Decadal change of the surface water pco 2 in the North Pacific: A synthesis of 35 years of observations

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 111,, doi: /2005jc003074, 2006 Decadal change of the surface water pco 2 in the North Pacific: A synthesis of 35 years of observations Taro Takahashi, 1 Stewart C. Sutherland, 1 Richard A. Feely, 2 and Rik Wanninkhof 3 Received 26 May 2005; revised 25 January 2006; accepted 7 February 2006; published 7 July [1] Surface water pco 2 data observed over the 3 decades between 1970 and 2004 are analyzed for space and time (mean decadal) variability in thirty-two box areas over the North Pacific Ocean north of 10 N. During this period, the pco 2 values at SST increased at a mean decadal rate of 12.0 ± 4.8 matm decade 1 in all but four areas located in the vicinity of the Bering and Okhotsk Seas, where they decreased at a mean rate of 11.1 ± 5.7 matm decade 1. The mean rate of increase for the open ocean areas is indistinguishable from the mean atmospheric CO 2 increase rate of 15 matm decade 1 (or 1.5 ppm yr 1 ) suggesting that the North Pacific surface waters as a whole have been following the atmospheric CO 2 increase. However, the rate of increase varies geographically, reflecting differences in local oceanographic processes including lateral mixing of waters from marginal seas, upwelling of subsurface waters and biological activities. The decrease observed in the southern Bering Sea and the peripheries of the Okhotsk Sea may be accounted for by the combined effects of intensified biological production and changes in lateral and vertical mixing in these areas. The natural logarithm of wintertime pco 2 values normalized to a constant temperature and salinity of 14.3 C and 34.0 (the basin mean values, respectively) is correlated with winter SST. Using this relationship, the wintertime TCO 2 in mixed layer can be expressed as a function of winter SST with a standard error of ±5 mmol kg 1. Citation: Takahashi, T., S. C. Sutherland, R. A. Feely, and R. Wanninkhof (2006), Decadal change of the surface water pco 2 in the North Pacific: A synthesis of 35 years of observations, J. Geophys. Res., 111,, doi: /2005jc Introduction [2] The global oceans are one of the major dynamic reservoirs for CO 2 containing about 50 times as much CO 2 as the atmosphere, and continuing to absorb a large portion of the excess CO 2 emitted by anthropogenic activities [Battle et al., 2000; Keeling and Garcia, 2002; Quay et al., 2003; Sabine et al., 2004a; Sarmiento et al., 2000; Takahashi et al., 2002]. The mean annual rate of CO 2 uptake by the oceans for the past several decades has been estimated to be about 2 Pg-C yr 1. It is therefore important to know how the oceans have responded to the increased loading of CO 2 into the atmosphere. Because of the availability of research and commercial ships of opportunity, the North Pacific Ocean has become one of the most frequently sampled regions of the world oceans for the investigation of seasonal and interannual variability of CO 2 and nutrient chemistry [Landrum et al., 1996; Sabine et al., 2004b; 1 Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York, USA. 2 Pacific Marine Environmental Laboratory, NOAA, Seattle, Washington, USA. 3 Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, Florida, USA. Copyright 2006 by the American Geophysical Union /06/2005JC Takahashi et al., 1993; Wong and Chan, 1991; Wong et al., 2002; Zeng et al., 2002]. [3] The purpose of this paper is to estimate decadal mean rates of change in the partial pressure of CO 2 in surface mixed layer waters, (pco 2 )sw, over the North Pacific Ocean and their spatial variability using the measurements made by many investigators over the past 35 years. Since the time-space variability of (pco 2 ) SW over the equatorial Pacific has been investigated recently [Feely et al., 2006; Takahashi et al., 2003], this paper deals with the areas north of 10 N. In addition, we will show that the wintertime (pco 2 )sw is related to the wintertime mixed layer temperature (SST), and that it is also correlated tightly with the winter time total CO 2 concentrations (TCO 2 ) in the mixed layer. These relationships may be used for interpolating observations, as well as for testing ocean-atmosphere carbon cycle models. 2. CO 2 Partial Pressure in Seawater [4] The pco 2 in seawater is a vapor pressure of CO 2, and the difference between pco 2 in seawater and that in the overlying air is one of the important factors governing the CO 2 transfer flux across the sea-air interface. It is a sensitive function of temperature, doubling with every 16 C [Takahashi et al., 1993]. It is also a sensitive function of the total concentration of CO 2 (TCO 2 ), a sum of CO 2 species dissolved in seawater: TCO 2 = [CO 2 ]aq + 1of20

2 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 [HCO 3 ] + [CO 3 = ]. [CO 2 ]aq represents the concentration of free CO 2 molecules in aqueous media and is proportional to pco 2. TCO 2 in seawater depends on the net biological community production, the rate of upwelling of subsurface waters rich in CO 2, and the air-sea CO 2 flux. The sensitivity of pco 2 to changes in TCO 2 may be expressed in terms of the Revelle factor (= (@ ln pco 2 /@ ln TCO 2 ) T, S, Alk ), which varies from 8 in tropical waters with lower TCO 2 concentrations to 15 in polar waters with higher TCO 2 concentrations [Takahashi et al., 1993]. [5] In the surface mixed layer, the effect of seasonal warming on pco 2 is counteracted by lower TCO 2 caused by photosynthetic fixation of CO 2, as often seen during spring bloom periods; the effect of winter cooling is counteracted by increasing TCO 2 caused by upwelling of subsurface waters rich in respired CO 2. Consider the following example: If a parcel of polar ocean water at 1.9 C is warmed to an equatorial temperature of 30 C without change in TCO 2 and other chemicals, its pco 2 is increased by a factor of 4. On the other hand, deep waters upwelled in polar regions contain high nutrient concentrations, typically 35 mmol kg 1 nitrate. If the nitrate in this water is completely utilized by biological growth with the Redfield N/C ratio of 16/106, then the TCO 2 in the same water would decrease from 2150 mmol kg 1 to 1920 mmol kg 1.Asa result, pco 2 is decreased by a factor of 3 ((1920/2150) , using a Revelle factor of 10). The reduction is partially counteracted by the growth of CaCO 3 shell