where values for λ are about 1 (λ ≈ 1). Those mass-independent fractionations occur in photochemical reactions among O2, CO2, and O3 in the stratosphere6 and may be transferred to water vapor and CO2 in the troposphere. To describe these small anomalies in the isotopic signature, one introduces a parameter called 17O - excess (∆17O) (eq 3) that describes the deviation of the measured δ17O value from the value predicted by the relationship given in eq 2 when taking the natural logarithm:7

∆17O ) ln(δ17O + 1) - λ ln(δ18O + 1)

(3)

Other formulas were proposed to characterize the deviation from mass-dependent fractionation processes,8 but for our use and accuracy it will not make any difference. Therefore we will work with the formula of eq 3 throughout this study. Because of the masking effect of the more abundant 13C value on mass 45, when measuring CO2 with a mass spectrometer (MS), the idea to measure δ17O and therefore also ∆17O values with the equilibration method has not yet been discussed. Hence, other (2) Craig, H. J. Geology 1954, 62, 115-149. (3) Meijer, H. A. J.; Li, W. J. Isot. Environ. Health Stud. 1998, 34, 349-369. (4) Barkan, E.; Luz, B. Rapid Commun. Mass Spectrom. 2007, 21, 2999-3005. (5) Landais, A.; Barkan, E.; Yakir, D.; Luz, B. Geochim. Cosmochim. Acta 2006, 70, 4105-4115. (6) Thiemens, M. H.; Jackson, T.; Mauersberger, K.; Schueler, B.; Morton, J. Geophys. Res. Lett. 1991, 18, 669-672. (7) Luz, B.; Barkan, E. Geochim. Cosmochim. Acta 2005, 69, 1099-1110. (8) Assonov, S. S.; Brenninkmeijer, C. A. M. Rapid Commun. Mass Spectrom. 2005, 19, 627-636. 10.1021/ac702436t CCC: $40.75

© 2008 American Chemical Society Published on Web 04/01/2008

methods were developed for measuring both of the oxygen water isotopes culminating in the water fluorination method of Barkan and Luz with a precision of 0.01 to 0.03‰ for both δ17O and δ18O as well as a sample size of about 2 µL.9 Facing these impressive numbers it does not seem to make sense to improve the equilibration method where the sample size is 1-5 mL and the precision is around 0.08‰. However, for certain application inconsistencies, estimating ∆17O values with the often used equilibration method leads to a valuable byproduct in addition to δ18O values. THE EQUILIBRATION METHOD The only way to measure water isotopes directly on water is the infrared laser spectrometry method, where the isotopic composition is derived from the absorption spectrum of the water in the gas phase.10 Since there is no way to measure water isotopes directly with a MS, the isotopic composition of the water needs to be transferred first to a gas that can easily be measured by MS instruments. In order to remove the gases from the water sample prior to equilibration, two methods are generally used: In older systems, the sample is frozen, the glass flask is evacuated, and the water is thawed and refrozen again (repeated procedure). In most commercial arrangements, the water sample is evacuated for a limited time through a capillary to minimize evaporation and possible fractionations. Which procedure is used is insignificant for the following discussion. After the evacuation of the glass flask with the water samples, CO2 is added. Due to chemical reactions of the CO2 molecules with the water molecules, the isotopic composition of the water is transferred to the CO2. The isotopic compositions of the water and the CO2 gas change until an isotopic equilibrium is reached. The mass balance for this reaction is given in eq 4 as i f i f aixRH + bixRCO ) afxRH + bfxRCO 2O 2 2O 2

(4)

and T is the temperature given in units of K.11 After the equilibration process, the equilibrated CO2 is analyzed for its isotopic composition with a dual inlet isotope ratio MS relative to a working CO2 gas. For CO2, one gets two ratios of ion beam currents, namely, the mass/charge ratios of 45/44 and 46/44. These two mass-to-charge ratios document different isotopic ion ratios, given in eqs 8 and 9. 45

R ) 13R + 217R

(8)

R ) 218R + 213R17R + 17R2

(9)

46

By taking the ratio of the corresponding mass/charge ratios of the sample and the reference gas, the influence of the ionization probabilities and the transmission and collection efficiencies of the molecules with masses 13, 17, and 18 on the isotope ratios is cancelled. To account for the cross contamination in the dual inlet MS12 and the sensitivity of the MS regarding different isotopic compositions, the measured δ values have to be corrected by a scaling factor r. This factor r can be lower or higher than unity depending on whether the scale is stretched or compressed. For the 18O measurements, the scaling factor r18 is estimated by at least two different standard waters in each series. The scaling factor for the 17O isotopes r17 does not need to be the same as that of r18. Using eqs 1, 8, and 9, one can write eq 5 as 17 i RH2O

