59

Variability of Surface Layer CO2 Parameters in the Westernand Central Equatorial Pacific

Masao ISHII1, Shu SAITO1, Takayuki TOKIEDA1, Takeshi KAWANO2,Kazuhiko MATSUMOTO2 and Hisayuki Y. INOUE1*

1Geochemical Research Department, Meteorological Research Institute,1-1 Nagamine, Tsukuba, Ibaraki 305-0052, Japan

2Ocean Observation and Research Department, Japan Marine Science andTechnology Center, 2-15 Natsushima, Yokosuka 237-0061, Japan

Abstract. We present the results of the oceanic and the atmospheric CO2measurements in the warm pool and in the divergence zone of the western andcentral equatorial Pacific (135°E to 160°W), made during 18 cruises fromJanuary 1994 to January 2003. The majority of these measurements were madein boreal winter (December to February), the parameters measured includingpartial pressure of CO2 in near-surface seawater (pCO2sw) and in the atmosphere(pCO2air), total inorganic carbon (TCO2), and total hydrogen ion concentration(pH). Total alkalinity (TA) was calculated from these parameters using otheroceanographic parameters.

In the warm pool in the western equatorial Pacific where salinity isrelatively low (S < 35), a moderate CO2 supersaturation (0 < ∆pCO2/µatm <+30: ∆pCO2 = pCO2sw – pCO2air) due to high temperature (28.5 < T/°C < 31)and low salinity-normalized TCO2 (NTCO2: ca. 1940 µmol kg–1 at S = 35) wastypically observed. In this region, the variability of pCO2sw is controlledmainly by the variation of temperature and salinity. However, when the barrierlayer has developed, pCO2sw decreased to turn to a slight CO2 subsaturation(–10 < ∆pCO2/µatm < 0) with the decrease of NTCO2 (<1930 µmol kg–1) dueto net biological uptake. In the equatorial divergence zone where salinity isrelatively high (S > 35), near-surface water is highly supersaturated with CO2(+30 < ∆pCO2/µatm < +120) as a results of the higher NTCO2 (1960 < NTCO2/µmol kg–1 < 2030) and low salinity-normalized TA (NTA: ca. 2320 µmolkg–1 at S = 35). The westward decrease of high pCO2sw in this zone is ascribedto the westward decrease of NTCO2 and a small variation in NTA, but ca. 60%of these effects are compensated by the westward elevation of temperature. Theabrupt change in pCO2sw at the boundary zone between the warm pool and thedivergence zone is explained largely by the change in NTCO2 and partly by thechange in salinity.

Empirical relationships between NTCO2 and temperature and betweenNTA and temperature in near-surface water were derived both for the warmpool and for the divergence zone based on a total of 5069 data from underwaymeasurements and 184 from discrete samples taken after 1990. Combiningthese relationships and a data set of temperature, fields of NTCO2, NTA, and

*Present address: Graduate School of Environmental Earth Science, Hokkaido University,N10W5, Kita, Sapporo 060-0810, Japan.

Global Environmental Change in the Ocean and on Land, Eds., M. Shiyomi et al., pp. 59–94.© by TERRAPUB, 2004.

60 M. ISHII et al.

pCO2sw were obtained monthly for the equatorial Pacific (135°E–95°W,10°S–5°N). The sea-to-air CO2 flux from this zone evaluated on the basis ofthese pCO2sw fields in the period 1990–2000 was +0.38 ± 0.16 PgC year–1 onaverage, and it varied between +0.04 and +0.80 PgC year–1 in conjunction withthe El Niño Southern Oscillation.

Relationships among NTCO2, NTA and concentrations of inorganicmacronutrients in the divergence zone also varied year by year. Using theserelationships and a simple scheme of zonal advection, ratios of sinking particulatecarbonate flux: net sea-to-air CO2 flux: net community production wereestimated to vary within the range (0.02 to 0.17):(0.39 to 0.45):1, which in turncorresponded to the sinking particulate carbonate flux of 0.4 to 4.3 mmolm–2day–1, net sea-to-air CO2 flux of 4.0 to 9.9 mmol m–2day–1, and netcommunity production of 7.6 to 25 mmol m–2day–1 in the western part of thedivergence zone during the La Niña period in the years 1999 to 2001.

Keywords: equatorial Pacific, total inorganic carbon, partial pressure of CO2,sea-to-air CO2 flux, net community production, particulate carbonate

1. INTRODUCTION

The equatorial Pacific has been attracting a great deal of interest from geophysicistssuch as oceanographers and meteorologists as a region of the complex ocean-circulation and high primary productivity as well as a region of the strongestconvection in the troposphere. The fluctuation of the atmosphere-ocean interactionknown as El Niño Southern Oscillation (ENSO) is a prominent phenomenon inthis region that affects the global climate. Concurrent changes in the variousaspects of the ocean and the atmosphere that might be a proxy for the changesassociated with the global warming indicate close interactions among climate,ocean circulation, chemical environment, biological activities, and behavior ofgreenhouse gases such as CO2.