forming organisms, which is on the average about 20% of the organic carbon production. This lowers the alkalinity from 2300 to 2200 meq kg 1 and results in an increase of pco 2 by about 40%. Therefore, over the global oceans, the effect of change in temperature is roughly compensated by changes in TCO 2 and alkalinity, and the time-space variation in surface seawater pco 2 is dictated by competing effects of temperature, net biological production and the deep water upwelling. [6] Figure 1 shows the climatological mean distribution of (pco 2 )sw for non El Niño years over the North Pacific in February and August, normalized to a reference year of This is an updated version of the earlier distribution maps by Takahashi et al. [2002], and is based upon an improved database consisting of about 450,000 pco 2 measurements made over the tropical and North Pacific since the late 1960s. Since the atmospheric pco 2 was about 360 matm in 1995, the blue-magenta areas in Figure 1 are a sink for atmospheric CO 2, the green areas are nearly neutral and the yellow-orange areas are a CO 2 source. The central and eastern equatorial Pacific between 5 N and 15 S is a strong CO 2 source throughout a year due to upwelling of deep waters. The areas of the Kuroshio and its extension are a strong CO 2 sink during winter due primarily to cooling, and are a weak source during the summer due to warming. The western subarctic areas along the Kuril and western Aleutian arcs are a strong CO 2 source during winter due to convective mixing of deep waters rich in respired CO 2 and nutrients, and these areas become a strong sink in spring and early summer due to intense photosynthesis fueled by the nutrients that were supplied by the upwelling during the previous winter. In subtropical gyres, the primary cause for seasonal changes in pco 2 is seasonal temperature changes, whereas those in subpolar and polar waters are due primarily to TCO 2 changes caused by winter upwelling of deep waters and spring time plankton blooms [Takahashi et al., 2002]. Using the NCEP-42 year mean monthly wind speeds (at 10 m above the sea surface) [National Centers for Environmental Protection, 2004] and the wind speed squared dependence of sea-air gas transfer coefficient of Wanninkhof [1992], the pco 2 distributions shown in Figure 1 yield a mean annual net sea-to-air CO 2 flux of about 0.01 Pg-C yr 1 over the area north of 50 N and an air-to-sea net flux of 0.5 Pg-C yr 1 over the area between 14 N and 50 N. These flux estimates are uncertain by about ±35%, about one half of which is due to variability in (pco 2 ) SW, and the remainder is due to wind speed variability (1s) of±2ms Data Sources and Computational Procedures for the Decadal Change of pco 2 and SST [7] For an investigation of time-space variability of CO 2 chemistry, surface water pco 2 is chosen for the following reasons. First, a large number of measurements are available over the North Pacific for all seasons over the past 3.5 decades since Second, the measurements which have been made by various investigators using the airseawater equilibration method are compatible as a result of the use of common CO 2 gas mixture standards for calibrations at sea. Third, (pco 2 ) SW is a major property that governs sea-air CO 2 flux. Fourth, pco 2 in seawater is a sensitive function of the total CO 2 concentration dissolved in seawater because of the chemical amplification represented by the Revelle factor, so that small changes in TCO 2 can be detected Sources of the Data [8] The surface water pco 2 data and associated measurements (e.g., SST, salinity) used for this study have been obtained by the following research groups during the period from 1968 to 2004; Lamont-Doherty Earth Observatory (LDEO), Pacific Marine Environmental Laboratory (PMEL) of NOAA, Atlantic Oceanographic and Meteorological Laboratory (AOML) of NOAA; Institute of Ocean Sciences (IOS), Canada; and Japan Meteorological Center (JMC) and National Institute for Environmental Studies (NIES), Japan. The original data files are available at the Web sites of the respective groups. The total number of pco 2 measurements (north of 10 N) used in this study is about 327,000. Since (pco 2 ) SW computed using TCO 2, alkalinity and/or ph are not always consistent with the directly measured values, only those measured by equilibration methods are used in this study. The assembled data files and the time trend plots (annual and seasonal) are available at hhttp:// columbia.edu/co2i. The distribution of these data is shown in Figure Spatial Resolution [9] The spatial resolution used in this study is selected primarily on the basis of the time-space density of the observations. While smaller box areas would permit the resolution of narrow oceanographic features such as the Kuroshio and equatorial currents, they would reduce the number of observations made in a box during different seasons and hence fail to demonstrate seasonal variation clearly. Accordingly, we have chosen to bin the data into area boxes. Each box is defined according to its 2of20

3 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 1. Distribution of surface water pco 2 for February and August in a reference year This has been constructed using the method described by Takahashi et al. [2002], but is based on an updated database containing about 450,000 pco 2 measurements over the tropical and North Pacific (15 S 75 N), which has been extracted from a global database of 1.7 million pco 2 measurements. The yelloworange areas are a strong source for atmospheric CO 2, and the blue-magenta areas are a strong CO 2 sink. center point and contains ±5 latitude and longitude inclusive of 5.0 borders, with the exception of the six irregular size boxes which are located near the Aleutian and Kuril arcs along the northern edge of the Pacific. Four of these six boxes have rectangular shapes with a 10 E-W width but with varying N-S dimensions: the 55 N 165 W box has a N-S dimension of 54 N 60 N; the 45 N 165 W box has a N-S dimension of 40 N 54 N; the 55 N 175 E box has a N-S dimension of 52 N 60 N; and the 45 N 175 E box has a N-S dimension of 40 N 52 N. Two adjacent boxes 3of20