)

1 ((r 45δ + 1)45RCO2,ref 2 17 i RCO2,equ)(b/a + R-λ) - b/a17RCO (10) 2

13

and eq 6 as 18 i RH2O

)

1 [(r 46δ + 1)46RCO2,ref 2 18 RCO2,equ((r1745δ + 1)45RCO2,ref - 13RCO2,equ) +

13

where x denotes the mass numbers 17 or 18, i stands for initial, f stands for final, R is the accordant isotopic ratio, and a and b are the moles of oxygen atoms in the water and in carbon dioxide molecules, respectively. Assuming that the mole amounts do not change during the equilibration and considering the fractionations that occur between H2O and CO2, one finds eqs 5 and 6 for the isotopic ratios of H2O related to the CO2 oxygen isotopic ratios as 17 i RH2O

f i ) 17RCO (b/a + 1/Rλ) - b/a17RCO 2 2

(5)

18 i RH2O

f i ) 18RCO (b/a + 1/R) - b/a18RCO 2 2

(6)

for which R (eq 7) is the temperature-dependent fractionation factor for 18O isotopes 18 f RCO2

R ) 18

f RH 2O

2

) exp(-20.6/T + 17.9942/T - 0.01997)

(7)

(9) Barkan, E.; Luz, B. Rapid Commun. Mass Spectrom. 2005, 19, 3737-3742.

1 ((r 45δ + 1)45RCO2,ref - 13RCO2,equ)2](b/a + 1/R) 4 17 i b/a18RCO (11) 2

where 45δ and 46δ are the raw measured δ values against the working gas. In order to solve eqs 10 and 11, we have to know the exact isotopic composition of the working gas, the 13C value of the CO2 molecule after the equilibration, the mole amounts (a, b) in the initial state, the temperature in the equilibration bath, and the mass-dependent fractionation parameter λ. Using eq 3, one can then derive ∆17O values. For the calculation, the 17R value of the reference gas has to be known. The absolute values for V-PDB and V-SMOW were taken from Assonov and Brenninkmeijer.13 For routine δ18O evaluations, eq 11 is not necessarily needed. In practice, eqs 8 and 9 are used and solved analytically (10) Kerstel, E. R. T.; van Trigt, R.; Dam, N.; Reuss, J.; Meijer, H. A. J. Anal. Chem. 1999, 71, 5297-5303. (11) Bottinga, Y. J. Phys. Chem. 1968, 72, 4338-&. (12) Meijer, H. A. J.; Neubert, R. E. M.; Visser, G. H. Int. J. Mass Spectrom. 2000, 198, 45-61. (13) Assonov, S. S.; Brenninkmeijer, C. A. M. Rapid Commun. Mass Spectrom. 2003, 17, 1017-1029.

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Figure 1. Effect of using the approximated formulas for calculating δ18O values for 5 mL water samples with different δ18O values (left panel) and for different amount of waters (right panel). In these calculations, the CO2 used for equilibration has the following isotopic composition: δ18O ) -13.07‰ V-PDB and δ13C ) -5.68‰ V-PDB.

by assuming λ to be 0.5. To see the order of magnitude of the influence of 13R and 17R values on the mass-to-charge ratio 46 and the influence of the amount of water, we approximate eq 11 with b/a ) 0 and 46R ) 218R instead of eq 9, resulting in eq 12.

where pCO2 is the pressure of CO2(g) in atm, K0 is the solubility in mol l-1 atm-1,15 T is the temperature in K, and pKi ) -log10(Ki). H2CO3 dissociates in water to HCO3- (eq 15) via K1

18 i RH2O

) 1/R(r1846δ + 1)18RCO2,ref ) m46δ + q

H2CO3 {\} H+ + HCO3-

(12)

(15)

and HCO3- to CO32- (eq 16) via The 18R value of the water sample could then be estimated by a linear equation. The unknown parameters m and q in eq 12 have to be determined by two waters with ideally large differences in their known δ18O values. The deviation from the true δ18O value when assuming b/a ) 0 is shown in Figure 1 for 5 mL volume water samples and a CO2 gas with the following isotopic composition: δ18O ) -13.07‰ V-PDB and δ13C ) -5.68‰ V-PDB. For smaller water samples, the deviations are getting larger (Figure 1) and one cannot neglect the ratio of the mole amounts b/a anymore, because the water is also changing its isotopic composition during the equilibration.14 The assumption that only 18O is contributing to the mass/charge 46 leads to shifted values in the δ18O values (Figure 1) and is not dependent on the amount of water. THE CARBON CHEMISTRY As soon as CO2 gets in contact with water, the reactions of eqs 13, 15, and 16 occur. K0

