Calibrated high-precision 17 O excess measurements using laser-current tuned cavity ring-down spectroscopy

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Introduction
Measurements of the stable isotope ratios of water are ubiquitous in studies of the earth's hydrological cycle and in paleoclimatic applications (Dansgaard, 1964;Dansgaard et al., 1982;Johnsen et al., 1995;Jouzel et al., 2007).Isotopic abundances are reported as deviations of a sample's isotopic ratio relative to that of a reference water, and expressed in the δ notation as: One important innovation was the development by Merlivat and Jouzel (1979) of a theoretical understanding of "deuterium excess": where δD is equivalent to δ 2 H.The parameter d excess is commonly used as a measure of kinetic fractionation processes such as the evaporation of water from the ocean surface.For example, d excess variations from ice cores have often been used to infer conditions at the moisture source from which polar precipitation is derived (Johnsen et al., 1989;Petit et al., 1991;Cuffey and Vimeux, 2001;Masson-Delmotte et al., 2005).The δ 18 O and δD isotopic ratios can be experimentally determined via a number of isotope ratio mass spectrometry (IRMS) techniques.For δ 18 O, equilibration with CO 2 has been the standard method for many decades (Cohn and Urey, 1938;McKinney et al., 1950;Epstein, 1953).For δD, reduction of water to H 2 over hot U (Bigeleisen et al., 1952;Vaughn et al., 1998) or Cr (Gehre et al., 1996) has typically been used.Simultaneous determination of δ 18 O and δD was made possible via the development of continuous-flow mass spectrometric techniques utilizing conversion of water to CO and H 2 in a pyrolysis furnace (Begley and Scrimgeour, 1997;Gehre et al., 2004).A recent innovation is the measurement of the difference between δ 18 O and δ 17 O at sufficiently high precision to determine very small deviations from equilibrium.In general, the nuclei mass difference of +1n 0 and +2n 0 implies that the fractionation factor for δ 17 O between two different phases will be approximately the square-root of the fractionation factor for δ 18 O (Urey, 1947;Craig, 1957;Mook, 2000): functions (Q), which for the oxygen isotope ratios is as follows (Matsuhisa et al., 1978): By analyzing a set of meteoric waters, Meijer and Li (1998) estimated the value of λ to be 0.528.Barkan and Luz (2005) used careful water equilibrium experiments to verify that the equilibrium value for λ is 0.529, in accordance with Matsuhisa et al. (1978), while Barkan and Luz (2007) showed that λ is 0.518 under purely diffusive (kinetic) conditions, also in good agreement with theory (Young et al., 2002).Thus, the Meijer and Li (1998) value of 0.528 for meteoric waters reflects the combination of equilibrium and diffusive processes in the hydrological cycle.
Based on these observations, Barkan and Luz (2007) defined the 17 O excess parameter as the deviation from the meteoric water line with slope 0.528 in ln(δ + 1) space: 17 O excess = ln(δ 17 O + 1) − 0.528 • ln(δ 18 O + 1) (5) Like d excess , 17 O excess is sensitive to kinetic fractionation, but unlike d excess , it is nearly insensitive to temperature and much less sensitive than δD and δ 18 O to equilibrium fractionation during transport and precipitation.Natural variations of 17 O excess are orders of magnitude smaller than variations in δ 18 O and δD and are typically expressed in per meg (×10 −6 ) rather than ‰ (×10 −3 ).The potential of 17 O excess in hydrological research is significant because it provides independent information that may be used to disentangle the competing effects of fractionation at the source, in transport, and in the formation and deposition of precipitation (Landais et al., 2008;Risi et al., 2010).It also has applications in atmospheric dynamics, because of the importance of supersaturation conditions that, during the formation of cloud ice crystals impart a strong isotope signature to water vapor (e.g., Blossey et al., 2010)  As a result the measurement of δ 17 O requires conversion of water to O 2 rather than equilibration with CO 2 or reduction to CO. Meijer and Li (1998) developed an electrolysis method using CuSO 4 .More recently, Baker et al. (2002) used a fluorination method to convert water to O 2 , which was analyzed by continuous flow IRMS; this approach was updated by Barkan and Luz (2005) for dual-inlet IRMS.
The dual-inlet IRMS method can provide high precision and high accuracy 17 O excess measurements.However, the technique is time consuming, resulting in significantly lower sample throughput when compared to the standard and relatively routine analysis of δ 18 O and δD.The fluorination procedure typically requires 45 min per sample, while the dual-inlet mass spectrometric analysis requires 2-3 h.In practice, multiple samples must be processed because of memory effects in the cobalt-fluoride reagent and other issues that can arise in the vacuum line (e.g.fractionation during gas transfer) (Barkan and Luz, 2005).Moreover, while this method provides the most precise available measurements of 17 O excess , measurements of individual δ 18 O ratios by this method are generally less precise than those obtained with more traditional approaches.
In recent years, laser absorption spectroscopy in the near and the mid-infrared regions has increasingly been used for isotope analysis.An overview of experimental schemes for different molecules and isotopologues can be found in Kerstel (2005).In the case of water, laser absorption spectroscopy constitutes an excellent alternative to mass spectrometry.The main advantage is the ability to perform essentially simultaneous measurements of the water isotopologues directly on a water vapor sample.
As a result, tedious sample preparation and conversion techniques are not necessary.
In this work we report on development of a new cavity ring-down laser absorption spectrometer that provides both high-precision and high-accuracy measurements of 17 O excess .This instrument is a custom modification of the Picarro, Inc. water isotopic analyzer, model L2130-i, a version of which has recently been made available as model L2140-i.Critical innovations include (1) the use of two lasers that measure absorption in two different infrared (IR) wavelength regions and (2) modifications to the spectroscopic measurement technique.We also developed a sample introduction system that permits the continuous introduction of a stable stream of water vapor from a small liquid water sample into the optical cavity.In combination with precise control of the temperature and pressure in the optical cavity of the instrument, data averaging over long integration times results in precision of better than 8 per meg in 17 O excess .This work can also be seen as a demonstration of state-of-the-art performance for laser absorption spectroscopy isotope ratio analysis for all four main isotopologues of water (H 16 2 O, H 17 2 O, H 18 2 O and HDO).We compare our results with high precision IRMS measurements and discuss the advantages as well as limitations of our approach.Introduction

