Factors Affecting the Efficiency ofCarbon Monoxide Photoproductionin the St. Lawrence EstuarineSystem (Canada)Y O N G Z H A N G , † , ‡ H U I X I A N G X I E * , ‡ A N DG U O H U A C H E N †

College of Chemistry and Chemical Engineering,Ocean University of China, Qingdao, China, 266003;Institut des sciences de la mer,Universite du Quebec a Rimouski, Quebec, Canada G5L3A1

This study examined the effects of water temperature andthe origin (terrestrial vs marine) and light history ofchromophoric dissolved organic matter (CDOM) on theapparent quantum yields of carbon monoxide (CO)photoproduction for water samples collected along asalinity gradient (salinity range: 0-33) in the St. Lawrenceestuarine system (Canada). The solar insolation-weightedmean apparent quantum yield of CO (Φh CO) decreasedas much as fourfold with increasing salinity and showeda strong positive correlation with the dissolved organic carbon-specific absorption coefficient at 254 nm. This suggeststhat terrestrial CDOM is more efficient at photochemicallyproducing CO than is marine algae-derived CDOM andthat aromatic moieties are likely involved in this photoprocess.CDOM photobleaching, mainly at the very early stage,dramatically decreased Φh CO (by up to 6.4 times) for low-salinity samples, but photobleaching had little effect on themost marine sample. For a 20 °C increase in temperature,Φh CO increased by ∼70% for low-salinity samples and 30-40% for saline samples. This study demonstrates that watertemperature, as well as the CDOM’s origin and lighthistory, strongly affect the efficiency of CO photoproduction.These factors should be taken into account in modelingthe photochemical fluxes of CO and other related CDOMphotoproducts on varying spatiotemporal scales.

IntroductionCarbon monoxide (CO) in the surface ocean is primarilyproduced from the photolysis of chromophoric dissolvedorganic matter (CDOM) and is lost by microbial consumptionand outgassing (1-3). The primary motivation of early studiesof seawater CO arose from the observation that the oceanis a net source of atmospheric CO (4, 5), which regulates theoxidizing capacity of the atmosphere (6). Recently, interestin the marine CO cycle has expanded and diversified. Thestrong diel variation of the CO concentration in the surfaceocean (1), imposed by the diurnal fluctuation of the solarinsolation and modulated by microbial removal, outgassing,and vertical mixing, renders CO a suitable probe for upper-ocean mixing dynamics, photochemistry, optics, biology, and

air-sea gas exchange (7). As the second most abundantinorganic carbon-containing product of CDOM photochem-istry, CO is of significance to marine carbon cycling (8). COis also considered a useful proxy for general CDOM pho-toreactivity (9) and for the difficult-to-measure photopro-duction of dissolved inorganic carbon (DIC) (8, 10) andbiolabile carbon (11), which together have been proposed tobe one of the major terms in the ocean carbon cycle (12).Therefore, any significant advances or modifications in ourknowledge of oceanic CO would affect our view on othermajor marine biogeochemical cycles.

To quantitatively assess the role of CDOM photooxidationin the fate of organic carbon in the ocean (10, 13, 14), twoapproaches have been employed most frequently: in situincubations (15) and optical-photochemical coupled model-ing based on apparent quantum yield (AQY) (13, 16-18).The former determines water column photochemical fluxesby directly incubating water samples at varying depths inthe photic zone; it requires laborious fieldwork, but is thoughtto closely simulate the natural photochemistry and the insitu light field. The latter calculates photochemical rates bycombining experimentally determined AQY spectra withCDOM absorption coefficient spectra and underwater irra-diance. As CDOM absorption coefficients can be retrievedfrom satellite ocean color measurements (19), the modelingapproach appears promising for large-scale investigations(12, 17). The reliability of this approach depends, to a largeextent, on the reliability of the AQY spectra used in the model.Potentially large uncertainties in published AQY spectra arepartly associated with lack of quantitative knowledge of theinfluences of CDOM quality and environmental conditionson the related photoprocesses, including CO photoproduc-tion. This study determined CO AQY (ΦCO) spectra on watersamples from the estuary and Gulf of St. Lawrence (Canada)and evaluated the effects of water temperature as well as theCDOM’s origin (terrestrial vs marine) and light history onΦCO. The implications of these influences for the mechanismsof CO photoproduction and for modeling the photochemicalfluxes of CO and other related CDOM photoproducts arediscussed.