Investigations of the variability of partial pressure of CO2 in surface water(pCO2sw) in the equatorial Pacific have been developing along with those of theatmospheric CO2. In 1957 and 1961, Keeling et al. (1965) measured pCO2sw bya method of direct equilibration in an extensive area in the Pacific Ocean,discovering a significant supersaturation of CO2 in the central equatorial zone.From late 1960s to early 1970s, Miyake et al. (1974) also surveyed pCO2sw overthe North and South Pacific, and documented significant CO2 supersaturation inthe equatorial region and off South America (Miyake et al., 1974; Inoue et al.,1999). In 1979–1980, Weiss et al. (1982, 1992) repeated measurements betweenHawaii and Tahiti, and they reported the seasonal variation in pCO2sw. Since thelate 1980s, oceanographic observations including measurements of pCO2sw havebeen made more frequently with the increase in concern for the El Niño event.These observations revealed that pCO2sw in the equatorial Pacific exhibits a greatvariability in conjunction with the ENSO, potentially affecting the increasing rateof atmospheric CO2 concentration (Feely et al., 1987; Fushimi, 1987; Inoue andSugimura, 1992; Wong et al., 1993). With the increase of attention to climatechange, more data were obtained in 1990s for the carbon cycle in the ocean. Data

Variability of Surface Layer CO2 Parameters 61

on pCO2sw in the equatorial Pacific have significantly increased, too (e.g. Goyetand Peltzer, 1994; Lefevre et al., 1994; Feely et al., 1995, 1997; Dandoneau,1995; Ishii and Inoue, 1995; Watai et al., 1999; Etcheto et al., 1999; Inoue et al.,2001). The flux of CO2 from the equatorial Pacific to the atmosphere wereevaluated as ca. 0.6 PgC year–1 (1 PgC = 1 × 1015 g) during a non-El Niño period,and this zone is recognized as being a dominant source of CO2 to the atmosphere(e.g. Takahashi et al., 2002). However, the CO2 flux from this zone decreases to0.2–0.4 PgC year–1 during the El Niño periods, increasing to 0.8–1.0 PgCyear–1 during the La Niña periods (Feely et al., 2002). Such a large interannualvariability is thought to have a major impact on the net CO2 flux from theatmosphere to the world oceans, which is considered as ca. 2 PgC year–1 onaverage (Le Quéré et al., 2000; Obata and Kitamura, 2003).

The equatorial Pacific is divided into two major domains. In the west,salinity is lower (S < 35) due to the largest annual precipitation in the worldoceans (You, 1998) and temperature is higher (T/°C > 28.5) in the surface layers.Consequently, vertical stratification is usually attained. One or more haloclinesare often formed above the thermocline (e.g. Ando and McPhaden, 1997) and theformation of the barrier layer—the layer between the surface mixed layer and thebottom of the isothermal layer—insulates the surface shallow mixed layer fromthe cold, CO2-rich water below the thermocline (Lukas and Lindstrom, 1991).Macronutrients such as nitrate are depleted due to the stratification. The level ofCO2 is usually close to the air-sea equilibrium value in surface layers so that thenet sea-to-air CO2 flux is small (e.g. Ishii and Inoue, 1995) except for the strongEl Niño periods, when CO2 is significantly supersaturated with respect to theoverlying atmosphere (Fushimi, 1987; Inoue et al., 2001).

In the central and eastern regions, on the other hand, the equatorial divergenceand turbulent mixing driven by the easterly trade wind bring the water of thenarrow Equatorial Under Current (EUC), which flows eastward across the basinaround the thermocline, into the surface layer. Accordingly, temperature is lower(22 < T/°C < 28.5), nutrients remain at moderate concentrations, and CO2 ishighly supersaturated in surface layers. The boundary between this equatorialdivergence zone (or “cold tongue”) and the western warm pool can often bedefined with the abrupt changes in the pCO2sw and salinity rather than temperature.It migrates east-west following the strengthening, weakening, and even reversalof the easterly (e.g. Johnson et al., 2000), and the longitude of the boundary isknown to be closely correlated with the Southern Oscillation Index (SOI) (Inoueet al., 1996; Le Borgne et al., 2002). The displacement of the equatorialdivergence zone by the warm pool during the El Niño periods and vice versaduring the La Niña periods is one of the main causes of the large variability of thenet sea-to-air CO2 flux from the equatorial Pacific (Boutin and Etcheto, 1997).

The factors controlling the variability of pCO2sw in the warm pool and in thedivergence zone of the equatorial Pacific, however, are not fully understood interms of the oceanic CO2 chemistry. Variability in the pCO2sw, which isproportional to the concentration of aqueous CO2, is physicochemically ascribedto the variability of total inorganic carbon concentration (TCO2), and total

62 M. ISHII et al.

hydrogen ion concentration (pH) or total alkalinity (TA) with information onother acids and bases, as well as temperature and salinity. In terms of oceanography,these properties are perturbed by the various phenomena related to the fluxes ofheat and freshwater that affect temperature and salinity, biological activities thatconsume and exhaust CO2 and carbonate (CO3

2–), and CO2 exchange across theair-sea interface. Horizontal and vertical circulation of water is also a criticalfactor that alters these properties. In the warm pool of the western equatorialPacific, hampered vertical mixing is an important factor that is responsible for thehigh temperature and lower pCO2sw. In the central and eastern equatorial Pacific,on the other hand, the equatorial divergence and its variability is a factorresponsible for the high, variable pCO2sw in the surface layers. The role ofbiological activities as consumers of CO2 and their significance relative to thesea-to-air CO2 evasion for the subsequent CO2 decrease from the surface layer aswell as the fate of organic carbon assimilated has, however, not been consistentlyquantified (Quay, 1997) and their seasonal and interannual variability has beenpoorly documented.