4 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 2. Numbers of the observations for surface water pco 2 in box areas used in this study over the North Pacific Ocean. K after numbers indicates a multiplier of There are about 327,000 pco 2 measurements in the areas north of 10 N. that straddle the Aleutian arc have a slanted border, so that the 55 N 175 W box (with its southern border slanted at an angle of about 40 ) contains only the data from the Bering Sea; and the 45 N 175 W box (with its northern border slanted at an angle of about 40 to fit the box located north of it) contains only the North Pacific data south of the Aleutian arc. Since we have no measurements inside the Okhotsk Sea, the boxes located along the Kuril chain contain only the Pacific data, although some of them reflect the outflow from the Okhotsk Sea Seasonal Correction and Annual Mean [10] In many of the box areas, measurements are available only in some months in a year, and fewer are available especially for pre-1997 periods. In order to estimate the decadal mean rate of change in (pco 2 ) SW based upon measurements made during various seasons in different years, it is necessary to eliminate the seasonal bias from the observations. The following method is used for correcting the (pco 2 ) SW data obtained during a certain month to yield an annual mean. Of a total of 55 boxes in the North Pacific north of 10 N, none of them has measurements in every month in a single year. However, 43 of them have observations for 6 months or more for the most recent 7-year period 1997 through Twelve boxes with less than 6 months data are not used in this study. We estimate a mean seasonal variability in each box using year composite data. Assuming that the seasonal variability thus estimated remained unchanged over time, we correct monthly mean values to deseasonalize them. [11] Figure 3 shows an example for our procedure. In each box, mean monthly values for (pco 2 ) SW and SST values are computed (open circles) using the 7- year composite data. If measurements are available for a given month in different years, an average is obtained (open squares). Values for the months with no observations are estimated by linearly interpolating two adjacent monthly mean values (solid squares). A single year is wrapped around from December to January. The annual mean (pco 2 ) SW value is computed using all 12 monthly values including the interpolated values (the thick horizontal line), and the difference between the annual mean and each monthly value (vertical arrows) represents a seasonal adjustment to be applied to mean monthly observations. In the bottom panel of Figure 3, the monthly mean SST values, that are computed from the 1 1 data of Reynolds et al. [2003], are shown in gray dots for the same box area. The variability of their values reflects mostly the SST trend over 10 latitude. Our SST values measured concurrently with (pco 2 ) SW are consistent with their values. In each of the 43 box areas, mean seasonal variation for (pco 2 ) SW and SST are similarly established. We assume that the seasonal variability is invariant with time, and apply the seasonal adjustment values thus derived to mean values observed for the corresponding months during different years (including the pre-1997 years) in order to deseasonalize the observations Mean Decadal Trends [12] The mean multidecadal trend of (pco 2 ) SW is estimated using observations made as far back as Since measurements were not always available for every month in each box, and since seasonal variability of (pco 2 ) SW is large, each mean monthly value is deseasonalized using the method described above. The deseasonalized mean monthly (pco 2 ) SW values are regressed linearly against time to obtain the mean decadal rate of change for (pco 2 ) SW. The SST values observed concurrently with (pco 2 ) SW are deseasonalized similarly, and the mean rate of change is 4of20

5 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 3. The year composite data for (pco 2 ) SW and SST in Box 25 N 165 W, showing the seasonal variability. The open circles are mean monthly values, and the open squares are a mean of the mean monthly values if measurements were made during the same month in different years. The solid squares are values interpolated linearly using two adjacent mean monthly values. The annual mean is shown with a thick horizontal line, and the differences between the monthly values and the annual mean (vertical arrows) are seasonal adjustments applied for mean monthly values to deseasonalize the observations made in different years. The gray lines in the bottom panel indicate the range of climatological monthly mean SST values on 1 1 grid reported by Reynolds et al. [2003] for this box area. computed by means of a linear regression. Uncertainties for the rates of change are computed using ±[s 2 /(S(X i 2 ) N(X mean ) 2 )]1 /2, where s 2 =[(S(Yi ax i b) 2 )/(N 2)] is the variance around the fitted equation Y = a X + b, and Y is (pco 2 ) SW or SST and X is year. 4. Decadal Trends of Surface Water pco 2 and SST [13] The mean decadal rates of change of (pco 2 ) SW and SST determined in 43 box areas are discussed in this section. The data and their analysis for the three areas which show contrasting trends are presented first in some detail. Overall features observed in the North Pacific will follow Area (55 ± 5 N, 145 ± 5 W) and Weather Station P [14] Figure 4 shows the surface ocean pco 2 and SST data obtained in Box 55 N 145 W (50 N 60 N, 140 W 150 W) between 1970 and This box includes the Weather Station P (50 N and 145 W), where measurements were made in by Wong and Chan [1991], in by Takahashi et al. [1991], and in by NIES, PMEL and others. This is one of the best documented areas in the North Pacific. One of the unique features of this site is that the seasonal amplitude of pco 2 is small (on the average about 50 matm) in spite of the fact that the seasonal amplitude of SST is as large as 13 C. The warming alone should increase the (pco 2 ) SW by about 70% from a typical winter value of 300 matm to 500 matm during summer. The effect of warming, however, is largely canceled by the biological drawdown of CO 2 in spring-summer months. The top panel shows all the pco 2 data at in situ temperature (solid dots) and the mean monthly values (open circles) that are deseasonalized using the method described in section 3.3. The solid line represents a linear regression line computed using the deseasonalized monthly mean values yielding a mean rate of increase of (pco 2 ) SW at SST of 19.9 ± 1.7 matm decade 1 (with 87 mean monthly values). Because of the irregular data distribution and the large amplitude of seasonal changes, we are unable to 5of20