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Reporting of water isotope ratios
Normalization to known standards is critical in the measurement of water isotope ratios.By convention, δ 18 O of a sample is relative to 18  = −55.5 ‰ is the value assigned by the International Atomic Energy Agency (Gonfiantini, 1978;Coplen, 1988).δD is normalized in the same manner, using δD assigned SLAP = −428 ‰.We normalize δ 17 O using: There is no IAEA-defined value for δ 17 O assigned SLAP , but Schoenemann et al. (2013) recommended that it be defined such that SLAP 17 O excess is precisely zero.We follow that recommendation here; that is, we define: ≈ −29.6986 ‰, well within the error of published measurements (Schoenemann et al., 2013;Lin et al., 2010;Barkan and Luz, 2005;Kusakabe and Matsuhisa, 2008).
Throughout this paper, all reported values of δ 18 O, δ 17 O, δD and 17 O excess have been normalized as described above, unless specifically noted otherwise.Subscripts are omitted except where needed for clarity.

17 O excess analysis with mass spectrometry
Isotope-ratio mass spectrometry (IRMS) measurements provide the benchmark for comparison with results from analysis of 17 O excess by CRDS.We used IRMS to establish accurate measurements of the 17 O excess , of four working laboratory standards and the IAEA reference water GISP, calibrated to VSMOW and SLAP.We also used both IRMS and CRDS measurements to determine the δD and δ 18 O of the same standards; δ 17 O is calculated from the 17 O excess and δ 18 O data.Table 1 reports the values, updated from those reported in Schoenemann et al. (2013).We used the method described in Schoenemann et al. (2013) to convert water to O 2 by fluorination, following procedures originally developed by Baker et al. (2002) and Barkan and Luz (2005).Two microliters of water are injected into a nickel column containing CoF 3 heated to 370 • C, converting H 2 O to O 2 , with HF and CoF 2 as byproducts.
The O 2 sample is collected in a stainless steel cold finger containing 5A molecular sieve following Abe (2008).To minimize memory effects, a minimum of 3 injections are made prior to collecting a final sample for measurements.
The O 2 sample is analyzed on a ThermoFinnigan MAT 253 dual-inlet mass spectrometer at mass/charge ratios (m/z) 32, 33, and 34 for δ 18 O and δ 17 O, using O 2 gas as a reference.Each mass spectrometric measurement comprises 90 sampleto-reference comparisons.Precise adjustment of both sample and reference gas signals (10 V ±100 mV) permits long-term averaging with no measurable drift, so that the analytical precision is given by simple counting statistics:   (Barkan and Luz, 2005;Schoenemann et al., 2013).External precision of the 17 O excess of repeated water samples ranges from 4 to 8 per meg (Schoenemann et al., 2013).