CO2(g) + H2O {\} CO2(aq) + H2O a H2CO3

(13)

CO2(g) is the carbon dioxide in the gaseous (g) phase, and CO2(aq) is the dissolved CO2 in water. H2CO3 is hardly abundant. But to keep track, we denote the amount of dissolved carbon dioxide by [H2CO3]. The equilibrium condition is described by Henry’s law:

K0 )

[H2CO3] with pK0 ) pCO2 -2622.38/T - 0.0178471T + 15.5873 (14)

(14) Meyer, H.; Schonicke, L.; Wand, U.; Hubberten, H. W.; Friedrichsen, H. Isot. Environ. Health Stud. 2000, 36, 133-149.

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K2

HCO3- {\} H+ + CO32-

(16)

where the equilibrium state is described by the dissociation constants15 given in eqs 17 and 18.

K1 )

[H+][HCO3-] [H2CO3]

with pK1 )

3404.71/T + 0.032786T - 14.8435 (17) K2 )

[H+][CO32-] [HCO3-]

with pK2 ) 2902.39/T + 0.02379T - 6.4980 (18)

The dissociation of water is described by eq 19:16

Kw ) [H+][OH-] with pKw ) 6013.6757/T + 23.6521 log10(T) - 64.7013 (19)

Actually, the dissociation constants, here given for freshwater, are not only a function of the temperature but also of the salinity. The salinity in water alters the distribution of the components of DIC (dissolved inorganic carbon, DIC ) [H2CO3] + [HCO3-] + [CO32-]) strongly and has an impact on the fractionation of the carbon isotopes. Since we are measuring distilled water samples as well as water from precipitation, we restrict the discussion to samples with salinity of zero and calculate with the formulas given above. (15) Harned, H. S.; Davis, R. J. Am. Chem. Soc. 1943, 65, 2030-2037. (16) Dickson, A. G.; Riley, J. P. Mar. Chem. 1979, 7, 89-99. (17) Mook, W. G. Neth. J. Sea Res. 1986, 20, 211-223.

Figure 2. Distribution of the different compounds of DIC subject to different pH values (left panel) and the resulting influence on the fractionation of the carbon isotope 13C in CO2 (right panel). ∆δ13C is defined as the difference between the equilibrated CO2 and the initial CO2 (∆δ13C ) δ13Ceq - δ13Ci). For pH values smaller than 4.3, ∆δ13C becomes rather stable. The values are calculated for a temperature of 20 °C, a CO2 filling pressure of 185 mbar, and 5 mL of water in a glass flask of 45 cm3, as used in routine measurements at our laboratory.

With eqs 14, 17, and 18, one can calculate the DIC content in the water resulting in the formula (eq 20):

(

DIC ) KHpCO2 1 +

K1 +

[H ]

+

K1K2 + 2

[H ]

)

(20)

The distribution of the different components of DIC is strongly dependent on the pH value of the water, as documented in Figure 2. The CO2 in the gas phase changes the isotopic composition depending on the distribution of the different carbonate species in the water. Because of the different fractionation coefficients for those carbon species, the isotopic composition of the CO2 gas is strongly dependent on the pH value of the water. The temperature-dependent fractionation factors are given in eqs 2123.17 13

13

RH2CO3/CO2(g) ) (-373/T + 0.19)/1000 + 1

(21)

RHCO3-/CO2(g) ) (9483/T - 23.89)/1000 + 1

(22)

13

RCO32-/CO2(g) ) (8616/T - 21.37)/1000 + 1

(23)

where 13Ru/v represents the fractionation factor of a compound u relative to a compound v. We can describe the changes in the carbon isotopes in the gas phase according to eq 24 (shown in Scheme 1) where pi is the initial and p(t) is the residual CO2 gas pressure at a time t. Granted that all of the carbon atoms are from

the embedded CO2 gas, one can calculate the final pressure of CO2 according to eq 25

pi(Vg - Vw) ) RT pf(Vg - Vw) K1K2 K1 + K0pf 1 + + + + 2 Vw (25) RT [H ] [H ]