In order to quantify the flux of carbon among carbon reservoirs in theatmosphere and in the surface and deeper layers of the ocean, and to evaluate thecontrolling factors on the variability of pCO2sw and sea-to-air CO2 flux in thecontext of the carbon cycle, it is critical to identify the temporal and spatialvariability of the oceanic CO2 parameters including TCO2, pH and/or TA as wellas pCO2sw. In the eastern equatorial Pacific to the east of 170°W, investigationsconducted as part of comprehensive studies, such as Joint Global Ocean FluxStudy (JGOFS) in early 1990s, documented the variability of CO2 parametersassociated with the ENSO (e.g. Wanninkhof et al., 1995; Archer et al., 1996)using a newly developed, precise coulometric technique for TCO2 analysis(Johnson et al., 1985). In the western and central equatorial Pacific on the otherhand, while much effort has been devoted to clarifying the variability of pCO2sw,studies on the other parameters of the oceanic CO2 chemistry had been rathersparse until 1980s. They include the pioneering works of Takahashi et al. (1982)during the Geochemical Ocean Sections Study (GEOSECS) Pacific Expeditionand of Feely et al. (1987) during and after 1982/1983 El Niño. In the early 1990s,several intermittent cruises along north-south transects were made as a part ofWorld Ocean Circulation Experiment (WOCE) and Northwest Pacific CarbonCycle Study (NOPACCS) (Tsubota et al., 1999). In addition, the studies oncarbon cycle that include measurements of CO2 parameters have also beenconducted along the equator (Le Borgne et al., 1995). Consequently, the increasein TCO2 on moving from the west to the east in the western and central equatorialPacific was recognized. However, they were still not sufficient to provide asatisfactory documentation of the temporal and spatial variability of CO2parameters.

Since January 1994, we have conducted concurrent underway measurementsof pCO2sw and TCO2 as well as TCO2 measurements at hydrographic stations ina total of 18 cruises in the western and central equatorial Pacific between 135°Eand 160°W (Table 1). Since January 1999, we have started precise

Variability of Surface Layer CO2 Parameters 63

spectrophotometric pH measurements as well. Many of these observations wereconducted along the equator in boreal winters. This paper reviews the results ofthe oceanic CO2 observations with an emphasis on the zonal distribution andtemporal variation of TCO2 in the warm pool and in the divergence zone of thewestern and central equatorial Pacific. Below we estimate monthly fields ofsurface TCO2, pCO2sw and sea-to-air CO2 flux in the equatorial Pacific based onthe empirical relationships found for this region. Finally, the role of biologicalCO2 uptake and sea-to-air CO2 transport in changing the oceanic CO2 parametersin the western part of the divergence zone is also quantified.

2. SAMPLING AND ANALYSIS METHODS

The majority of data for the oceanic CO2 parameters were obtained duringthe cruises of R/V Kaiyo (1994–1999) and R/V Mirai (1998–2003) of JapanMarine Science and Technology Center (Table 1).

For pCO2 analyses, underway measurements of the CO2 concentration (molefraction of CO2 in air; xCO2) in air equilibrated with a great excess of near-surfaceseawater and in the overlying atmosphere were made using the automated air-seaCO2 analyzer (Nippon ANS Co.) (Inoue, 1999) equipped with a non-dispersiveinfrared (NDIR) gas analyzer BINOS 4. Equipment of the same design has beenused in the international intercomparison exercises for pCO2 measurements heldat the Scripps Institution of Oceanography in June 1994 (Dickson, 1994) and onboard the R/V Meteor in June 1996 (Körtinger et al., 1999, 2000). Themeasurements were made continuously in a 1-hour or 1.5-hour cycle, including

Table 1. List of cruises in the equatorial Pacific (to the south of 5°N) in which both pCO2sw andTCO2 were measured concurrently.

Month year Latitude Longitude Ship Cruise

Jan–Feb 1994 5°N–4°S 135°E–165°W Kaiyo KY9401Aug 1994 5°N–3°S 139°E–142°E Ryofu maru RY9407(WHP-P9)Nov–Dec 1994 5°N–0° 146°E–165°W Kaiyo KY9411Feb 1995 5°N–3°S 147°E–153°E Hakuho maru KH94-4Dec 1995–Jan 1996 5°N–0° 147°E–165°W Kaiyo KY9512Feb 1997 3°N 137°E Ryofu maru RF9701Oct 1997 4°N 146°E Ryofu maru RF9709Dec 1997–Jan 1998 5°N–0° 136°E–180° Kaiyo KY9712Feb 1998 5°N–0° 176°E–164°W Mirai MR97K2Jan 1999 5°N–0° 135°E–167°W Mirai MR98K2Jan–Feb 1999 5°N–1°S 137°E–168°E Kaiyo KY9901Oct–Nov 1999 5°N–5°S 137°E–156°E Mirai MR99K6Nov–Dec 1999 5°N–0° 135°E–167°W Mirai MR99K7Jan 2001 5°N–0° 145°E–160°W Mirai MR00K8Feb–Mar 2001 5°N–5°S 147°E–156°E Mirai MR01K1Jan–Feb 2002 5°N–0° 145°E–160°W Mirai MR02K1Mar 2002 5°N–5°S 145°E–156°E Mirai MR02K2Nov 2002–Mar 2003 5°N–5°S 138°E–160°W Mirai MR02K6