Instrument design
We use modified versions of a cavity ring-down spectroscopy (CRDS) analyzer designed for δ 18 O and δD, commercially available as Model L2130-i, manufactured by Picarro, Inc.The L2130-i is an update to the water-isotope analyzers originally discussed in Crosson (2008).Its uses an invar (Ni-Fe) optical cavity coupled to a nearinfrared laser.Optical resonance is achieved by piezoelectric modifications to the length of the cavity.When the intensity in the cavity reaches a predetermined value, the laser source is turned off and the intensity then decays exponentially.The time constant of this decay is the "ring-down time".The ring-down time depends on the reflectivity of the mirrors, the length of the cavity, the concentration of the gas being measured, and the absorption coefficient which is a function of frequency.The frequency is determined with a wavelength monitor constructed on the principle of a solid etalon (Crosson et al., 2006;Tan, 2008).we refer to as the L2130-i-C, we added a second laser that provides access to another wavelength region, centered on ∼ 7193 cm −1 where there are strong H 17 2 O and H 18 2 O absorption lines (Fig. 1).Rapid switching between the two lasers allows measurement of all three isotope ratios essentially simultaneously.About 200-400 ring-down measurements are made per second, and complete spectra covering all four isotopologues are acquired in ∼ 0.8 s intervals.
For isotope measurements with the L2130-i or L2130-i-C under normal operating conditions, water vapor in a dry air or N 2 carrier gas flows continuously through the cavity to maintain a cavity pressure of 50 ± 0.1 Torr at a temperature of 80 ± 0.01 • C, normally at a concentration of 2 × 10 4 ppm.The flow rate of 40 sccm −1 is maintained by two proportional valves in a feedback loop configuration up-and down-stream of the optical cavity.The spectral peak amplitudes are determined from the least-squares fit of discrete measurements of the absorption (calculated from measurements of the ring-down time) to a model of the continuous absorption spectrum.
The spectroscopic technique utilized for the acquisition and analysis of the spectral region relevant to the measurement of the isotopologues of interest is essentially the same as the one used in the earlier commercially available L2130-i analyzer.One of the main features of this technique is that optical resonance is obtained by dithering the length of the cavity.As discussed in Results (Sect.3), we found that drift on timescales longer than a few minutes limited the achievable precision of 17 O excess measurements to about 20 per meg; this drift is ascribed largely to small but detectable drift in the wavelength monitor.
To improve measurement precision, we developed an updated version of the L2130i-C, hereafter referred to as model L2140-i, which incorporates a different spectroscopic method.In the new method, the length of the optical cavity is kept constant during the acquisition a spectrum.Resonance is obtained by dithering of the laser frequency by means of laser current modulation.In this method, known as "laser-current tuning" the frequency for each ring-down measurement is determined directly from the resonance itself, based on the principle that resonance will occur only at frequencies spaced by Introduction

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Full integer multiples of the free spectral range (FSR) of the cavity.The FSR is inversely proportion to the (fixed) cavity length.The FSR under normal operating conditions is ≈ 0.02 cm −1 , and varies by no more than 10 −5 cm −1 owing to the precisely controlled temperature and pressure conditions.The wavelength monitor is still used for feedback to the laser frequency control electronics, thus allowing for rapid tuning to a frequency near the desired resonance, but the fine frequency spacing for a given narrow spectral region (e.g., that for H 17 2 O) is determined only by the FSR.In this way, each ring-down measurement can be unambiguously assigned to a stable and equidistant frequency axis.Additionally, the new scheme yield higher cavity excitation rates -typically 500 ring-downs per second.Further details on the laser-current tuning method are provided in a US patent application (Hoffnagle et al., 2013).