(

)

where R is the universal gas constant, pf is the final CO2 pressure at equilibrium, Vg is the gas volume, and VW is the amount of water in a sample flask. With the formulas of eqs 3, 10, 11, and 24 we can theoretically estimate the influence of the pH value, the amount of water, the temperature, and the CO2 filling pressure on the precision of the 17O excess measurements and δ18O measurements with the equilibration method. For a normal routine measurement, we equilibrate 5 mL of water with about 7-8 mL STP of CO2 ((185 ( 10) mbar of CO2) admitted to a glass flask of 45 cm3 at a temperature of 20 °C. This temperature is held constant by a water bath into which the flasks are dipped. Assuming a reference water with a pH value of 4 (see below) and the conditions described above, we can calculate differences in the ∆17O and δ18O values for samples with slightly different end or start conditions relative to that of the reference water (Figure 3). From Figure 3 we conclude that the most influential factor for δ18O measurements is the temperature (0.19‰/°C). For measurements of δ18O values with a precision better than 0.05‰, the temperature has to be controlled in a range better than (0.26 °C. For the ∆17O values, the most critical factors are the pH value of the water and the amount of water (-0.43‰/

Scheme 1

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Figure 3. ∆∆17O and ∆δ18O defined as the difference between the water with different start or end conditions (amount of water, temperature, pH value, and filling pressure) relative to reference water with standard conditions (5 mL of water, 20 °C, pH 4, and 185 mbar CO2).

Figure 4. One sigma standard deviation for different pH values and 1 or 5 mL of water samples for δ18O and ∆17O, expected from theoretical considerations and values given in the text.

mL). For a pH value lower than 4, the dependency is negligible but becomes very large for a pH greater than or equal to 5. Different filling pressures between samples have no significant effects on the isotope measurements, because the fractionation is dependent on the amount of dissolved gas relative to the CO2 in the gas phase. This ratio, however, is constant for different filling pressures. ERROR ESTIMATE Additionally, the theory allows us to retrieve error estimates for the ∆17O and the δ18O values via error propagation. The temperature in the water bath is controlled to (0.2 °C during the measurement cycles. The amount of water in each glass is filled using a pipet with a precision of better than (0.02 mL. The CO2 starting pressure has no measurable influence on the ∆17O value 3248 Analytical Chemistry, Vol. 80, No. 9, May 1, 2008

and only a small influence on the δ18O value. The CO2 starting pressure generally changes by (10 mbar between samples. Assuming that the pH value of the water after the equilibration between the different samples can differ by (0.3 pH units, we can calculate an error of ∆17O for different pH values (Figure 4). From Figure 4 we conclude that if one wants to measure ∆17O values with a precision better than 0.1‰, one has to keep the pH values of the different waters below 4.3. Below this pH value, small differences in the final pH value between sample to sample or sample to reference water will not lead to significant impacts on the carbonate chemistry. The error for smaller amounts of water samples is smaller for the 17O excess but higher for δ18O measurements for a pH below 6. Since the major task of the equilibration method is to measure δ18O and the 17O excess comes

along as a valuable byproduct, the amount of water was chosen to be 5 mL. The error for the ∆17O is smaller for small amounts of water because it is mostly dependent on δ13C changes of the CO2. Less water results in a higher amount of CO2 in the glass flask and therefore to enhanced stability of pH values and fractionations. Furthermore, one can also conclude that the best possible precision for measuring ∆17O would be 0.04‰ for 5 mL of water and 0.01‰ for 1 mL at a pH value of 4 with the described equilibration method. Not included in this estimate are the internal precision of the MS and uncertainties associated to gas handling. The overall precision would therefore be higher than 0.04‰ and will be ascertained in the Experimental Section (see below). THE TIME-DEPENDENT EQUILIBRATION PROCESS IN THEORY Until now, we have only looked at the equilibrium condition. In this part we focus on the time dependency of the equilibration process. Instead of solving the system of differential equations of the reactions described in eqs 13 and 15, we search for the ratelimiting step. In the water-CO2 system, the CO2 transfer from the gas phase to the water is limiting since the hydration of the CO2(aq) is much faster (a few seconds) and accelerates with decreasing pH value.18,19 The flux between the gas phase and the water can be described with eq 26:

F ) k([DIC]s - [DIC])

DCO2 z

[cm‚s-1]

(27)

DCO2 is the diffusion coefficient of CO2 through water. Using the equilibrium condition of eq 25 and eq 26, we get the simplified differential equation for the CO2 gas pressure during the equilibration process (eq 28):

dpCO2 dt

) -kA

(

)

(p(t) - pf)(Vg - Vw) RT ) Vg - Vw RTVw -k

dRCO2 dt

)

RCO2(t) N(t)

[u(t)(R2 - 1) +

v(t)(R3 - 1)] -

1 k (R (t) - R1RH2O) (30) 3 CO2 CO2

(26)

where [DIC]s is the saturated concentration in equilibrium with the gas phase and k is the transfer coefficient. In a simple, empirical model one can interpret the flux as diffusion through a thin water film on the surface of the water with a thickness z. Then the transfer coefficient k can be interpreted as

k)