64 M. ISHII et al.

calibration with four standard gases (300–450 ppm CO2 in air, Nippon Sanso Co.,Ltd.) followed by replicate analyses of the seawater-equilibrated air and theatmosphere in the marine boundary layer. Seawater was taken continuously fromthe bottom of the ship located ca. 5 m below sea level and diverted into the MRI-shower-type equilibrator (~6 dm3min–1 for 10–12 minutes) where it wasequilibrated with an aliquot of air in a closed circuit of the analyzer including theheadspace of the equilibrator. The atmosphere in the marine boundary layer wastaken continuously from the foremast (ca. 10 m above sea level) and an aliquot(~500 cm3min–1) was introduced into the analyzer for analysis.

The concentration of CO2 is calculated based on the MRI87 scale (Inoue etal., 1995). Primary standard gases of MRI87 scale have been prepared with agravimetric method following the procedure reported by Tanaka et al. (1987).The difference in CO2 concentration scale between WMO mole-fraction andMRI87 has tentatively been determined as Eq. (1):

(xCO2WMO – xCO2

MRI87)/ppm= –0.015 – 0.569·(xCO2

MRI87 – 370)·10–2 + 0.237·(xCO2MRI87 – 370)2·10–4

+ 0.341·(xCO2MRI87 – 370)3·10–6 (271 < xCO

2MRI87/ppm < 405). (1)

Although the difference exceeds 0.4 ppm below 300 ppm, it is insignificant indiscussing the variability of pCO2 in near-surface water.

Corrections were made for the drift of NDIR output voltage and the drift ofCO2 concentration in the working standard gases, if any. Corrections for thetemperature-rise of near-surface water within the inner piping from the seawaterintake to the equilibrator were made using the equation given by Copin-Montegut(1988). The partial pressure of CO2 was calculated from xCO2 by taking thesaturated-water vapor pressure (psw) and the atmospheric pressure (P) intoaccount,

pCO2 = xCO

2·(P – p

sw) (2)

where P is assumed constant at 1 atm.Measurements of TCO2 were made coulometrically (Johnson et al., 1985)

using the automated TCO2 analyzer equipped with a carbon coulometer (UIC Co.,Ltd., model 5011 or 5012) (Ishii et al., 1998) and its improved design (NipponANS Co.). For underway measurements, a portion of seawater taken continuouslyfrom the bottom of the ship was introduced into the pipette (~22 cm3) of theanalyzer twice every 1 or 1.5 hour concurrently with the pCO2 measurement ofsea water. At hydrographic stations we took discrete samples for TCO2 analysisfrom depths using Niskin bottles. These samples were stored in 250 cm3 borosilicateglass bottles with ground-glass stoppers lubricated with Apiezon-L grease afterpoisoning with 0.2 cm3 of saturated mercury(II) chloride solution. Proceduresand a method for calculating TCO2 were basically identical to those described inDOE (1994).

To establish an identical TCO2 concentration scale among cruises and

Variability of Surface Layer CO2 Parameters 65

laboratories, we also analyzed TCO2 in the Certified Reference Material (CRM)provided by Dr. A. Dickson of the Scripps Institution of Oceanography (http://www.mpl.ucsd.edu/people/adickson/CO2_QC/index.html) and in several batchesof the reference seawaters we prepared from western North Pacific oligotrophicwater by a similar method to that of Dickson (1991). The concentration of TCO2in the reference seawater and its stability have been determined by occasionalcalibration with a suite of sodium carbonate solutions (Ishii et al., 1998) andanalyses of CRM. The analytical results of CRM based on our calibration withsodium carbonate solutions usually agreed with its certified value within ±1.5µmol kg–1, but sometimes exceeded ±3.0 µmol kg–1. The systematic error in ourcalibration was corrected if it was statistically significant. During the series ofsample analyses, we also analyzed CRM or our reference seawater typically twiceduring the each run of the coulometric cathode- and anode-solution. The precision(±1σ) of analysis estimated for each cruise ranged between ±1.0 and ±2.1 µmolkg–1.