Integrated absorption
The use of laser-current turning permits greater accuracy in the determination of the width of spectral lines than was achievable with the L2130-i or L2130-i-C instruments.This allows us, with the L2140-i, to use the integrated absorption under the spectral lines, rather than the height of spectral peaks, to determine isotopologue abundances.The integrated absorption is given by: where κ(ν, T , P ) is the monochromatic absorption coefficient (cm 2 molecule −1 ) and u is the number density of absorbers (molecules cm −2 ).The integrated absorption is directly related to the absorption strength, S, via: where f is the line shape function due to Doppler and pressure spectral line broadening, T is temperature and P is pressure.The integral independent of pressure (e.g, Varghese and Hanson, 1984;Rothman et al., 1996).
The ratios A i /A j for two different absorbing isotopologues i and j -and therefore in principle the isotope ratios (δ 18 O, δ 17 O, etc.) -are also independent of pressure.This makes the integrated absorption superior to the spectral peak amplitude used in earliergeneration instruments.
Values of A are obtained by a least-squares fit of the measurements to an empirically-determined spectral model.The spectral model describes the measured absorption as the sum of a baseline and molecular absorption lines.Free parameters in the baseline are an offset, slope and quadatric curvature term.The molecular absorption spectrum is modeled as the superposition of Galatry profiles, which describe the shape function, f , for each spectral line.The Galatry profile, G is given by the real part of the Fourier transform of the correlation function, Φ (Galatry, 1961): where x is the frequency separation from the line center normalized by the Doppler width, y and z are collisional broadening and narrowing parameters, respectively, and τ is dimensionless time.
The parameters that determine the shape of the lines are obtained from spectra acquired by operating the analyzer in a fine-scan mode where ring-downs are acquired with a frequency spacing much smaller than the line width and using the wavelength monitor to determine the frequency axis.This determines the relationship between the collisional broadening and narrowing parameters, y and z, and the relationship between y for the "normal" water peak (H 16  2 O) and the values of y for each of the isotopologues.The Doppler width is a known function of temperature (e.g.Galatry, 1961) and is therefore a fixed parameter.This leaves three or four free parameters Introduction

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Full needed to describe absorption for unknown samples in each spectral region: one yparameter and one value for the integrated absorption, A, for each independent isotopologue spectral line of interest (e.g., one each for the H 18 2 O, H 16 2 O and HDO lines in the 7200 cm −1 wavenumber region).

Determination of isotope ratios
For the determination of δ where Values of δD are determined similarly using data from the first laser only:

Sample inlet system
We use two different inlet systems for the introduction of water into the CRDS optical cavity.To obtain measurements of the same water sample continuously over several hours, we use a "custom vaporizer".The custom vaporizer comprises a continuous flow inlet system similar to that described by Gkinis et al. (2010Gkinis et al. ( , 2011) ) and used previously for δ 18 O and δD.In this design, water is pumped continuously through a capillary and into a stainless steel "Tee" heated to 170 • C. In our application, the "pump" is a simple air pressure system, with a double needle that is used to puncture septum-sealed vials; air pressure introduced into the vial through a small steel tube pushes water through a fused-silica capillary and into the heated Tee.Within the Tee, the liquid water is mixed with dry air and exits the Tee as water vapor that is introduced into the CRDS optical cavity through an open split (Fig. 2).Water vapor concentrations as measured by the CRDS analyzer are maintained at a target value (normally 2 × 10 4 ppm) to within better than ±100 ppm.
For discrete injections of water into the CRDS we use a commercial vaporizer available from Picarro, Inc. as model A0211 and described by Gupta et al. (2009).The vaporizer, operating at 110 • C, mixes dry carrier gas with 1.8 µL of water, which is (LEAP Technologies LC PAL).In our experiments, a complete vial analysis consists of ten repeated injections from the same vial, for a total analysis time of about 1200 s.

Measurement precision and drift
We use the custom vaporizer to obtain CRDS analyses of the isotope ratios of the same water over several hours.The Allan variance statistic provides a convenient way to assess the analytical precision and drift for the resulting long integrations.The Allan variance is defined as (Werle, 2011): where τ m is the integration time and δ j +1 , δ j are the mean values (e.g.δ = δ 18 O or 17 O excess ) over neighboring time intervals.Here, we use the "Allan deviation" (σ Allan , square root of the Allan variance) which can be interpreted as an estimate of the achievable external precision as a function of integration time.
Figure 3 shows σ Allan for measurements made both with the L2130-i-C instrument using peak amplitudes, and with the L2140-i instrument using laser-current tuning and