This documents that the equilibration time is also dependent on the geometry of the cylinder, namely, the ratio of the surface to the total volume of the cylinder. To estimate the equilibration time for the 17O and 18O molecules of the water, we have to consider that this equilibrium is reached only after the carbon isotopes are in equilibrium. This is due to the hydration pathways and the fact that a water molecule contains one oxygen atom whereas the carbon dioxide molecule contains two oxygen atoms. After the hydration of the CO2 molecule and the following reactions in the water, the newly formed and released CO2 molecule carries a mixed oxygen isotope signal with one-third of the value from the water signal (RH2O) and two-thirds of the value from the carbon dioxide signal (RCO2). Therefore, the reaction constant for the equilibration of the oxygen isotopes is one-third of the reaction constant for the CO2. Assuming that the isotopic composition of the water remains constant, this assumption is valid for a 5 mL volume of water (see Figure 1), and one can then derive a differential equation for the 18O isotope of the CO2 signal and similarly for the 17O isotope for the equilibration process (eq 30).

A (p(t) - pf) (28) Vw

A is the surface of the water, and for the definitions of the other variables, see above. This differential equation can be solved, and we obtain eq 29 for the time-dependent pressure in a cylindrical volume:

p(t) ) pf + (pi - pf) exp-A/VWkt ) pf + (pi - pf) exp-k/ht (29)

R1 (eq 7) is the fractionation factor for 18O isotopes between H2O and CO2(g), R2 is the fractionation factor20 between CO2(g) and CO2(aq), R3 is the fractionation factor21 between CO2(g) and HCO3-, u(t) is the time-dependent flux from CO2(g) to CO2(aq) derived from eq 29, and v(t) is the flux from CO2(g) to HCO3-. The temperature dependence of R2 and R3 is given by the following equations.

103 ln R2 ) 0.004107T - 2.2937

(31)

103 ln R3 ) 0.0002903T2 - 0.1872T + 22.253

(32)

In the next section we will use these formulas to compare the measured data with a model based on the equations derived in the previous sections. MEASUREMENTS AND RESULTS Reproducibility Test. In a first test we measured 60 replicates of the same water with our standard procedure: 5 mL of water equilibrated with ∼8 mL STP CO2 gas at 20 °C. Our home-built equilibration system allows us to run 16-20 samples automatically. The standard waters and the samples are distributed to three lines of eight connectors each. At the beginning of each line, Bern standard water is measured and used for calibration. The glass vessels with the water samples are shaken in the water bath for 5 h. After this time, shaking is stopped and the equilibrated CO2 is expanded sequentially to the MS. These repeatability tests indicate a precision of 0.05‰ for the δ18O value and a precision of 0.4‰ for the deduced ∆17O value.

where h is the height up to which the cylinder is filled with water. (18) Leuenberger, M.; Huber, C. Anal. Chem. 2002, 74, 4611-4617. (19) Kern, D. M. J. Chem. Educ. 1960, 37, 14-23.

(20) Vogel, J. C.; Grootes, P. M.; Mook, W. G. Z. Phys. 1970, 230, 225-238. (21) Beck, W. C.; Grossman, E. L.; Morse, J. W. Geochim. Cosmochim. Acta 2005, 69, 3493-3503.

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3249

Figure 5. Conductivity values and ∆17O of EPICA Dome C samples. The smooth lines represent the mean spline of Monte Carlo simulations based on the measured values and their assigned uncertainties. An anticorrelation is observed due to different buffer capacities of the sample and the reference water resulting in different pH values and therefore different fractionations in the 13C values.