Experimental SectionSampling. Sampling stations were dispersed along a salinitygradient from the upstream limit of the St. Lawrence estuarynear Quebec City through the Gulf of St. Lawrence and tothe open Atlantic off Cabot Strait. Thirteen stations weresampled for absorbance and DOC measurements and six forthe AQY study (Figure 1). Water samples (2 m deep) weretaken in late July 2004 for Stations 1-12 and in mid-June2005 for Station 13 using 12-L Niskin bottles attached to aCTD rosette. Samples were gravity-filtered upon collectionthrough Pall AcroPak 1000 capsules sequentially containing0.8 µm and 0.2 µm polyethersulfone membrane filters. Thefiltered water was transferred in darkness into acid-cleaned,4 L clear glass bottles, stored in darkness at 4 °C, and broughtback to the laboratory at Rimouski. Samples were re-filteredwith 0.22 µm polycarbonate membranes (Millipore) im-mediately prior to irradiations, which were carried out within2 months of sample collection.

Photobleaching. In order to evaluate the effect of theCDOM’s light history on ΦCO (i.e., dose dependence), filteredsamples, placed in a clear glass container covered with aquartz plate, kept at 15 °C and continuously stirred, wereirradiated with a SUNTEST XLS+ solar simulator equippedwith a 1.5 kW xenon lamp. Radiations emitted from the xenonlamp were screened by a Suprax long band-pass cutoff filter

* Corresponding author phone: (418) 724-1767; fax: (418) 724-1842; e-mail: huixiang_xie@uqar.qc.ca.

† Ocean University of China.‡ Universite du Quebec.

Environ. Sci. Technol. 2006, 40, 7771-7777

10.1021/es0615268 CCC: $33.50 2006 American Chemical Society VOL. 40, NO. 24, 2006 / ENVIRONMENTAL SCIENCE & TECHNOLOGY 9 7771Published on Web 11/01/2006

to minimize radiations <290 nm, and the spectral composi-tion of the solar simulator closely matched that of naturalsunlight reaching the earth’s surface. The output of the lampwas adjusted to 765 W m-2 (280-800 nm) at the irradiationsurface as determined with an OL-754 UV-vis spectrora-diometer (Optronics Laboratories) fitted with an OL IS-2702 in. integrating sphere. Irradiation time varied from 20 minto 175.0 h to obtain various photobleaching regimes.

Irradiation for ΦCO Determination. The irradiation setupand procedure for determining ΦCO spectra were modifiedfrom Ziolkowski (20). Briefly, water samples were irradiatedin gastight quartz-windowed cylindrical cells in a temper-ature-controlled incubator using a SUNTEST CPS solarsimulator equipped with a 1-kW xenon lamp. Eight spectraltreatments were examined employing successive Schott longband-pass glass filters. Spectral irradiance under each filterwas measured with the OL-754 spectroradiometer. To assessthe effect of temperature on CO photoproduction, the originalsamples were irradiated at five temperatures: 0.5, 7.0, 15.0,24.0, and 32.0 °C. The dose dependence was evaluated onlyat 15 °C. A more detailed description of the irradiationexperiments can be found in the Supporting Information.