Concentration sensitivity
Laser spectroscopy instruments used for water isotope measurements exhibit dependence of δ 18 O and δD values on the the water vapor concentration (Gkinis et al., 2010), and similar dependence is expected for δ 17 O and 17 O excess .This dependence arises primarily from the effect of pressure broadening on peak shape.As noted in Sect.2.3.2,use of the integrated absorption in place of peak amplitude in the calculation of isotope ratios with the L2140-i instrument should theoretically eliminate the water-vapor concentration dependence.We used the custom vaporizer to obtain measurements on the L2140-i over a wide range of water vapor concentrations.Figure 5 shows that there is a significant reduction in the sensitivity of isotope ratios to concentration when using the integrated absorption measurement, as expected.to that seen in the L2130-i and other earlier-generation instruments -to less than 0.04 ‰/1000 ppm.Sensitivity for δ 17 O is comparably reduced, from ≈ 0.4 ‰ to less than 0.08 ‰/1000 ppm.Finally, the sensitivity of 17 O excess to water vapor concentration is reduced from > 250 per meg to less than 30 per meg/1000 ppm.The concentration sensitivity of δD, at about 1 ‰/1000 ppm, however, is not significantly changed between earlier models and the L2140-i.

Calibration against VSMOW and SLAP
We performed two independent types of calibration experiments with the L2140-i.In the first experiment, we analyzed standard waters SLAP2 and VSMOW2, along with reference waters GISP, VW and WW, and used the two-point calibration line defined by Eqs. ( 7) and ( 8) to determine the value of the references waters treated as "unknowns".
(Note that the δ 18 O and δ 17 O of VSMOW2 and SLAP2 are indistinguishable from those of VSMOW and SLAP (Lin et al., 2010).)In the second experiment, we analyzed lab standards PW, VW, SW and WW and used the IRMS δ 18 O and δ 17 O values of PW and VW as calibration points.The resulting calibrated δ 18 O and δ 17 O values are then used to calculate 17 O excess , using Eq. ( 5).
In both types of calibration experiments, we used the commercial vaporizer and sets of four or five consecutive 2 mL vials, from which ten 1.8 µL injections were made.Two sets of each type of water were analyzed in each calibration experiment.The uncertainty of the measurements is given by the standard error of the mean values of repeated sets.To reduce the potential for memory effects influencing the results, we exclude data from the first two or three vials in each set of repeated waters.
The results of the calibration experiments are tabulated in Table 2. value of 28 ± 2 per meg (Schoenemann et al., 2013).Further, we find that both KD (Kona Deep), which is a deep ocean water sample, and VW, which is a meteoric water sample from the interior of East Antarctica (VW) have 17 O excess values indistinguishable from 0. The CRDS data thus provide independent support for defining 17 O excess of SLAP to equal that of VSMOW (≡ 0) following Schoenemann et al. (2013).The rootmean square difference between all the IRMS and CRDS data in our analyses is < 3 per meg.

Discussion
Our results demonstrate that analysis of 17 O excess using cavity ring-down laser absorption spectroscopy, as implemented in the L2140-i instrument, can be competitive with analyses by mass spectrometry.The precision of repeated individual measurements made over ≈ 30 min is better than 8 per meg, similar to the external precision reported for IRMS (e.g., Luz and Barkan, 2010;Schoenemann et al., 2013), and calibrated values of reference waters are indistinguishable between the two methods.
Achieving 17 O excess measurements at the < 10 per meg level with CRDS requires relatively long integration times when compared with the more common δ 18 O or δD measurements, which for typical applications require lower precision (≈< 0.1‰ and ≈ 1‰, respectively).Nevertheless, the new method is significantly less time consuming, less labor intensive, and safer than the IRMS method requiring the use of fluorination.Measurements of 17 O excess with a laser spectroscopy instrument with a different design (off-axis integrated cavity output spectroscopy, or OA-ICOS) were reported recently by Berman et al. (2013).It is beyond the scope of this paper to compare the ICOS and CRDS methods, but we note that the results reported by Berman et al. (2013) were obtained by frequent calibration of the instrument with standard waters to correct for drift.Measurements of the IAEA reference water GISP reported by Berman et al. (2013), when calibrated to VSMOW and SLAP, are somewhat lower than ours (23 ± 2 per meg, compared with our value of 27 ± 4), but compatible within 2σ of most 10208 Introduction