Ice Measurements. With our standard procedure we also measured 41 different melted ice samples from the EDC9622 ice core in the age interval from 3000 to 16200 years B.P. Each sample was measured twice. The reason for measuring the ∆17O excess in the ice was to see whether there exists a signal in the ice affecting the δ13C measurements on CO2 from the air in ice cores. ∆17O values in meltwater deviating from zero would imply even larger signals in atmospheric CO2 due to the exchange of isotopes with the water. ∆17O signals in CO2 have been detected in stratospheric and mesospheric CO2.23,24 Fluxes between the stratosphere and the troposphere could affect the signature in tropospheric CO2.25 Existing ∆17O signals in the ice would have large impacts on δ13C gas measurements due to altered 17O correction procedures.26 The correction that would have to be (22) Augustin, L.; Barbante, C.; Barnes, P. R. F.; Barnola, J. M.; Bigler, M.; Castellano, E.; Cattani, O.; Chappellaz, J.; DahlJensen, D.; Delmonte, B.; Dreyfus, G.; Durand, G.; Falourd, S.; Fischer, H.; Fluckiger, J.; Hansson, M. E.; Huybrechts, P.; Jugie, R.; Johnsen, S. J.; Jouzel, J.; Kaufmann, P.; Kipfstuhl, J.; Lambert, F.; Lipenkov, V. Y.; Littot, G. V. C.; Longinelli, A.; Lorrain, R.; Maggi, V.; Masson-Delmotte, V.; Miller, H.; Mulvaney, R.; Oerlemans, J.; Oerter, H.; Orombelli, G.; Parrenin, F.; Peel, D. A.; Petit, J. R.; Raynaud, D.; Ritz, C.; Ruth, U.; Schwander, J.; Siegenthaler, U.; Souchez, R.; Stauffer, B.; Steffensen, J. P.; Stenni, B.; Stocker, T. F.; Tabacco, I. E.; Udisti, R.; van de Wal, R. S. W.; van den Broeke, M.; Weiss, J.; Wilhelms, F.; Winther, J. G.; Wolff, E. W.; Zucchelli, M.; Members, E. C. Nature 2004, 429, 623-628.

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made because of ∆17O values unequal to zero on the measured δ13C is described by eq 33. 17

δ13Ccorrected ) δ13C - 2 13

RV-PDB

RV-PDB

∆17OV-PDB )

δ13C - 0.07∆17OV-PDB (33) We found variations of ∆17O values in the range of (0.4‰ for melted EDC96 ice samples during the mentioned time interval (Figure 5). When comparing the values with the conductivity record27 measured on the same ice core, an anticorrelation between the measured ∆17O and the conductivity values is observed (Figure 5). Higher conductivity leads to a higher buffer capacity of the water. Therefore, the pH value of these waters is (23) Lammerzahl, P.; Rockmann, T.; Brenninkmeijer, C. A. M.; Krankowsky, D.; Mauersberger, K. Geophys. Res. Lett. 2002, 29, 1581-1584. (24) Kawagucci, S.; Tsunogai, U.; Kudo, S.; Nakagawa, F.; Honda, H.; Aoki, S.; Nakazawa, T.; Gamo, T. Anal. Chem. 2005, 77, 4509-4514. (25) Liang, M. C.; Blake, G. A.; Lewis, B. R.; Yung, Y. L. Proc. Natl. Acad. Sci. U.S.A. 2007, 104, 21-25. (26) Assonov, S. S.; Brenninkmeijer, C. A. M. Rapid Commun. Mass Spectrom. 2003, 17, 1007-1016. (27) Wolff, E.; Basile, I.; Petit, J. R.; Schwander, J. Ann. Glaciol. 1999, 29, 8993.

Figure 6. Time dependence of the equilibration process with (right panel) and without (left panel) shaking of the water samples. The smooth lines document the model results. Note that different waters were used, with an equilibrated δ18O value of -55.18‰ V-SMOW (left panel) and 25.2‰ for the isotopically enriched water (ew). Different CO2 filling pressures slightly shift the pH value of the water which leads to shifted ∆17O values.

higher than the pH values of the standard water, resulting in different fractionations of the 13C. Without pH control, this signal is then spuriously interpreted as a ∆17O signal. Equilibration Process. Usually, the water samples are shaken in a water bath for about 5 h until the isotopic equilibrium is reached. In an additional test, shaking of the samples was omitted and the CO2 was measured after 2, 13, 24, 35, 47, 57, 69, 80, 99, 117, 135, 159, 184, 216, 246, 278, 342, 392, 478, and 579 min (Figure 6). The δ18O value of the water used for this test was -55.18‰ V-SMOW. After about 100 min, the ∆17O value of the water did not change any more. The model can explain the observed offset of the ∆17O if one assumes that the final pH value of the standard water (pH ) 5.75) was slightly higher than the pH value of the sample waters (pH ) 5.53). Another explanation could be the influence of the scaling factors r17 and r18, which are normally derived from the equilibrium state and may therefore not be applicable for this transient situation. In addition, those scaling factors have a larger impact on the results for more negative δ values. In the case that the equilibrium is not reached, these factors are difficult to estimate. The thin water film on the surface of the water, describing the gas exchange in the model, would have a thickness of 2.5 µm (see eq 27) to fit the data best. The δ18O value is far from equilibrium even after 600 min. From theoretical considerations we concluded that the transfer coefficient for the δ18O value should be 3 times smaller than the transfer coefficient for the δ13C value. In fact, the equilibration for the oxygen isotopes without shaking is about 24 times slower than that for the carbon isotopes. This can be explained by assuming that the CO2 equilibrates with only a thin layer on the surface of the water and migrates downward into the water body only by diffusion. The pH of the surface layer is constant after 100 min and does not change anymore. The standard deviation for the corrected ∆17O values is 0.17‰. From this test we further conclude that the equilibration time needs to be the same for the 17O and the 18O isotopes, since the ∆17O value is constant after