Analysis. CO was measured using a headspace methodfor gas extraction and a modified Trace Analytical TA3000reduction gas analyzer for CO quantification (21). Absorbancespectra were recorded from 200 to 800 nm at 1 nm incrementsusing a Perkin-Elmer lambda-35 dual beam UV-visiblespectrometer fitted with 10 cm quartz cells and referencedto Nanopure water. A baseline correction was applied bysubtracting the absorbance value averaged over an intervalof 5 nm around 685 nm from all the spectral values (22).Absorption coefficients (a) were calculated as 2.303 timesthe absorbance divided by the cell’s light path length inmeters. The lower detection limit of the absorption coefficientmeasurement was 0.03 m-1. This detection limit permittedmeasuring a at least up to ∼600 nm for Stations 1-10 in theestuary (a600: 0.051 -0.14 m-1) and up to ∼500 nm for Stations11-13 in the Gulf and Atlantic (a500: 0.054-0.084 m-1).Dissolved organic carbon (DOC) was measured using aShimadzu TOC-5050 carbon analyzer calibrated with potas-sium biphthalate. The coefficient of variation (c.v.) ontriplicate injections was <5%. Salinity was determined usinga PortaSal 8410A salinometer.

Calculation of ΦCO. The spectral CO apparent quantumyield, ΦCO(λ), was defined as the number of moles of COphotochemically produced per mole of photons absorbed

by CDOM at wavelength λ. A Matlab-coded iterative curve-fit method (20, 23) was employed to derive ΦCO(λ). Briefly,this method assumes an appropriate mathematical form withunknown parameters to express the change in ΦCO as afunction of wavelength. Decreasing exponential functionsare usually chosen for AQY spectra of CDOM photoprocesses(20, 23, 24). The amount of CO produced in an irradiationcell over the exposure time can then be predicted as theproduct of the assumed ΦCO(λ) function and the number ofphotons absorbed by CDOM integrated over the 250-600nm wavelength range, assuming no CO photoproductionabove 600 nm. We followed Hu et al.’s (25) recommendationsto calculate the number of photons absorbed by CDOM. Theoptimum values of the unknown parameters in the assumedΦCO(λ) function are obtained by varying these parametersfrom initial estimates until the minimum difference betweenthe measured and predicted CO production is achieved. Thefollowing quasi-exponential form was adopted to fit the data:

where m1, m2, and m3 are fitting parameters. This functionhas been demonstrated to perform generally better (18, 26),particularly in the UV-A and visible wavelengths, than themore frequently used single exponential form. To facilitateanalysis of local ΦCO variability, we defined a solar spectrum-weighted mean CO quantum yield (18), Φh CO, over the 280-600 nm range:

where Q(λ) is the noontime cloudless spectral solar photonflux recorded at Rimouski (48.453° N, 68.511° W), Quebec,on 24 May 2005. The rationale for this normalization is toreduce the ΦCO spectrum to a single value that accounts forboth the magnitudes and shapes of the ΦCO(λ) and Q(λ)spectra, thereby giving more weight to the wavelengths atwhich CO production is maximum (i.e., 320-340 nm). Froman environmental relevance perspective, Φh CO correspondsto the solar insolation-normalized CO production in the watercolumn in which all solar radiation over 280-600 nm isabsorbed by CDOM. Note that the Φh CO values presented hereare specific to this study since they more or less depend onthe specific Q(λ) spectrum used.

Results and DiscussionDOM Mixing Dynamics. Surface salinity (S) increased from0.004 at Station 1 to 32.55 at Station 13; the elevated salinitiesat Stations 7 and 9 were likely indicative of recently upwelledwaters (27) (Figure 5SI, Supporting Information). The DOCand absorption coefficient (a350) distributions nearly mirroredthe salinity distribution. The same was true with the DOC-specific UV absorption coefficient at 254 nm (SUVA254), anindicator of the aromatic carbon content of DOM (28), exceptat Stations 2 and 3, where the SUVA254 values were higherthan expected considering the salinity distribution trend(Figure 5SI, Supporting Information). The SUVA254 vs S curve(Figure 2) further confirms the presence of local DOM sourcesat these two sites, which are enriched with aromatic carbonrelative to the DOM transported from further upstream. SinceStations 2 and 3 were located in the maximum turbidity zoneof the upper St. Lawrence estuary, which is generatedprimarily by two-layer estuarine circulation (29), local DOMinputs could be from the dissolution of trapped particulateorganic materials and the injection of DOM into the watercolumn during sediment resuspension. The release of

FIGURE 1. Sampling locations in the St. Lawrence estuarine systemand in the Atlantic Ocean off Cabot Strait. Triangles represent stationsthat were sampled for the CO quantum yield study. All stationswere sampled for CDOM absorbance and DOC measurements.