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Full reported IRMS values from the literature; the weighted average of previously-reported measurements (IRMS only) was 22 ± 11 per meg Schoenemann et al. (2013).The mean VSMOW-SLAP normalized value for GISP for all recent IRMS and CRDS measurements from four different laboratories is 17 O excess = 27±3 per meg (Schoenemann et al., 2013;Berman et al., 2013).
High precision 17 O excess measurements are achieved without drift correction on the L2140-i.Indeed, the precision and drift characteristics of the 17 O excess results are better than would be expected from the simple combination of noise in the δ 18 O and δ 17 O measurements, both of which show evidence of some drift in their Allan variances (cf.Fig. 3).The relationship between δ 18 O, δ 17 O and 17 O excess errors can be understood as a combination of correlated and uncorrelated noise contributions (Schoenemann et al., 2013): where σ xs is the precision of 17 O excess , σ xs is the precision of ln(δ 18 O+1), and ν 17 is the residual in ln(δ 17 O + 1) from a best-fit line through the data having slope m.
In general, the uncorrelated errors (ν 17 ) are small.At higher frequencies, m tends towards higher values.We find that for the 0.8 s averages of ≈ 400 individual ring-down measurements, the slope is 0.82 ± 0.02, or 1.0 ± 0.1 if a "model 2" regression that accounts for variance in both the δ 18 O and δ 17 O measurements is used (e.g.York, 1969).A slope of precisely 1.0 would be expected, for example, if all measurement error were due to noise in the H 16 2 O spectral line, since this measurement is shared equally in the calculation of both δ 18 O and δ 17 O.The noise in the high frequency data is indistinguishable from Gaussian, and is consequently reduced as a function of the square root of the integration time.For longer measurement times (integrations of 10 3 s or longer), m is ≈ 0.5, so that the term (m − 0.528)σ 18 is small.As for IRMS measurements, it is the combination of the very small magnitude of uncorrelated noise, ν 17 , combined with m ≈ 0.5, that leads to the very high precision for 17 O excess measurements, even where the δ We note that frequency-dependence of the error slope, m, is not observed in IRMS measurements.As discussed in Schoenemann et al. (2013), in both the H 2 O fluorination procedure and in the mass spectrometer source, likely sources of error will involve some combination of diffusive and equilibrium fractionation processes, both of which will lead to values of m close to 0.5 (e.g., Miller, 2002).That the relationship between δ 18 O and δ 17 O errors in the CRDS also tends towards m =≈ 0.5 at longer integration times suggests that low-frequency drift in these measurements is similarly attributable to fractionation effects, rather than, for example, drift in the optical cavity temperature or other aspects of the CRDS instrument itself.Fractionation of the δ 18 O and δ 17 O values could be associated with diffusion of water vapor, incomplete evaporation, or condensation and re-evaporation during the vaporization process, or possibly in the optical cavity.These observations suggest that the current practical limit of precision for isotope measurements on the L2140-i is set by the sample introduction system, rather than the CRDS analysis itself.As illustrated in Fig. 8, which compares IRMS and CRDS measurements, the magnitudes of the uncorrelated residuals (ν) are very small -similar to those obtained with high-precision IRMS -while the magnitude of σ 18 is much smaller than that obtained with IRMS measurements of O 2 prepared by fluorination.This was not the case with our original prototype instrument (L2130-i-C), for which analyzer noise was dominant even for long integration times (Fig. 8).We note that since the term (m − 0.528)σ 18 is very small, < 1 per meg, for vial-average measurements, changes to the sample introduction system that would significantly improve 17 O excess precision will be challenging.These comparisons attest to the significant improvement in the spectroscopic measurements achieved in the L2140-i, as well as to the stability of the water vapor delivery and minimal amount of fractionation occurring both in the commercial vaporizer and in our custom vaporizer design.
The L2140-i should be useful in a variety of applications, such as in the measurement of ambient water vapor concentrations in the atmosphere, currently done with laser spectroscopy instruments for δ 2009), or in the high-resolution analysis of ice core samples using in-line continuous melting systems (Gkinis et al., 2011).The low sensitivity to water-vapor concentration achieved with the integrated-absorption measurement would be an advantage in such applications, though we note that there is still some sensitivity that may become important for concentration variability greater than ≈ ±100 ppm.In the current commercial version of the L2140-i instrument, a water-concentration correction is available in the instrument software that uses a bilinear relationship of the form: where A(1) and A(2) refer to the integrated absorption for peaks 1 and 2 (Fig. 1), and a 0 and a 1 are empirically-determined coefficients.The coefficients are determined by varying the water concentration over a large range and then applied to each measurement.A similar correction is applied to A(3), A( 11), and A(13).A simple linear correction following instrument-specific empirical measurements such as illustrated in Fig. 5 could be used as an alternative.We note, however, that we have not evaluated the performance of the instrument at low water vapor concentrations (< 5000 ppm).
As an example of an application of the L2140-i, we performed a simple experiment in which 42 vials containing identical water, open to the air, were measured sequentially using 10 injections each.Because the vials were open to a relatively low humidity laboratory atmosphere, evaporation of the vials would be expected to raise the δ 18  of sample preparation time, and > 100 h of analysis time using the traditional fluorination and IRMS method.Note also that the progressive lowering of the 17 O excess value is clearly detectable from vial to vial at the 1-2 per meg level; this would probably not be possible to observe using the IRMS method.We suggest that the laser spectroscopy method for 17 O excess could be used in a number of hydrological and atmospheric sciences applications that were previously impractical.