100 min. This is only possible when the rate of change is the same for the 17O and the 18O isotopes. In a following test we produced an isotopically enriched water sample to examine the equilibration process toward more positive δ values. Therefore, we evaporated 3 L of water until only 130 mL was left. During evaporation, the salinity of the water was increased. Therefore, this water should have a larger buffer capacity against CO2 than that of the standard water used for calibration. Indeed, large negative ∆17O values relative to those of the standard water were measured (Figure 6). We repeated the measurement for the enriched water but with a nearly doubled CO2 filling pressure (320 mbar instead of 170.3 mbar) (Figure 6). In these experiments, the samples were continuously shaken in the water bath; hence, the expected transfer coefficient for the oxygen isotopes is indeed 3 times smaller than that for the carbon isotopes. Therefore, equilibrium is reached after about 200 min. The model parameters were 2.5 µm for the diffusion layer and a pH value of 5.2 for the standard water. For the sample water, the final pH values were 6.17 and 5.93 for the measurements with original and doubled CO2 filling pressures, respectively. The standard deviation for ∆17O of the last 10 measurements is 0.14‰ with doubled CO2 filling pressure and 0.22‰ for samples with a filling pressure of 170.3 mbar. This confirms that a more stabilized pH can improve the precision. It is noteworthy that the different filling pressures do not influence equilibration time. pH Control. On the basis of these results and our findings documented in Figure 3, we mixed a water buffered around pH 4. The standard mixture to use is 200 mL of 0.1 M potassium hydrogen phthalate (C8H5KO4) and 0.4 mL of 0.1 M hydrochloric acid (HCl). For arbitrary amounts of sample waters with volume V in milliliters we can use eq 34 to calculate the amount of 0.1 M HCl and eq 35 to calculate the amount of potassium hydrogen phthalate we have to add to the water sample to create water with a pH of 4. Analytical Chemistry, Vol. 80, No. 9, May 1, 2008

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VHCl ) 0.002V [mL]

(34)

mC8H5KO4 ) 0.020423V [g]

(35)

The 0.1 M HCl should be made with water which has an isotopic composition close to the composition of the water samples. One has to correct the measured δ18O value of the buffered water when measuring samples with much different isotopic signatures than those of the HCl solution. The δ18O value of the HCl solution can be estimated via a mixture with a water sample with known isotopic values. It would be sufficient to know the δ18O value of the HCl solution to better than (1‰ to ensure a correct δ18O value of the water sample. In contrast, even large values of ∆17O ((10‰) in the hydrochloric acid would not change the ∆17O value of the water samples by more than 0.02‰. In addition, the oxygen isotopes of the phthalate do not change the isotopic composition of the water, because they do not exchange with water isotopes. The ∆17O value, or more precisely the δ13C value, is in equilibrium after about 100 min (Figure 7). The model parameters were 2.5 µm for the diffusion layer and a pH value of 5.2 for the standard water. We did not buffer the standard water for this test, but only the sample waters. The standard deviation for the δ13C corrected samples is 0.12‰ and reduces to 0.08‰ for the last 10 samples of this sequence. The offset of the ∆17O value can be explained by differences in the final pH values. The standard deviation for the δ18O value, as soon as the equilibration is reached, is 0.03‰. In the future, one should test whether the results would be the same if one acidifies the water to a pH below 4 instead of using a buffer solution. This would minimize the amount of external water that has to be added to the sample and would remove possible hydration of the buffer (which is not expected anyway when applying a specifically selected buffer) and its potential isotope effects. In our test, however, we saw no effects of hydration when using potassium hydrogen phthalate. Precipitation Water. From the previous test, we conclude that one should not use a nonbuffered standard water to determine ∆17O values, since this leads to shifted values in the ∆17O values due to differences in the pH levels. Therefore, we mixed a water with pH 4 to be used as the new standard in the measuring process. In our lab we routinely measure water samples from the Swiss precipitation network. As an example, we show here the results of source water measured against the nonbuffered standard and against buffered standard water to pH 4. The water “Lucens” was measured two times in a routine sequence against the nonbuffered standard. From these measurements we derived the value ∆17O ) (-21.02 ( 0.8)‰. The water was then buffered to a pH of 4 and measured three times against the standard water with pH 4, which resulted in ∆17O ) (0.03 ( 0.02)‰. Here, the influence of the carbonate chemistry is clearly detectable. The large negative ∆17O value disappears as soon as one forces the pH to a value of 4. The nonbuffered standard water measured against the buffered standard water results in a value of ∆17O ) (-0.55 ( 0.13)‰. Interpreting this as a result of different pH values, would give a pH value of (5 ( 0.1) that is in agreement with the model assumption made in the previous tests. For laboratories performing δ18O measurements of natural water samples, it is important to know whether such large calculated ∆17O values, like those in the natural water “Lucens”, 3252