For the measurement of pH, we introduced a spectrophotometric method(Clayton and Byrne, 1993) using m-cresol purple as an indicator dye. In January1999 during the R/V Mirai MR98K2 cruise, we used a JASCO V-550spectrophotometer with a temperature-controlled (25.00 ± 0.05°C) photometriccell of 10 cm light-path length. Since November 1999 during the R/V MiraiMR99K7 cruise, we have been using an automated pH analyzer (Nippon ANSCo., Ltd.) equipped with a Cary 50 spectrophotometer (Varian Australia PtyLtd.). Samples for pH analysis were collected in the same manner as for TCO2analyses. Perturbation by the addition of indicator-dye solution was corrected forby the empirical method described in DOE (1994). Details of the equipment anda method of correcting for the perturbation by the addition of mercury(II) chloridesolution for discrete samples will be reported elsewhere. Repeatability as deducedfrom the analyses of reference seawater and duplicate samples was within ±0.001.

We calculated TA in near-surface waters from pCO2sw and TCO2, and thatin water columns from TCO2 and pH, with temperature and salinity using thesolubility of CO2 in seawater given by Weiss (1974), dissociation constants ofcarbonic acid given by Roy et al. (1993) and equilibrium constants for other acidsand bases recommended in DOE (1994).

3. VARIABILITY IN SURFACE LAYER CO2 PARAMETERS

3.1 Surface layer CO2 parameters in the warm pool

Longitudinal distributions of pCO2sw and pCO2 in the atmosphere (pCO2air),NTCO2 (TCO2 normalized at S = 35), NTA (TA normalized at S = 35), temperatureand salinity in near-surface water along the equator are shown in Figs. 1 to 4 for8 cruises conducted in boreal winters. We have observed extensive eastwardprogress of the warm pool that extended across the Dateline in January 1994(KY9401), in December 1994 (KY9411), in December 1997 to January 1998(KY9712 and MR97K2), in January/February 2002 (MR02K1), and in January2003 (MR02K6) during the El Niño and non-La Niña periods. As has been

66 M. ISHII et al.

320

360

400

440

480

130E 140E 150E 160E 170E 180 170W 160W

pCO2

pC

O2 /

µa

tm

Jan 1994

(KY9401)

Dec 1994

(KY9411)

1920

1940

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NT

CO

2 /

µm

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-1

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NT

A /

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ol kg

-1

24

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Te

mp

. /

de

gC

33.5

34.0

34.5

35.0

35.5

130E 140E 150E 160E 170E 180 170W 160W

Sal

Sa

linity

Fig. 1. Partial pressure of CO2 in near-surface water (pCO2sw) and in the overlying atmosphere(pCO2air), salinity-normalized (S = 35) total inorganic carbon concentration (NTCO2) and totalalkalinity (NTA), temperature and salinity in near-surface water along the equator in January 1994during the R/V Kaiyo’s KY9401 cruise (triangles) and in December 1994 during the R/V Kaiyo’sKY9411 cruise (plus symbols).

Variability of Surface Layer CO2 Parameters 67

Fig. 2. Partial pressure of CO2 in near-surface water (pCO2sw) and in the overlying atmosphere(pCO2air), salinity-normalized (S = 35) total inorganic carbon concentration (NTCO2) and totalalkalinity (NTA), temperature and salinity in near-surface water along the equator in December 1997to January 1998 during the R/V Kaiyo’s KY9712 cruise and R/V Mirai’s MR97K2 cruise (triangles)and in January 1999 during the R/V Mirai’s MR98K2 cruise (plus symbols).

320

360

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130E 140E 150E 160E 170E 180 170W 160W

pCO2

pC

O2 /

a

tm

Dec 1997

(KY9712)

Jan 1999

(M R98K2)

Jan 1998

(MR97K2)

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NT

CO

2 /

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A /

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130E 140E 150E 160E 170E 180 170W 160W

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Sa

linity

68 M. ISHII et al.

320

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pCO2

pC

O2 /

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tm

Jan . 2001

(MR00K8)N ov-Dec 1999

(MR99K7)

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1940

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NT

CO

2 /

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A /

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mp

. / d

eg

C

33.5

34.0

34.5

35.0

35.5

130E 140E 150E 160E 170E 180 170W 160W

Sa l

Sa

linity

Fig. 3. Partial pressure of CO2 in near-surface water (pCO2sw) and in the overlying atmosphere(pCO2air), salinity-normalized (S = 35) total inorganic carbon concentration (NTCO2) and totalalkalinity (NTA), temperature and salinity in near-surface water along the equator in November-December 1999 during the R/V Mirai’s MR99K7 cruise (triangles) and in January 2001 during theR/V Mirai’s MR00K8 cruise (plus symbols).

Variability of Surface Layer CO2 Parameters 69

320

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pC

O2 /

µa

tm

Dec 2002-Jan 2003

(M R02K6)

Jan-Feb 2002

(M R02K1)

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eg

C

33.5

34.0

34.5

35.0

35.5

130E 140E 150E 160E 170E 180 170W 160W

Sa l

Sa

linity

Fig. 4. Partial pressure of CO2 in near-surface water (pCO2sw) and in the overlying atmosphere(pCO2air), salinity-normalized (S = 35) total inorganic carbon concentration (NTCO2) and totalalkalinity (NTA), temperature and salinity in near-surface water along the equator in January-February 2002 during the R/V Mirai’s MR02K1 cruise (triangles) and in December 2002–January2003 during the R/V Mirai’s MR02K6 cruise (plus symbols).