ΦCO(λ) ) m1 * exp( m2

λ + m3) (1)

Φh CO )∫280

600Q(λ)ΦCO(λ)dλ

∫280

600Q(λ)dλ

(2)

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aromatic carbon-rich DOM from adjacent mudflats mightalso have contributed to the high SUVA254 values there,especially at Station 3. The a350 value decreased linearly withsalinity, but the slope of the line changes at Station 6 (Figure

2), slightly downstream from the mouth of the SaguenayFjord, where the topography changes abruptly from anaverage of ∼60 m in the upper estuary to >200 m in thelower estuary. This a350 distribution pattern agrees with thefinding by Nieke et al. (30). Tidal and wind-driven upwellingin and around the head of the lower estuary of CDOM-depleted deep water originating from the Atlantic Ocean (27)could be mainly responsible for this feature. Relatively moreintense in situ photobleaching in the lower estuary, asexpected from the longer residence times of surface watersthere, might also have played a role. However, the lineara350-S relationship across the entire lower estuary suggeststhe absence of significant photobleaching. The [DOC]-Srelationship resembles the a350-S relationship except atStation 3, where [DOC] is lower than inferred from the DOCmixing line, resulting in the elevated SUVA254 value at thisstation (Figure 2).

ΦCO of Terrestrial vs Marine CDOM. A compilation ofthe fit parameters for eq 1 is shown in Table 1. ΦCO spectrarepresentative of the upstream limit of the St. Lawrenceestuary (Station 1), the Gulf (Station 11), and the AtlanticOcean (Station 13) are displayed in Figure 3. Across the UV-visible regimes, the freshwater had the highest ΦCO values,the open-ocean water the lowest, and the Gulf waterintermediate. However, the differences between these spectraprogressively diminished with decreasing wavelength, apattern in accordance with previous ΦCO spectra determinedon water samples from widely varying geographic regions(Figure 3). These observations suggest the presence of

FIGURE 2. Plots of a350 (in m-1), [DOC] (in mg L-1), SUVA254 (in L (mgC)-1 m-1), and Φh CO vs salinity. The best fit of a350 vs salinity splitsinto two segments: salinity 0.0043-26.2 (y ) -0.116 + 5.02, R2 )0.995) and salinity 26.2-32.55 (y ) -0.267 + 8.86, R2 ) 0.975). Insetis the Φh CO vs SUVA254 plot and the best fit. The Φh CO values shownhere are those determined at 15 °C on the original (not pre-faded)samples. The 15 °C temperature was chosen since the mean ((s.d.)temperature of the sampled stations was 14.2 °C ((4.4 °C).

TABLE 1. Fit Parameters for Function ΦCO(λ) ) m1 × exp(m2/(λ + m3)) (eq 1 in Text)

temperature series photobleaching series

station no. T (°C) m1 m2 m3 f330a m1 m2 m3

1 0.5 3.46 × 10-11 6205.7 123.3 1.000 6.55 × 10-10 4036.1 39.317 3.81 × 10-11 6259.2 127.1 0.988 9.91 × 10-10 3295.1 -11.4815 6.55 × 10-10 4036.1 39.31 0.984 3.40 × 10-11 5281.2 56.6224 4.00 × 10-10 4894.5 90.37 0.751 2.10 × 10-11 5370.0 58.4532 1.03 × 10-10 6969.1 195.8 0.307 9.97 × 10-12 5401.2 45.44

0.169 5.88 × 10-12 5711.6 55.043 0.5 4.10 × 10-11 6600.9 151.9 1.000 5.42 × 10-11 6875.6 172.1