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Figure 7. Equilibration of a water sample with a pH of 4. The precision for the last 10 measurements is (0.08‰. In this test, the standard water was not buffered.

have an impact on the value of δ18O. As long as one uses the formulas of eqs 8 and 9 for the calculation of the δ18O value and not approximation formulas, the influence of the carbon isotopes is cancelled, since the mass 45 as well as mass 46 are affected. The situation is contrariwise as soon as one measures samples that deviate from mass-dependent behavior and as soon as one wants to estimate ∆17O values. When measuring water samples with ∆17O unequal to zero with the common equilibration method, the δ18O value has to be corrected according to the formula of eq 36.

δ18Ocorrected ) δ18Omeasured - 0.002∆17O

(36)

When measuring ∆17O values of water samples with the equilibration method, one has to know the δ13C value of the equilibrated CO2. This value can be derived from reference water measurements. However, this requires that the δ13C value has to be the same for the reference and the sample water, which makes a buffer necessary. When deriving δ18O and ∆17O values by assuming a wrong δ13C value, the ∆17O values are wrong and in addition the δ18O values also require slight correction. Critical samples that deviate from normal behavior relating to carbon fractionation or ∆17O values unequal to zero can easily be found as outliers from the fitted line through the measurements where the raw δ45 data are plotted against the raw δ46 data. Waters that are more strongly buffered against CO2 have δ13C values and consequently δ45 values that are shifted to more negative values. These waters can be clearly seen as outliers in a δ46-δ45 plot. Not every outlier is necessarily a result of the carbonate chemistry where the pH of the equilibrated water is different from that of the standard water. If the water is contaminated, for instance with drilling fluid in the case of ice cores, it would also lead to a large deviation from the fitted line through the raw data plot. Changes in the experimental setup and small differences between the lines

and the water samples during the evacuation process as well as the expansion of the CO2 gas from and to the water samples can also lead to fractionations in the δ13C value of the added CO2. Also, exceptional behavior of one of the Faraday cups and its electronics could be seen in this plot.

around a pH of 4, but in any case below a pH of 4.3. Waters that are not buffered, as those used in routine measurements for δ18O, should be checked for anomalous δ45 values due to fractionations of the carbonate isotopes by plotting the raw δ45 values against the raw δ46 values.

CONCLUSION From this study we conclude that the best way to measure ∆17O values with the equilibration method includes the following steps: The bracket with the water samples has to be shaken in a water bath during the whole measuring processsnot only during equilibration, but until the equilibrated CO2 is measured for all samples. This provides a more stable and well-homogenized temperature in the water bath for all samples. The samples should not be measured before an equilibrium of the carbon isotopes as well as the oxygen isotopes between water and carbon dioxide is reached. In our lab, an approved time for the equilibration is 300 min. The precision for ∆17O determinations with the equilibration method was found to be 0.4‰ for routine measurements and could be improved to better than 0.08‰ for a CO2 isotope determination of 10 min. The ultimate theoretical precision is 0.01‰. In order to improve precision, one has to consider the following precautions: To control the effects of different final pH values between the samples and standard waters, the waters should be buffered

ACKNOWLEDGMENT We thank T. Stocker for his continuous support of our MS laboratory. Many thanks to Peter Nyfeler and Ru¨diger Schanda for their help and work with the setup and the water measurements. We also thank Serge Bogni for cutting the ice samples. This work was supported by the Swiss National Science Foundation and Bundesamt fu¨r Umwelt BAFU. This work is a contribution to the European Project for Ice Coring in Antarctica (EPICA), a joint European Science Foundation/European Commission scientific program, funded by the EU (EPICA-MIS) and by national contributions from Belgium, Denmark, France, Germany, Italy, The Netherlands, Norway, Sweden, Switzerland, and the United Kingdom. Received for review November 28, 2007. Accepted February 6, 2008. AC702436T

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