70 M. ISHII et al.

reported earlier, pCO2sw in the warm pool was mostly within the range –10 µatmand +30 µatm of pCO2air. Concurrent underway measurements of TCO2,temperature and salinity revealed that such a moderate level of pCO2sw under thehighest temperature condition (28.5 < T/°C < 31) was ascribed to the low TCO2(1920 < NTCO2/µmol kg–1 < 1955) and somewhat to the low salinity (33.5 < S <35) in this region. Although the variability is smaller than in the divergence zone,some features of the variability of the CO2 parameters can be identified in thewarm pool.

In the western region and sometimes even in the eastern region in the warmpool, both TCO2 and TA in surface layer varied with salinity, and their salinity-normalized values showed little zonal variation. The consequence of such auniform NTCO2 and NTA was that the zonal variation of pCO2sw in surface layerwas predominantly controlled by the variation of temperature and salinity. Forexample, in December 1997 during the strongest El Niño event on record, near-surface NTCO2 and NTA on the equator between 143°E and 180° showed verylittle variation (NTCO2 = 1947.7 ± 2.9 µmol kg–1, NTA = 2316.6 ± 3.2 µmolkg–1) but T was +0.6°C higher and S was +0.41 higher in the western part of thewarm pool than in the eastern part (Fig. 2). A westward increase in T and Sconcurrently raised pCO2sw to 376.6 ± 3.8 µatm (∆pCO2 = 26.7 µatm) between145°E and 152°E. Similarly, near-surface NTCO2 and NTA between 145°E and179°E in late December 2002 to January 2003 were 1948.3 ± 2.3 µmol kg–1 and2318.0 ± 3.0 µmol kg–1, respectively (Fig. 4). At this time, temperature andsalinity again showed a small westward increase of 0.6°C for T and 0.7 for Swithin the warm pool. These zonal gradients of temperature and salinity resultedin the zonal gradient of pCO2sw at 380.8 ± 3.6 µatm (pCO2sw – pCO2air = ∆pCO2= +22.4 ± 3.8 µatm) in 145°E–152°E and 356.5 ± 6.0 µatm (∆pCO2 = –3.0 ± 6.1µatm) in 170°E–179°E. It has been reported that the warm pool in the westernequatorial Pacific sometimes became a weak source of CO2 (Fushimi, 1987; Ishiiand Inoue, 1995; Inoue et al., 2001). While the increase in surface NTCO2 due tothe erosion of halocline by enhanced turbulence has been suggested at 137°Eduring the 1982/83 El Niño (Fushimi, 1987) and at 3°–5°N of 144°E during the1997/98 El Niño (data not shown), the zonal change in pCO2sw to the east of145°E in the warm pool can primarily be attributed to the change in temperatureand salinity.

In the eastern part of the warm pool, however, a slight subsaturation of CO2was also often observed. In this region, while temperature was the highest (T >30°C) and variation in NTA was insignificant, the low NTCO2 (1920 < NTCO2/µmol kg–1 < 1935) and low salinity (~34) were responsible for the lower pCO2sw.For example, in January 1994 during the non-El Niño period, lower pCO2swvalues due to low NTCO2 were observed between 155°E and 170°E (Fig. 1). Inthis region, near-surface NTCO2 (=1931 ± 3 µmol kg–1) was 12 µmol kg–1 lowerthan the average NTCO2 in the western region (1943 ± 5 µmol kg–1 in 145°E–152°E) and was 16 µmol kg–1 lower than that in the warm pool in December 1997(1947 ± 3 µmol kg–1) (Fig. 3), while no variability in NTA was detected either inspace or time (2315 ± 4 µmol kg–1 in 155°E–170°E in January 1994; 2318 ± 5

Variability of Surface Layer CO2 Parameters 71

19

20

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19

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72 M. ISHII et al.

µmol kg–1 in 145°E–152°E in January 1994; 2317 ± 4 µmol kg–1 in 145°E–179°Ein December 1997). Consequently, the eastern region in the warm pool had beensubsaturated with CO2. Similarly, a distinct NTCO2 minimum was observedbetween 180° and 175°W just to the west of the salinity front in January 2003(Fig. 4). The lowest NTCO2 (1936 µmol kg–1) at 178°43′ W was 11 µmol kg–1

lower than the average NTCO2 between 160°E and 179°E. Such a lower pCO2swin the warm pool that was ascribed to the lower NTCO2 has been observed in allphases of the ENSO; in December 1994 during the El Niño period: 1935.7 ± 2.1µmol kg–1 between 170°W and 166°W (Fig. 1); in January 1999 during the LaNiña period: 1928 µmol kg–1 at 153°16′ E (Fig. 2); and in January 2002 during thenon-El Niño period: 1938.9 ± 3.5 µmol kg–1 at around 160°E (Fig. 4).

The vertical profiles of NTCO2 and hydrographic properties in the warmpool show that the region of such a lower NTCO2 about 1920 to 1940 µmolkg–1 with the tendency of CO2 subsaturation coincided with the region where thebarrier layer has been formed (Fig. 5). Although nitrate had been completelydepleted in the isothermal layer, NTCO2 and trace amounts of phosphate increasedwith depth within and above the barrier layers, suggesting that the net biologicalCO2 uptake is responsible for the lower NTCO2 in this region. It is likely that thenet biological uptake of CO2 possibly by the nitrogen fixers appeared as a resultof the hampered vertical mixing at the site of barrier layer formation.