7 4.53 × 10-11 6714.2 160.2 0.990 3.72 × 10-11 6324.2 131.915 5.42 × 10-11 6875.6 172.1 0.953 1.51 × 10-10 4728.9 56.4524 6.74 × 10-11 7332.9 205.5 0.735 1.87 × 10-11 5800.1 88.1132 7.20 × 10-11 7928.2 249.3 0.385 8.64 × 10-12 5453.1 52.36

0.189 1.82 × 10-11 5019.5 47.488 0.5 1.12 × 10-09 3280.0 0.00107 1.000 2.01 × 10-09 3076.2 -6.73

7 9.27 × 10-11 5136.1 80.25 0.984 1.27 × 10-11 6403.7 116.115 2.01 × 10-09 3076.2 -6.73 0.981 1.13 × 10-11 6069.5 93.3924 6.20 × 10-11 5901.4 119.2 0.966 1.04 × 10-11 6234.2 104.232 1.14 × 10-09 3719.4 30.77 0.886 8.15 × 10-12 6154.5 96.88

0.680 6.35 × 10-12 5958.3 81.880.424 3.49 × 10-11 4554.7 33.250.177 6.49 × 10-10 2738.8 -39.19

11 0.5 1.83 × 10-08 1321.0 -133.87 3.71 × 10-09 2097.7 -83.5515 1.18 × 10-08 1716.2 -100.024 8.77 × 10-11 4854.9 53.9732 3.61 × 10-10 3801.4 4.68

12 0.5 3.43 × 10-11 5096.6 56.94 1.000 1.43 × 10-11 5863.2 81.827 2.54 × 10-11 5400.6 68.13 0.886 2.07 × 10-11 4943.5 31.93

15 1.43 × 10-11 5863.2 81.82 0.649 3.78 × 10-10 3084.7 -33.8324 1.28 × 10-11 6060.3 85.79 0.422 6.80 × 10-09 1599.9 -102.732 2.48 × 10-11 5830.3 87.11 0.287 1.09 × 10-08 1366.3 -119.9

13 0.5 8.38 × 10-11 3957 -2.73 1.000 5.94 × 10-11 4166.2 2.897 1.21 × 10-10 3545.8 -27.65 0.943 1.48 × 10-10 3920.2 2.72

15 5.94 × 10-11 4166.2 2.89 0.699 7.97 × 10-10 2920.9 -35.3124 3.40 × 10-11 4690 22.10 0.347 2.43 × 10-08 1339.4 -114.932 1.02 × 10-10 4569.1 36.88

a Fraction of original a330 values. The original a330 values (in m-1) are 6.33 for Station 1, 6.37 for Station 3, 2.6 for Station 8, 1.2 for Station 11,0.61 for Station 12, and 0.38 for Station 13.

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multiple CO precursors that were less selectively photolyzedby UV-B radiation than by UV-A and visible radiations. It isalso possible that metal ions (e.g., iron and copper), whichare known to promote photodegradation of CDOM, couldhave played a role in this phenomenon since the concentra-tions of these metal ions are usually higher in high-CDOMestuary waters than in oceanic waters. Our spectra for theGulf of St. Lawrence and the Atlantic Ocean almost perfectlymatch those for the Gulf of Maine (20) and the Pacific Ocean(2), respectively. Nevertheless, our ΦCO values for thefreshwater sample (Station 1) from the head of the St.Lawrence estuary are considerably lower, particularly in theUV-A and visible spectral regions, than those for the morecolored inland lake and river waters studied by Valentineand Zepp (16). This indicates that CDOM photoreactivitycan vary substantially among freshwater ecosystems, likelydue to differences in the quality of the CDOM. For example,SUVA350 (i.e., a350/[DOC]) for Valentine and Zepp’s samples(2.2 L (mg C)-1 m-1; 16) is on average 1.7 times that of oursample from Station 1.