6 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 4. Surface water pco 2 and SST observations in Box 55 N 145 W. This area includes the Weather Station P. The dots indicate individual observations, and the solid lines indicate the linear regression lines computed using the deseasonalized mean monthly values (open circles). (top) The pco 2 at SST, (middle) the SST measured concurrently with pco 2, and (bottom) the pco 2 values normalized to a constant temperature of 7.91 C, the mean of the monthly mean values. The mean decadal rate of change in each property is shown along the bottom of each panel, and N indicates the number of mean monthly values used. identify the effects of the North Pacific Decadal Oscillation and the El Niño events in terms of a slope change and displacement of the time trends. During the same period , the atmospheric pco 2 has been increasing with decadal rate ranging from 12 to 19 matm decade 1 (or 1.5 ppm yr 1 ranging from 1.2 to 1.9 ppm yr 1 over the 30-year period) as a result of anthropogenic emissions. Although the surface water pco 2 at this box area appears to be increasing at a rate consistent with the atmospheric CO 2 increase, other contributing factors will be evaluated below. [15] Since seawater pco 2 is a sensitive function of temperature, we need to examine the contribution of SST changes to (pco 2 ) SW. The middle panel in Figure 4 shows the SST data (solid dots) obtained concurrently with pco 2, and the open circles indicate the deseasonalized monthly mean SST. The straight line indicates a linear regression line computed using these monthly mean values, with a mean 6of20

7 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 temporal rate of 0.1 ± 0.1 C decade 1, that includes zero change. The validity of this estimate is tested by comparing the time trend obtained using the monthly mean SST data set from Reynolds et al. [2003], that has complete monthly coverage for Their data, when processed in the identical way used for this study, yield a mean temporal rate of 0.11 ± 0.07 C decade 1 for this box, which is consistent with that based upon our own SST data. Thus, although the time-space distribution of our data is irregular and incomplete, it appears to yield a credible estimate for the SST time trend for this box. Over the past 3 decades, the mean SST appears to have stayed constant within 0.1 C. This means that the observed increase rate in (pco 2 ) SW is caused primarily by change in seawater chemistry, and is not affected significantly by SST changes. Below, the nature of chemical changes will be further explored. [16] Freeland et al. [1997] reported the following changes in upper ocean layers at and near the Station P from 1970 to 1994; (1) a slight warming of mixed layer at a mean rate of +0.2 ± 0.1 C decade 1 (which differs from our estimate of 0.1 ± 0.1 C decade 1 for representing a broader sampling area and a longer time span), (2) a decrease in salinity at a mean rate of 0.04 ± 0.03 decade 1, (3) a decrease in the winter mixed layer thickness at a mean rate of 6.3 ± 2.8 m decade 1, and (4) a decrease in the winter-average mixed layer nitrate concentration by about 30% the from 16.2 mmol kg 1 in 1970 to 12.3 mmol kg 1 in 1994 at a mean rate of 1.6 (±1.1) mmol kg 1 decade 1. These observations are consistent mutually: As the winter mixed layer becomes shallower, subsurface waters with lesser nutrient concentrations should be mixed into the mixed layer. Their one-dimensional vertical mixing model (without biological effects) yields a rate of decrease in nitrate of 0.43 mmol kg 1 decade 1 that is nearly consistent with the field observations with the lower limit of the estimated uncertainty. The low model value suggests that other processes such as change in lateral transport may be involved. We test below whether their observations are consistent with our increasing pco 2 trend observed in this area. [17] Changes in the carbon and nutrient concentrations that are associated with the observed decrease in nitrate concentration may be estimated using the properties of sub-mixed layer waters. We assume that the chemical properties of the sub-mixed layer water that is added to the winter mixed layer is represented by the measurements made near a depth of 200 m at a WOCE Station (48.2 N and W): TCO 2 (2200 mmol kg 1 )/NO 3 (34.0 mmol kg 1 ) = 64.70, TALK (2262 meq kg 1 )/NO 3 = 66.53, SiO 3 (56.0 mmol kg 1 )/NO 3 = 1.65 and PO 4 (2.15 mmol kg 1 )/ NO 3 = The reduction of these properties corresponding to the nitrate reduction may be estimated by multiplying the observed rate of nitrate reduction of 1.60 mmol kg 1 decade 1 with each property/nitrate ratio: 106 meq kg 1 decade 1 for the total alkalinity (TALK), 103 mmol kg 1 decade 1 for TCO 2, 2.6 mmol kg 1 decade 1 for silica and 0.1 mmol kg 1 decade 1 for phosphate. Using an inorganic chemical equilibrium model with a typical winter mixed layer chemistry, we obtain a pco 2 increase of about +13 matm decade 1 at a mean winter SST of 4.0 C and salinity of The increase rate in pco 2 as subsurface water addition is reduced may be accounted for by the fact that the alkalinity is reduced faster than TCO 2 ( 106 meq kg 1 versus 103 mmol kg 1 ), and that an increasing effect on pco 2 of the lower alkalinity wins out the reducing effect of lower TCO 2. This indicates that the decrease in nitrate observed in the winter mixed layer water is consistent with the observed increase in the mixed layer pco 2. [18] Wong et al. [2002] observed that during the period, the production of CaCO 3 was low or negligibly small in the western subarctic Pacific, whereas it was high in the eastern sector ranging from 6% to 75% of the organic carbon production. This indicates that during the post-winter growing season, the surface water pco 2 was controlled by a competition of the decrease in mixed layer pco 2 by biological CO 2 fixation against the increasing effects of warming SST and alkalinity reduction by CaCO 3 production. This may contribute to the small seasonal pco 2 amplitudes observed in this region. [19] C. S. Wong et al. (unpublished manuscript, 2006) reported cooling and significant increases in surface water salinity and concentrations of nutrients and TCO 2 in the Station P area in coincidence with the1976 and 1989 La Niña periods. As shown in Figure 4, the pco 2 data failed to show the 1976 event (and no pco 2 data for the 