3.2 Surface layer CO2 parameters in the equatorial divergence zone

We have observed the western part of the equatorial divergence zone,extending more than 20 degrees to the west of 160°W in January 1994 (KY9401;Fig. 1), in January 1996 (KY9511; data not shown), in January 1999 (MR98K2;Fig. 2), in November/December 1999 (MR99K7; Fig. 3) and in January 2001(MR00K8; Fig. 3) during the non-El Niño and the La Niña periods.

Concurrent measurements of TCO2 have shown that the high pCO2sw (385to 465 µatm) in this zone is ascribed to the moderately high NTCO2 (1960 to 2030µmol kg–1), salinity (35.0 to 35.5) and temperature (24.5 to 29.3°C), and low NTA(ca. 2320 µmol kg–1). In general, near-surface NTCO2 decreased to the west withthe elevation of temperature, but there was little zonal variation in salinity and

Table 2. Zonal gradients of temperature (∂T/∂x), NTCO2 (∂C/∂x) and NTA (∂NTA/∂x) in the westernequatorial divergence zone.

Month Year Cruise Longitude ∂T/∂x ∂C/∂x ∂NTA/∂x n°C degree–1 µmol kg–1deg–1 µmol kg–1deg–1

Jan 1994 KY9401 175°E–172°W –0.18 ± 0.01 +3.72 ± 0.11 +0.29 ± 0.13 55Jan 1999 MR98K2 164°E–170°W –0.14 ± 0.00 +2.33 ± 0.14 +0.14 ± 0.12 12Dec 1999 MR99K7 160°E–170°W –0.08 ± 0.00 +1.14 ± 0.16 +0.02 ± 0.02 164Jan 2001 MR00K8 160°E–160°W –0.09 ± 0.00 +1.34 ± 0.02 –0.05 ± 0.01 254Jan 2002 MR02K1 177°W–160°W –0.07 ± 0.00 +0.97 ± 0.07 –0.07 ± 0.04 79Jan 2003 MR02K6 170°W–160°W –0.12 + 0.00 +3.00 + 0.10 +0.15 ± 0.08 54

Variability of Surface Layer CO2 Parameters 73

NTA. The zonal gradient of NTCO2 as well as temperature were changing yearby year (Table 2). Although NTCO2 and temperature in the near-surface watersometimes fluctuated according to the effect of the Tropical Instability Wave (seedata from MR99K7 in Fig. 3, for example), mean zonal gradient of near-surfaceNTCO2 in the western divergence zone varied more than three-fold from 0.97 ±0.07 µmol kg–1degree–1 in January 2002 (Fig. 4) to 3.72 ± 0.11 µmol

Fig. 6. Vertical sections of (a) NTCO2, (b) pH, (c) NTA, (d) temperature and (e) salinity in the top200 m along the equator in the western and central equatorial Pacific in January 1999 during the R/V Mirai’s MR98K2 cruise.

74 M. ISHII et al.

kg–1degree–1 in January 1994 (Fig. 1).These zonal gradients of NTCO2 and temperature were seen over the surface

layer above the thermocline (Fig. 6). This suggests that the upwelling from thethermocline is weaker in the west than in the east, and biological uptake and CO2evasion to the atmosphere is removing TCO2 while solar energy input warms thesurface layer in the course of the westward advection in the South EquatorialCurrent. The westward decreases of nitrate, phosphate and silicic acid above thethermocline (data not shown) also suggest the zonal difference in the upwardtransport of the thermocline-water and the importance of net biological uptake forthe zonal TCO2 gradient in the surface layer. By contrast, NTA showed only asmall gradient, both vertically and horizontally, in the top 300 m of the watercolumn across the thermocline (see Fig. 6 and Table 2). It is consistent with theprevious results that the equatorial divergence does not accompany a largeincrease in surface NTA (Archer et al., 1996; Millero et al., 1998). The smallNTA gradient implies that the activities of calcareous algae such ascoccolithophorid are playing rather minor role in perturbing the oceanic CO2system in this zone. The role of biological activities and sea-to-air CO2 evasionin reducing the surface NTCO2 is quantified in Section 5.

3.3 Boundary between the warm pool and the equatorial divergence zone

The boundary between the warm pool and the divergence zone in theequatorial Pacific is often recognized by a marked change in pCO2sw (from 360–390 µatm to 390–430 µatm) within a few degrees in longitude (Inoue et al., 1996;Le Borgne et al., 2002). It is also marked by a zonal salinity front (34.1–34.4 to35–35.5) but not by temperature. We observed the boundary zone with theunderway measurements of pCO2sw and TCO2 as well as temperature and salinityin five R/V Mirai’s cruises took place in January 1999 (MR98K2), in November/December 1999 (MR99K7), in January 2001 (MR00K8), in January 2002(MR02K1), and in January 2003 (MR02K6) (see Figs. 2 to 4). These observationsrevealed that the boundary zone is also a zone of abrupt change in NTCO2, andthe abrupt change in pCO2sw is therefore ascribed mainly to the changes inNTCO2 and partly to the change in salinity. However, these underwaymeasurements with fine spatial resolution further revealed that the boundary zonecould be categorized into two domains: the western boundary domain with alarger salinity change and the eastern boundary domain with a larger NTCO2change.