Φh CO, as defined in eq 2, increased seaward initially (fromStation 1 to Station 3) but decreased monotonically withsalinity downstream of Station 3 (Figure 2). Since the watermass characteristics in the Gulf are typical of Case 1 watersof oceanic origin (30), the Φh CO-S relationship demonstratesthat marine algae-derived CDOM is less efficient thanterrestrial CDOM at producing CO photochemically. A linearregression reveals that Φh CO correlates well with SUVA254 (insetin Figure 2); the negative intercept suggests that not allaromatics are CO precursors. This Φh CO-SUVA254 correlationpoints to an important role of aromaticity in controlling ΦCO,which is in line with the study of Hubbard et al. (9)demonstrating that many specific aromatic compounds areefficient CO producers. As terrestrial DOM usually containsa greater fraction of aromatic carbon than does marine DOM(31, 32), the higher CO production efficiency of terrestrialDOM observed in the present study is likely a general featurefor aquatic environments. Mopper et al. (33) found thatincreasing salinity reduced the photoreactivity of a high-CDOM swamp sample, including CO photoproduction.However, as SUVA254 could account for 98% of the varianceof Φh CO (Figure 2), salinity was probably not a prevailingdeterminant of ΦCO in the St. Lawrence estuary, at least forthe season sampled.

Temperature Dependence. The Φh CO-temperature (T)relationship followed the linear Arrhenius behavior for allstations except Station 13 for which a concave Arrheniusplot is evident, showing relatively constant Φh CO valuesbetween 0.5 and 7 °C. (Figure 4). The mean activation energyfor Stations 1 and 3 in the upstream area (18.3 kJ mol-1) was78% higher than the mean activation energy for the rest ofthe stations (12.2 ( 0.8 kJ mol-1) in the lower estuary and theGulf. For a 20 °C increase in T, Φh CO increased by approximately70% for Stations 1 and 3 and by 30-40% for the other stations.These changes were relatively small compared to thedoubling-per-20 °C T dependence of hydrogen peroxidephotoproduction in Antarctic waters (34) and dimethylsulfidephotolysis in the Sargasso Sea (35), both of which are knownto be secondary photoreactions (i.e., photosensitized reac-tions or reactions of substrates with free radicals).

The T dependence of ΦCO demonstrates that secondaryphotoreactions were involved in the CO production, sup-porting the speculation by Hubbard et al. (9) that aromaticswithout the carbonyl group are the dominant CO precursorsbut contradicting the supposition that CO is primarilyproduced via the direct cleavage of DOM carbonyl groups,i.e., the Norrish type I mechanism (36, 37). However, COproduction through primary photoreactions could also existto a certain extent since simple carbonyl compounds, suchas formaldehyde and acetaldehyde, with photochemical (38)and perhaps biological (39) sources in natural waters, arewell-known to undergo direct photodecarbonylation in thesolar UV spectrum (40). The lower T dependence for thesaline samples suggests that primary reactions might be moreimportant for marine CDOM than for terrestrial CDOM, orthat there was a significant difference in the secondaryreaction mechanism between the two DOM pools. Since thephotoproduction of carbonyl compounds from DOM is alsocorrelated to its UV absorbance (38), an aromatics f carbonylcompounds f CO pathway, whether via primary or secondaryphotoreactions, agrees with the positive correlation betweenΦCO and aromaticity observed in the present study.

Dose Dependence. The dependence of ΦCO on CDOMphotobleaching is depicted as plots of Φh CO vs the fraction ofthe original a330 (Figure 5A) (note that a330 is selected because330 nm is near the peak CO production wavelength as inferredfrom Figure 3SI in the Supporting Information). The dosedependence varied widely among different samples and atdifferent stages of photobleaching. Φh CO for the upper- and

FIGURE 3. Comparison of ΦCO spectra for three representativestations in this study with previously published ΦCO spectra. TheΦCO spectrum for average freshwater is from ref 16, for the Gulf ofMaine from ref 20, and for average Pacific blue water from ref 2.The spectra from this study were those determined at 24 °C on theoriginal samples. The 24 °C temperature was chosen sinceirradiations for the previous studies were performed at room orlaboratory temperatures (refs 2, 20; the temperature was not reportedin ref 16).

FIGURE 4. Arrhenius plots of the solar insolation-weighted meanCO quantum yield, Φh CO. Lines are the best fits of the data. Linearregression was not performed for Station 13 since its Arrheniusplot is nonlinear.