Conclusions
Cavity ring-down laser spectroscopy (CRDS) is commonly used for measurements of the 18    Full Discussion Paper | Discussion Paper | Discussion Paper | λ ≈ 0.5 and the subscripts s and r refer to different phases or samples.For thermodynamic equilibrium, the value of λ is given theoretically by the ratio of the partition Discussion Paper | Discussion Paper | Discussion Paper | e (0.528 ln(−55.5×10−3 +1)) Discussion Paper | Discussion Paper | Discussion Paper | standard deviation of the individual sample/reference comparisons.

Determination of δ 18 O
and δD ratios on the Model L2130-i is obtained by measurements of the amplitude of H 18 2 O and H 16 2 O and HDO spectral lines from a laser operating in the area of 7200 cm −1 (wavelength ≈ 1400 nm).In a modified version, which Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper |

∞
0 f (ν, T , P )dν = 1 and S is 10201 Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | the value of R ref is an instrument-specific estimate of the ratio of integrated absorption of H 17 2 O or H 18 2 O to that of H Discussion Paper | Discussion Paper | Discussion Paper | injected through a septum.The resulting water vapor is introduced into the optical cavity via a three-way valve after a ∼ 60 s equilibration.Analysis of a single injection pulse takes approximately 120 s, excluding injection, vaporizer purging and equilibration time.Automated sampling from 2 mL vials is accomplished with an autosampler Discussion Paper | Discussion Paper | Discussion Paper | the integrated absorption measurement.In both cases, σ Allan values for 17 O excess of < 20 per meg are achieved after integration times of ≈ 5 × 10 2 s, and σ Allan values for δ 18 O, δ 17 O and δD are below 0.03, 0.03 and 0.04 ‰, respectively.However, these values represent the limits with the L2130-i-C; no additional improvements in precision were achieved with longer integrations times, and in general σ Allan begins to rise af-Discussion Paper | Discussion Paper | Discussion Paper | For δ 18 O, sensitivity is reduced from ≈ 0.2 ‰ for a 1000 ppm variation in H 2 O vapor concentration -comparable Discussion Paper | Discussion Paper | Discussion Paper | Figure 6 shows the calibrated mean values and uncertainties in 17 O excess for the two different types of calibration experiment.The results show that the 17 O excess values of the "unknowns" in each experiment with the CRDS are indistinguishable from the values previouslydetermined using IRMS.Note in particular that the CRDS value of the IAEA reference water, GISP (27 ± 4 per meg), calibrated independently, is nearly identical to the IRMS 10207 Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | 18 O and δ 17 O measurements are comparatively imprecise.Discussion Paper | Discussion Paper | Discussion Paper | 18 O and δD (e.g.,Noone et al., 2011;Sayres et al., Discussion Paper | Discussion Paper | Discussion Paper | O values through time, and the 17 O excess value should decrease; furthermore, the relationship between ln(δ 17 O + 1) and ln(δ 18 O + 1) would be expected to evolve along a slope intermediate between the equilibrium value (0.529) and the value for diffusion into dry air (0.518) (Barkan and Luz, 2007).These features are indeed observed in the experiment: 17 O excess decreases by ≈ 90 per meg and evolves along a slope of 0.523 (Fig. 9).The slope of ln(δ 17 O + 1)/ ln(δ 18 O + 1) is 0.5232 ± 0.0005, distinguishable at > 99% confidence from the "meteoric water line" slope of 0.528, as measured with reference water samples and IAEA water standards using the same instrument.A simple experiment like this, which was run fully-automated over ≈ 60 h, would take many hours 10211 Discussion Paper | Discussion Paper | Discussion Paper |