1989 event). During this event, SST decreased by about 1.5 C, while TCO 2 and salinity increased by about 20 mmol kg 1 and 0.1, respectively. Thus the increase in TCO 2 should have increased pco 2 by 11% (40 matm), while the cooling should have decreased pco 2 by about 6.3% (22 matm). In addition, an increase of about 7 meq kg 1 in the alkalinity that is estimated assuming its proportionality with the salinity would have decreased pco 2 by about 3% (11 matm). Therefore the effect on pco 2 of the TCO 2 increase by the La Niña event is compensated largely by the cooling and alkalinity increase. This suggests that we must measure as many carbon-nutrient parameters as possible in order to understand the regulatory mechanisms for surface water pco 2 and hence the sea-air CO 2 exchange. [20] In summary, the observed pco 2 increase rate of matm decade 1 appears to be regulated not only by sea-air gas transfer and net community production of organic matter and CaCO 3, but also by changes in the upper layer dynamics including a reduction in mixed layer depth and changes in lateral transport with time Area (25 ± 5 N and 155 ± 5 W) and Station ALOHA [21] Next, we examine the 25 N 155 W box, which includes the Station ALOHA (22.7 N and 158 W) of the Hawaii Ocean Time-series (HOT) project. The top panel of Figure 5 shows individual observations (solid dots) and deseasonalized mean monthly pco 2 values (open circles) in surface waters. A linear regression for the monthly mean values yields a mean slope of 13.0 ± 2.2 matm decade 1 for the period. As observed for Box 55 N 145 W, the rate of increase in surface water pco 2 is consistent with the mean rate of increase for atmospheric pco 2. The SST in this area is virtually unchanged over the 30-year period (the middle panel, Figure 5): 0.1 ± 0.2 C decade 1 on the basis of our SST data obtained concurrently with the pco 2 measurements and 0.04 ± 0.04 C decade 1 based on the monthly data of Reynolds et al. [2003]. Hence the 7of20

8 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 5. Surface water pco 2 and SST observations in Box 25 N 155 W. This area includes the Station ALOHA of the HOT program. The dots indicate individual observations, and the solid lines indicate the linear regression lines computed using the deseasonalized mean monthly values (open circles). The crosses indicate the (pco 2 ) SW and SST values at the Station ALOHA reported by Dore et al. [2003]. Since pco 2 was not measured during the HOT program, these pco 2 values have been computed using the measured alkalinity and TCO 2 values, and hence are not used in our analysis. (top) The pco 2 at SST, (middle) the SST measured concurrently with pco 2, and (bottom) the pco 2 values normalized to a constant temperature of C, the mean of the monthly mean values, that are used to estimate changes in TCO 2. observed pco 2 increase rate is due mainly to changes in the water chemistry. [22] On the basis of time series measurements of the alkalinity, TCO 2, temperature and salinity in surface waters at the Station ALOHA, Dore et al. [2003] computed (pco 2 ) SW and reported a mean increase rate of 24.6 ± 2.8 matm decade 1 for the 13-year period, Keeling et al. [2004] used also the alkalinity and TCO 2 data, that were determined by them independently for the Station ALOHA samples, and reported that (pco 2 ) SW increased at a mean rate of 14 ± 2 matm decade 1 for the first 8-year period, , and at a much faster rate of 32 ± 4 matm decade 1 for the following 5-year period, On the basis of a thorough analysis of the data 8of20

9 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 using a diagnostic box transport model, they concluded that the post-1997 increase in (pco 2 ) SW was a result of local decrease in precipitation as well as to a regional change in water mass distributions, that is perhaps related to the Pacific Decadal Oscillation (PDO). Thus their mean rate of increase of 25 ± 1 matm decade 1 for their 14-year study period of is in agreement with that reported by Dore et al. [2003]. [23] The 14-year mean rate of (pco 2 ) SW increase determined by these investigators are twice as large as our 35-year mean rate 13.0 ± 2.2 matm decade 1 estimated using the directly measured pco 2 values, whereas the pre year mean rate of 14 ± 2 matm decade 1 obtained by Keeling et al. [2004] is consistent with ours. Two reasons may be considered to account for the difference: The first is our undersampling problem, and the second is incompatibility of the computed with the measured pco 2 values. [24] First, since we have no pco 2 observations during and a limited number of observations during as shown in Figure 5, our observations are not sufficient for documenting the transition to a higher rate of pco 2 increase reported by the previous investigators. Furthermore, since the amplitudes for the pco 2 seasonal variability (about 60 matm) and are much greater than the mean decadal rate of pco 2 increase (14 to 34 matm decade 1 ), a trend change is difficult to detect unless observations over a sufficiently long period are available. If our data are linearly regressed for the preand post-1996 periods separately, they yield a rate of 9.8 ± 4.7 matm decade 1 (N = 29) for , and 15.1 ± 10.3 matm decade 1 (N = 18) for These rates are statistically indistinguishable owing to the large uncertainties which result from undersampling, large seasonal variability and our broader sampling area. Although our data are not sufficient to resolve the trend change, they allow us to obtain a mean rate of pco 2 change over the 3 decades. [25] Second, Dore et al. [2003] and Keeling et al. [2004] both computed (pco 2 ) SW using an inorganic chemistry model, whereas we measured the pco 2 directly using an air-water equilibration method. In inorganic models, the ionization effects of organic acids are neglected. Organic acids may affect the carbonate equilibria by the ionic dissociation (i.e., organic acid alkalinity) and/or by forming complex ions with other ions. Considering the fact that the total dissolved organic carbon (DOC) ranges from about 60 to 100 mmol kg 1, the organic carbon alkalinity could be of an order of 10% of DOC concentration. This must be subtracted from the total (or titration) alkalinity in order to compute the carbonate alkalinity, and hence it alters the computed pco 2 values by as much as 5% [Millero et al., 2002]. However, the organic alkalinity cannot be determined