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lower-estuary samples (Stations 1, 3, and 8) decreaseddramatically at first (within 5% a330 loss), continued to declinethereafter at gradually reducing rates, and eventually becamerelatively constant, a pattern resembling that of the dosedependence of photochemical O2 consumption in dilutedShark River water (13). Station 12 in the Gulf exhibited asimilar pattern, but the initial decrease in Φh CO was muchsmaller and occurred over a much broader bleaching range(over 22% a330 loss). Stn 13 outside the Gulf showed noconsistent dose dependence. These observations imply (1)that there appeared to be two distinct classes of COprecursors: one was very reactive, with its photolysis beingfaster than photobleaching, while the other was much lessreactive, with its photolysis being slower than photobleach-ing; (2) that terrestrially derived DOM contained a muchhigher proportion of the reactive class relative to marineDOM. The lack of dose dependence for the most marinesample could be due to the nature of the algae-derived CDOMor to the possibility that the sample from Station 13 hadalready been considerably photobleached in situ.

Unlike the strong linear Φh CO-SUVA254 correlation foundfor the original samples (Figure 2), the Φh CO-SUVA254 rela-tionship observed for the photobleaching study (Figure 5B)is nonlinear and resembles the pattern of Φh CO vs a330 (Figure5A). This suggests that the reactive CO precursors, as proposedabove, contained aromatic moieties but that the cleavage oftheir aromatic rings was not required for CO production (sinceSUVA254 did not decline very much during the initial rapiddrawdown of ΦCO). It also implies that the linear Φh CO-SUVA254

correlation observed for the original samples may not holdif CDOM is subjected to significant photobleaching in theenvironment.

Implication for Modeling. We are not aware of anypublished AQY-based modeling studies of photochemicalfluxes in natural waters (e.g., see refs, 2, 12, 13, 16-18, 20)that have adequately taken into account the effects of theCDOM source, CDOM light history, and water temperatureon the efficiencies of these photoprocesses. Results from thepresent study, however, point toward the necessity of adetailed mapping of ΦCO in relation to these variables inorder to more accurately estimate the photochemical COfluxes and evaluate the impact of photooxidation on thecycling of marine and terrestrial DOM. Clearly, if one usedfreshwater samples with little prior photobleaching, ΦCO

spectra would substantially overestimate the CO productionrates in the open oceans while severe underestimation wouldoccur if blue-water ΦCO spectra were applied to organic-richestuarine and coastal areas, where DOM undergoes littlephotochemical processing due to limited exposure to solarradiation. Ideally, experimentally determined AQY spectraonly apply to timescales over which the spectral compositionand amount of solar radiation absorbed by CDOM in thefield are equivalent to those absorbed by CDOM in theirradiation cells in the laboratory. Practically, it is almost notfeasible to apply this approach, since it is difficult to obtainreasonably accurate information on the light history of CDOMin the field on widely varying timescales. The rapid decreasein ΦCO with an increasing absorbed light dose during theinitial fading of the low-salinity samples indicates that usingunfaded freshwater ΦCO spectra would strongly exaggeratethe role of photochemistry in removing terrestrial DOM ifCO is used as a proxy for CO2 photoproduction. In this regard,ΦCO spectra determined on significantly faded samples shouldbe employed since the faded samples showed much lessdose dependence (Figure 5A) and the initial fading, whichled to the major drawdown of ΦCO, caused little DOC loss(<2%). In low- and mid-latitude open oceans, CDOM insurface waters is exposed to year-round solar irradiation andthe input of “fresh” CDOM is relatively slow, dependinglargely on the renewal of the surface waters (19). CDOMthere is expected to be significantly photobleached withrespect to CO production for most of the year. This mayexplain why ΦCO spectra for widely varying open-oceanregions and different seasons are similar in both magnitudeand shape (2). It therefore seems acceptable to neglect thedose dependence for low- and mid-latitude blue waters(possibly with the exception of upwelling areas where unfadedCDOM in the deep ocean is transported to the surface). Inhigh-latitude blue waters, however, the dose dependencemay not be trivial at the start of the spring season, when“fresh” CDOM from the preceding long, dark winter isexposed to prolonged solar irradiation. However, if theinsensitivity of CO production in our Atlantic water sample(Station 13) to photobleaching was a consequence of thenature of marine CDOM, then the effect of light history islikely inconsequential in blue waters regardless of locationsand seasons.