δ
17 O and δD with precision competitive with previous-generation instruments.Liquid samples are introduced into the optical cavity using an automated vaporization system that requires no prior sample preparation.Direct analysis of ambient water vapor in air is also possible.Calibration against the IAEA standard waters VSMOW and SLAP yields values for the reference water GISP of 27±4 per meg, indistinguishable from the value of 28 ± 2 obtained by Schoenemann et al. (2013) using isotope ratio mass spectrometry (IRMS).Our results establish CRDS measurements of 17 O excess as a viable alternative to conventional IRMS methods that require the use of fluorination to convert H 2 O samples to O 2 prior to analysis.Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper |
Fig. 6.Comparison of 17 O excess data from two independent sets of calibrations of reference waters measured by laser spectroscopy on the L2140-i (CRDS, open squares) with previouslydetermined values from mass spectrometry (IRMS, filled circles). 17O excess data are plotted vs. δ 18 O.Error bars (1σ) on the CRDS calibrated values are the standard error of the measurements (see Table 2).Values and error bars (1σ) on the IRMS values are from Table 1, updated from Schoenemann et al. (2013).The calibration points VSMOW, SLAP, PW and VW are shown as open circles for reference.
Fig. 6.Comparison of 17 O excess data from two independent sets of calibrations of reference waters measured by laser spectroscopy on the L2140-i (CRDS, open squares) with previouslydetermined values from mass spectrometry (IRMS, filled circles). 17O excess data are plotted vs. δ 18 O.Error bars (1σ) on the CRDS calibrated values are the standard error of the measurements (see Table 2).Values and error bars (1σ) on the IRMS values are from Table 1, updated from Schoenemann et al. (2013).The calibration points VSMOW, SLAP, PW and VW are shown as open circles for reference.

Fig. 7 .Fig. 8 .Fig. 9 .
Fig. 6.Comparison of 17 O excess data from two independent sets of calibrations of reference waters measured by laser spectroscopy on the L2140-i (CRDS, open squares) with previouslydetermined values from mass spectrometry (IRMS, filled circles). 17O excess data are plotted vs. δ 18 O.Error bars (1σ) on the CRDS calibrated values are the standard error of the measurements (see Table 2).Values and error bars (1σ) on the IRMS values are from Table 1, updated from Schoenemann et al. (2013).The calibration points VSMOW, SLAP, PW and VW are shown as open circles for reference.
Compared to the routine nature of δ18O and δD analysis, isotopic ratio measurements of 17 O, the second heavy isotope of oxygen in terms of natural abundance, are challenging.The greater abundance of 13 C than 17 O effectively precludes the measurement of δ 17 O in CO 2 equilibrated with water by IRMS at mass/charge ratio (m/z) 45.
Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | values for δ18O, δ 17 O and δD.In fact, triple isotope-ratio measurements of water have been presented in the past via the use of various laser sources utilizing different optical and data analysis techniques The resulting precision of repeated measurements of O 2 gas is 0.002 ‰, 0.004 ‰, and 0.0037 ‰ (3.7 per meg) for δ 17 O, δ 18 O, and 17 O excess respectively.Reproducibility of the δ 17 O and δ 18 O ratios of water samples is, in practice, less precise than these numbers indicate, because fractionation can occur during the fluorination process or during the collection of O 2 .However, because this fractionation closely follows the relationship ln(δ 17 O+1) = 0.528 ln(δ 18 O+1), the errors largely cancel in the calculation of 17 O excess O/ 16 O and D/H isotope ratios of water and water vapor, reported as δ 18 O and δD deviations from Vienna Standard Mean Ocean Water (VSMOW).

Table 1 .
Schoenemann et al. (2013)topic ratios of reference waters analyzed at the University of Washington "∆*IsoLab". 17excess values are from long-term average IRMS measurements, updated fromSchoenemann et al. (2013)to reflect the inclusions of additional data.δ 18 O and δD values are from long term average laser spectroscopy measurements.δ 17 O values are calculated from 17 O excess and δ 18 O (Eq. 5).Precision (±) is the standard error.n = sample size.Introduction

Table 2 .
VSMOW-SLAP normalized 17 O excess , δ 18 O, δ 17 O and δD values for reference waters determined by CRDS using (a) IAEA standards VSMOW2 and VSLAP2 as calibration points and (b) using University of Washington standards PW and VW as calibration points.IRMS-measured 17 O excess values are shown for comparison.Precision (±) is the standard error.n = sample size.
a VSMOW2 and SLAP2 calibration.b PW and VW calibration.Errors take into account uncertainty in calibration points.Introduction