since the compositions of organic acids within DOC and their dissociation constants are not known. Hence the information needed for evaluating the systematic differences between the computed and measured (pco 2 ) SW values are not available presently. Keeling et al. [2004] cited Dore s personal communication that the computed values are systematically greater than a limited number of directly measured pco 2 values by 3.6 ± 5.7 matm. However, whether the differences are distributed randomly or correlated with salinity is not clear from their descriptions. Thus we are unable to evaluate whether the pco 2 trend change in association with salinity changes is real or due to changes in organic alkalinity associated with changes in salinity and water masses Area (50 N 60 N and W) in the Central Bering Sea [26] The data in this box (Figure 6) are entirely from the southern Bering Sea, which is free of ice cover year round. The open circles indicate the deseasonalized mean monthly values, which are used for computing the mean rate of change using a linear regression. In contrast to the two previous areas in the open Pacific, the surface water pco 2 decreases with time at a mean rate of 17 ± 12 matm decade 1. The mean SST, meanwhile, stayed virtually unchanged (Figure 6, middle panel): 0.1 ± 0.2 C decade 1 based on our measurements made concurrently with pco 2 and 0.01 ± 0.05 C decade 1 on the basis of the data of Reynolds et al. [2003]. Using the pco 2 data corrected to a constant temperature of the area mean 5.76 C, we estimate that pco 2 decreased at a rate of 19 ± 13 matm decade 1 (Figure 6, bottom panel) as a result of changes in chemistry. If this change were assumed to be due solely to TCO 2 change, a TCO 2 decrease of 8 ± 5 mmol kg 1 decade 1 is expected using a Revelle factor of 14.5 (see Table 2 in section 4.5). Similar decreasing trends are observed in other areas located within the Bering Sea and in waters just outside the Okhotsk Sea. We discuss below various factors which contribute to the observed trend. [27] On the basis of the Coastal Zone Color Scanner (CZCS; ) and Sea-viewing Wide Field-of view Sensor (SeaWiFS; ) ocean color measurements from satellites, Gregg et al. [2003] reported that the ocean primary production increased substantially in the Bering and Okhotsk Seas over the decade. However, their results could be uncertain because of persistent cloud cover in the region, which could bias the CZCS and SeaWiFS data differently. Although their observations reflect the gross production rather than the net community production which is relevant to changes in pco 2 and TCO 2 in seawater, the reported trends are considered to be a factor that contributes to the observed pco 2 trend. The increase in the productivity may reflect changing supplies of nutrients into the Bering Sea caused by changes in land hydrology or by increases in airborne or river inputs of nutrients mediated by anthropogenic activities. [28] In the southeastern Bering Sea, the mixed layer depth decreased and stratification increased as a result of warming and freshening of surface waters from May 2001 through October 2004 [Wirts and Johnson, 2005]. Significant warming was also observed at a mooring (56.8 N and 164 W) in the eastern shelf area in [Overland and Stabeno, 2004]. This may be due to the combined effects of weak wind forcing and increased inflow of the warmer, nutrient-rich Alaskan Stream water. As a result, while the supply flux of nutrients and CO 2 into the surface layer by the wintertime vertical mixing was decreased, lateral supply of nutrients into the Bering Sea was increased. This, combined with an increase in productivity due to earlier ice melt, may have caused the decadal decrease in surface water pco 2. Even if the community production per unit ocean surface area remains unchanged, a reduction in mixed 9of20

10 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 6. Surface water pco 2 and SST observations in Box 55 N 175 W located in the central Bering Sea. The dots indicate individual observations, and the solid lines indicate the linear regression lines computed using the deseasonalized mean monthly values (open circles). (top) The pco 2 at SST, (middle) the SST measured concurrently with pco 2, and (bottom) the pco 2 values normalized to a constant temperature of 5.76 C, the mean of the monthly mean values. layer depths due to warming (i.e., a smaller volume of photic waters per unit area) could cause a greater lowering of pco 2 and TCO 2 in the mixed layer. [29] The observed decrease in (pco 2 ) SW in the Bering Sea translates to a ph increase of about 0.02 ± 0.01 decade 1, suggesting that the seawater has become more alkaline and that the H + ion concentrations have decreased by 14% for the past 30 years. This is in contrast to the acidification of 0.1 ph unit in open ocean that is estimated for surface ocean waters in equilibrium with atmospheric pco 2 of 280 matm during the pre-industrial period and today s 380 matm [Feely et al., 2004]. Thus the Bering Sea may provide an interesting environment for investigating the contrasting ph effects on marine ecosystems Decadal Change of pco 2 in the North Pacific Surface Waters [30] The mean decadal rates of change in (pco 2 ) SW observed in 43 box areas are tabulated in Table 1. As mentioned earlier, all of these boxes have observations in 6 or more months during the most recent 7-year period, However, since many of the boxes have been 10 of 20

11 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Table 1. Decadal Mean Rate of Change of Surface Water pco 2 and SST in Box Areas a N. Lat. N Long. E or W Climatological Mean SST Rate, C decade 1 pco 2 Data Mean SST Rate, C decade 1 Difference, C decade 1 Surface Water pco 2 Mean Rate of Change, matm decade 1 Increase Decrease 15 ± 5 135E ± ± ± ± ± 5 145E ± ± ± ± ± 5 165W ± ± ± ± ± 5 155W ± ± ± ± ± 5 145W ± ± ± ± ± 5 125W ± ± ± ± ± 5 115W ± ± ± ± ± 5 105W ± ± ± ± ± 5 95W ± ± ± ± 9.9 b 25 ± 5 135E ± ± ± ± ± 5 145E ± ± ± ± ± 5 165E ± ± ± ± ± 5 165W ± ± ± ± ± 5 155W ± ± ± ± ± 5 135W ± ± ± ± ± 5 125W ± ± ± ± ± 5 115W ± ± ± ± ± 5 145E ± ± ± ± ± 5 155E ± ± ± ± ± 5 165E ± ± ± ± ± 5 175E ± ± ± ± ± 5 175W ± ± ± ± ± 5 165W ± ± ± ± ± 5 155W ± ± ± ± ± 5 145W ± ± ± ± ± 5 135W ± ± ± ± ± 5 125W ± ± ± ± 7.7 b 45 ± 5 145E ± ± ± ± ± 5 155E ± ± ± ± ± 5 165E ± ± ± ± ± 5 175E ± ± ± ± ± 5 175W ± ± ± ± ± 5 165W ± ± ± ± ± 5 155W ± ± ± ± ± 5 145W ± ± ± ± ± 5 135W ± ± ± ± ± 