The dose dependence of CO photoproduction is expectedto occur on relatively short time scales and is principallyrestricted to interfaces (e.g., land-water interface, plant-water interface, coastal and upwelling zones, melting ice-water interface), where unexposed CDOM enters from shadedenvironments to unshaded ones (14). In contrast, the thermaleffect on this process is much milder but spatiotemporallymore extensive. In mid- and high-latitude inland, estuarine,and near-shore aquatic systems, seasonal T variations cansignificantly influence ΦCO. In the St. Lawrence estuary, thesummer-winter surface T difference is ∼22 °C at theupstream limit (Quebec City) and ∼18 °C in the Gulf, causingΦh CO to be 86% and 35% higher, respectively, in summer thanin winter. The effect of T seasonality in the open ocean, whichis maximum in the mid-latitudes (∼6 °C), is relatively small

FIGURE 5. Effect of pre-fading on the CO quantum yields as illustratedby the plots of (A) Φh CO vs the fraction of the original absorptioncoefficient at 330 nm and (B) Φh CO vs the specific absorptioncoefficient at 254 nm (Stations 1 and 3 only; SUVA254 data for therest of the stations are not available).

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(14%) if our Φh CO-T relationship for Station 13 is applied (theΦh CO-T data were interpolated to temperatures that were nottested). Latitudinally, however, ΦCO spectra determined attropical SST (annual mean T: 27.3 °C) would overestimateΦh CO by 42% in the 30°-45°N (S) zone (annual mean T: 16.5°C) and 60% in the 45-60°N (S) zone (annual mean T: 6.2°C).

The strong linear Φh CO-SUVA254 correlation for the originalsamples (Figure 2) suggests that, for aquatic systems, in whichCDOM is little photochemically processed due to self-shading, strong vertical mixing, and/or low solar insolation,SUVA254 may serve as a predictive tool for ΦCO. After takinginto account the T dependence, we derived the followingempirical equation for predicting the CO photoproductionefficiency in the St. Lawrence estuarine system:

where Φh CO is defined in eq 2, T is in Kelvin, and SUVA254 inL (mg C)-1 m-1. Statistically, SUVA254 and T can explain 96%of the variance of Φh CO (Figure 6SI, Supporting Information).

AcknowledgmentsWe thank A. Rochon for collecting the water sample fromStation 13 as well as our colleagues and the crew of the CoriolisII for their participation in the mission. C. Belzile providedconstructive comments on the manuscript. Y. Zhang wassupported by a Quebec-China Merit Fellowship. This workwas supported by grants from the Natural Sciences andEngineering Research Council of Canada (NSERC), and theCanada Foundation for Innovation (CFI). This is a contribu-tion to the research programs of the Institut des sciences dela mer de Rimouski.

Supporting Information Available(1) A more detailed description of the approach for ΦCO

determination; (2) the reproducibility of ΦCO determination;(3) the performance of the curve-fit method for deriving ΦCO

spectra; (4) absorption spectra of the original water samples;(5) spectral response curves of CO photoproduction; (6)spatial variations of salinity, DOC, a350, and SUVA254; (7) theperformance of eq 3 for predicting ΦCO; (8) sample transferfor CO analysis; and (9) the solar irradiance spectrum usedfor calculating Φh CO. This material is available free of chargevia the Internet at http://pubs.acs.org.

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ln(Φh CO × 106) )

- 1626.3T

+ 2.42 × ln(SUVA254) + 2.62 (3)

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Received for review June 27, 2006. Revised manuscript re-ceived September 6, 2006. Accepted September 15, 2006.

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