5 125W ± ± ± ± ± 5 175E ± ± ± ± ± 5 175W ± ± ± ± ± 5 165W ± ± ± ± ± 5 155W ± ± ± ± ± 5 145W ± ± ± ± 2.3 b 55 ± 5 135W ± ± ± ± 3.6 Totals Climatological Mean SST Rate, C decade 1 pco 2 Data Mean SST Rate, C decade 1 Difference, C decade 1 Surface Water pco 2 Mean Rate of Change, matm decade 1 Increase Decrease Basin Mean Standard Deviation ±0.21 ±0.80 ±0.85 ±4.8 ±5.7 Number of Box Areas Standard Error ±0.03 ±0.12 ±0.92 ±2.9 a Climatological Mean SST Rate is the SST change rates computed using the complete monthly data for by Reynolds et al. [2003]; pco 2 Data Mean SST Rate is the SST change rates computed using the observations made concurrently with pco 2 measurements. Difference is the difference between climatological mean SST rate and pco 2 data mean SST rate, and the boxes which exceed a difference of ±0.85 C decade 1 are rejected as insufficient observations. The box areas to be rejected are marked with bold numbers. Surface Water pco 2 Mean Rate of Change is obtained by a linear regression of the deseasonalized mean monthly pco 2 values: the positive rates are on the left side column, and the negative rates are in italics listed on the right side column. b Eastern boundary coast area. sampled at irregular intervals with varying degrees of sampling density during the study period of , the reliability of our method for computing mean rates of change needs to be tested. For this purpose, we established the following criteria using the available SST data. We compute the decadal mean rate of change for SST in each box using the complete monthly set (1 1 spatial resolution) obtained by Reynolds et al. [2003] for (Table 1, third column), and compare it with that computed using the SST values measured concurrently with 11 of 20

12 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 Figure 7. Mean decadal rate of change of surface water pco 2 (at SST) in the North Pacific. These changes in pco 2 include the effects of changes in SST and other factors, and represent actual rates of change in the ocean pco 2. The numbers in bold letters indicate the rate of increase in matm decade 1, the light italic numbers indicate the rate of decrease in matm decade 1, and the values in parentheses indicate the uncertainty in the same unit. The gray curves in the western Pacific show the approximate locations of the Polar Front (PF) [Belkin et al., 2002] and Subtropical Front (STF). pco 2 (Table 1, fourth column). The mean for the 43 pairs of the differences between columns 3 and 4 is 0.18 C decade 1 with a standard deviation of 0.85 C decade 1 (Table 1, fifth column). The boxes which have SST change rates that differ from the climatological SST rates of the third column by more than 0.85 C decade 1 are considered to be unreliable, and hence 11 boxes are rejected (as indicated with bold numbers in fifth and sixth columns), and the remaining 32 boxes are accepted for further study. [31] Figure 7 shows the geographical distribution of (pco 2 ) SW change rates in these 32 boxes, of which 4 have negative rates and 28 have positive rates. The three boxes with negative rates are located in the Bering and one is outside the Okhotsk Sea along the Kuril chain, and possible causes for the negative trends have been discussed in section 4.3). The four boxes listed in the sixth column of Table 1 give a mean rate of 11.1 ± 5.7 matm decade 1, whereas the mean of the 28 open ocean boxes is 12.0 ± 4.8 matm decade 1 (Table 1). The map shows that a western temperate zone between 30 N 40 N and 150 E 180 appears to have slower rates ( matm decade 1 ) than the open ocean mean. They are located in the areas of the Kuroshio Current between the polar and subtropical fronts and also are influenced by the Oyashio Current which contains the outflow from the Okhotsk Sea with negative pco 2 change rates. Furthermore, according to Chen and Wang [1999] and Chen et al. [2004], the alkalinity in the East China Sea is high as a result of the input from the two major rivers draining the Chinese mainland and from the oxidation of organic debris in the broad shelf sediments. Hence we speculate that the East China Sea waters would have a negative rate of pco 2 change caused by an increasing influx of alkalinity, and hence that entrainment of these waters into the Kuroshio Current might lower the rate of pco 2 increase. A long-term monitoring of chemical properties of the river and shelf waters is needed for our improved understanding of the influence of river waters and continental processes on the open ocean water carbon chemistry. [32] Coastal areas along the North and Central Americas exhibit widely varying rates ranging from 2.0 ± 7.7 matm decade 1 off the Oregon coast, to 19.9 ± 2.3 matm decade 1 in the Gulf of Alaska and 25.5 ± 9.9 matm decade 1 in the Guatemala Basin. In these eastern boundary areas, the rate of pco 2 change depends largely on vertical and lateral circulation rates of water, and is influenced primarily by basin- and local-scale meteorological and climatic conditions rather than the sea-air CO 2 flux [e.g., Friederich et al., 2002; Hales et al., 2005; van Geen et al., 2000]. Although our SST data in these areas give decadal rates consistent with those from Reynolds et al. [2003], the wide range of decadal (pco 2 ) SW rates is most likely due to undersampling of highly variable coastal systems. [33] The mean rate of (pco 2 ) SW increase for all 28 box areas is 12.0 ± 4.8 matm decade 1 (Table 1). If the above three boxes located in the North American coastal upwelling areas (footnote b in Table 1) are excluded, the remaining 25 open ocean boxes give 12.0 ± 3.8 matm decade 1. These rates are some what slower than the mean rate of atmospheric CO 2 increase of about 15 matm decade 1 (or 1.5 ppm yr 1 ) observed at the Mauna Loa Observatory over the 33-year period [Keeling and Whorf, 1994]. We are thus unable to resolve unequivocally whether the North Pacific mixed layer waters as a whole are 12 of 20

13 TAKAHASHI ET AL.: NORTH PACIFIC SURFACE WATER PCO 2 taking up CO 2 from the atmosphere at the same or slower rate as the atmospheric CO 2 increase. The ocean mean rate of 12.0 matm decade 1 is, however, consistent with the results of a simple time-dependent box model for the sea-air CO 2 exchange with a constant gas transfer rate of 20 mol CO 2 m 2 yr 1 [Broecker et al., 1986]: The surface water pco 2 lags behind the atmospheric pco 2 by about 3 years, and the rate of increase in surface water pco 2 should be only a few percent slower than that for the atmospheric increase rate. Greater differences between the increase rates of atmospheric and surface ocean pco 2 would imply changes in ocean circulation, wind intensity, and/or marine ecosystems such as a shift in phosphate-nitrate dynamics observed in the North Pacific [Karl, 1999; Karl et al., 2001] Decadal Change of Surface Water CO 2 Chemistry [34] Large effects of temperature on (pco 2 ) SW may be removed by normalizing it to a constant temperature (the mean SST of each box area), so that the variation in chemical properties in seawater may be identified. For this purpose, we use a constant (@ ln pco 2 /@ T) of C 1 [Takahashi et al., 1993] and its integrated form, pco 2 (at mean SST) = pco 2 (observed) Exp [ (mean SST observed SST)]. The bottom panels of Figures 4, 5 and 6 show the (pco 2 ) SW values that are normalized to the mean SST in respective box areas. The significance of the temperature-normalized pco 2 values may be illustrated by the following cases. Consider that a parcel of surface seawater (with a constant alkalinity and salinity) is always in equilibrium with increasing atmospheric CO 2 at a rate of 15 matm decade 1. If the water is warmed at a rate of 1 C decade 1, its pco 2 should increase at a rate of about 15 matm decade 1, same as atmospheric CO 2. Hence this water has the same pco 2 as in the atmosphere CO 2, inferring that no net transfer of CO 2 occurs throughout the decade, and the TCO 2 in the water remain unchanged. The (pco 2 ) SW values that are normalized to the initial temperature remain unchanged through the decade. On the other hand, if the water is kept at a constant temperature, the pco 2 in water increases by taking up atmospheric CO 2, and also its TCO 2. Conversely, a decrease in the temperature-normalized pco 2 values indicates a reduction in TCO 2 in seawater with time. In real oceans, the rate of change of temperature-normalized pco 2 depends not only on the sea-air CO 2 flux, but also on many other processes that govern the carbon dynamics in the mixed layer. These processes include seasonal and interannual changes in SST, alkalinity, ecosystem structure (which affects the Redfield ratios and carbon export from the mixed layer), nitrification of atmospheric nitrogen, depths and rate of upwelling waters (which affect preformed concentrations of CO 2 and nutrients) and river water flux. [35] In Table 2, the decadal mean rate of change of the temperature-normalized pco 2 in each box area is listed, and its geographical distribution is shown in Figure 8. The rates in open ocean areas are positive indicating that TCO 2 is increasing, and show some geographical variability. In the western North Pacific areas north of the Subtropical Front (STF), the observed rates of increase (3.5 to 11.4 matm decade 1 ) appear to be smaller than the mean atmospheric CO 2 increase rate of about 15 matm decade 1. Increasing flux of nutrients from the East China Sea might cause an increase in net community production, and hence a reduction in TCO 2. In contrast, the rates tend to be greater than the atmospheric rate in the northeastern areas. This may be related to change in the rate and source depth for upwelling waters. The negative rates of change observed in the Bering Sea and outside the Okhotsk Sea have been discussed previously. A negative rate is also observed in a box area (35 ± 5 N) along the west coast of North America. The upper layer dynamics and primary production in this area are affected significantly by changes in wind regimes and ocean current structure associated with El Niño events [Goes et al., 2004]. Because of the complexity in the area, a simple answer for the negative trend cannot be offered on the basis of the available data. [36] The mean rate of increase for 27 box areas in the open Pacific (Table 2) is 12.6 ± 6.6 matm decade 1. If this is assumed to be caused by change in TCO 2 without temporal changes in alkalinity and salinity in each box area, these pco 2 rates may be converted to TCO 2 changes using the Revelle factor, which is estimated for each box using the available pco 2 and TCO 2 data (see Table 2). The mean rate of TCO 2 increase in the surface layer of the North Pacific thus estimated is 7.4 ± 3.3 mmol kg 1 decade 1. This is somewhat lower but is indistinguishable from the areaweighted mean rate of increase of 8.1 ± 1.0 mmol kg 1 decade 1 for the North Pacific surface waters that are assumed to be in equilibrium with the current atmospheric CO 2 increase rate of 15 matm decade 1 and have no decadal change in the alkalinity Comparison of TCO 2 Changes With Other Studies [37] Our estimate of TCO 2 increase may be compared with the rate estimated based on direct measurements of TCO 2 in the northeastern Pacific. On the basis of direct TCO 2 measurements made during seven expeditions spanning from 1973 to 2000, Feely et al. [2003] and Sabine et al. [2004b] estimated a mean rate of TCO 2 increase of 13 ± 2 mmol kg 1 decade 1 for surface mixed layer waters over 30 N 50 N and 140 W 180. Analyzing the depth profile data, Sabine et al. [2004b] used a multiple linear regression analysis for the station data, and reported an average TCO 2 increase rate of 7.9 ± 4 mmol kg 1 decade 1 for the North Pacific waters below the mixed layer down to 1250 m. This difference is noteworthy because the surface water data yielded a rate much greater than 8 mmol kg 1 decade 1 expected from equilibration with atmospheric CO 2 increasing at a mean rate of 15 matm decade 1. [38] Using the data listed in Table 2, we computed a mean rate of increase in TCO 2 of 6.0 ± 4.5 mmol kg 1 decade 1 for the same area studied by Feely et al. [2003] and Sabine et al. [2004b]. While this is about one half of their surface water value, this is consistent with their rate based on the depth profile data. The strength of Feely-Sabine approach is that they measured TCO 2 directly. On the other hand, their measurements were made during seven expeditions that took place over several years during nonwinter months, and the effects of seasonal changes in mixing and biological utilization were corrected using a multiple linear regression with salinity and nutrient data and a fixed P:N:C stoichiometry ratio for net production. Their surface TCO 2 increase may be affected by the lack of full-seasonal cycle observations. In contrast, the pco 2 data used in this study 13 of 20