In the western boundary domain, near-surface NTCO2 increased from 1940–1945 µmol kg–1 to 1960 µmol kg–1 on moving from the west to the east, whilepCO2sw also increased from 360–370 µatm to 380–400 µatm and salinity fromless than 34.4 to 35.0–35.3. Near-surface temperature varied between 29.6 and30.2°C. The concentration of phosphate in the surface layer also changed fromless than 0.16 µmol kg–1 to about 0.22 µmol kg–1 but nitrate was completelydepleted throughout this domain. In the eastern boundary domain, on the otherhand, near-surface NTCO2 changed from 1960 µmol kg–1 to 1970–1980 µmol

Variability of Surface Layer CO2 Parameters 75

kg–1, pCO2sw changed from 390–400 µatm to 415–430 µatm, and salinitychanged from 35.1–35.2 to 35.3–35.4. The temperature near the surface had atendency to decrease 0.3–1.0°C to the east from 29.6–29.8°C to 28.6–29.5°C. Thechange in salinity was much smaller in the eastern boundary domain (0.15–0.2)than in the western one (0.7–1.0), but the change in pCO2sw (15–40 µatm) wascomparable. In near-surface water of the eastern boundary domain, concentrationsof nutrients also changed from 0.22 µmol kg–1 to 0.29–0.38 µmol kg–1 forphosphate and from complete depletion to 1.5–3.0 µmol kg–1 for nitrate.

It turned out that the zonal variation in the near-surface CO2 parameters inthe boundary zone could be ascribed to the zonal extent of these western andeastern domains. In January 1999 during the La Niña period, the abrupt changein pCO2sw that corresponded to the western boundary domain was clearlyobserved at 155°E–158°E (Fig. 2). However, the eastern boundary domain had abroad zonal extent (158°E–170°E) so that pCO2sw and NTCO2 increased rathergradually to the east, and depletion of nitrate was observed in the surface layer ofthe hydrographic stations located in the eastern boundary domain. It is likely thata similar situation had also occurred to the east of the dateline in September/October 1994 during the El Niño period (Rodier et al., 2000). In December 1999,again in the La Niña period, the western boundary domain and the easternboundary domain were observed at 150°E–152°E and 160°E–162°E, respectively(Fig. 3). They had been separated from each other so that pCO2sw in the boundaryzone increased stepwise from the west to the east. By contrast, the transition fromthe warm pool to the divergence zone was less pronounced in January 2001 (Fig.3). At this time, both the western boundary domain (148°E–158°E) and theeastern boundary domain (158°E–165°E) had a rather broad zonal extent, so thatpCO2sw and NTCO2 as well as salinity changed less abruptly. In January 2002,on the other hand, the abrupt zonal change in pCO2sw was again observed (Fig.4). Although the western boundary domain had a rather broad extent (168°E–180°), the abrupt change in near-surface NTCO2 in the eastern boundary domain(180°–176°W) caused an abrupt change in pCO2sw at the boundary zone.Similarly, in January 2003, the western boundary domain (176°W–171°W) andin particular the eastern boundary domain (171°W–170°W) had limited zonalextent and these two domains were connected to each other (Fig. 4). Consequently,sharp zonal discontinuities of both pCO2sw and salinity appeared in the narrowboundary zone between warm pool and the divergence zone.

The key processes controlling the zonal distribution of CO2 parameters in thewestern boundary domain are unclear. Linear but different relationships betweensurface NTCO2 and salinity (Fig. 7) for each cruise suggest that the mixing of thedifferent waters from the western warm pool and the eastern boundary domain isdetermining the zonal distribution of the CO2 parameters as well as theconcentration of phosphate in the western boundary zone. Therefore, the westernboundary domain could be regarded as a hydrological frontal zone.

On the other hand, since the change in salinity is small, it is hypothesized thatthe eastern boundary domain is formed through biogeochemical processes andfluctuations of vertical and horizontal TCO2 transport in the equatorial divergence.

76 M. ISHII et al.

When the eastern boundary domain had rather broad zonal extent (e.g. in January1999, Fig. 2; in December 1999, Fig. 3; in January 2001, Fig. 3), it is ratherrecognized as being a part of the equatorial divergence zone. As mentioned in theprevious section, the large-scale feature of the westward decrease in NTCO2 inthe divergence zone is generated by the stronger divergence in the east. However,the horizontal transport of NTCO2, viz. the velocity of SEC, and the rate of non-conservative CO2 decrease due to biological uptake and the evasion to theatmosphere should also determine the horizontal NTCO2 gradient (see Section 5).The steeper horizontal NTCO2 gradient is expected when the westward advectionis slower and the net biological CO2 uptake and/or the net sea-to-air CO2 evasionto the atmosphere is faster.

In addition, biological uptake rates of CO2 and nitrate could overcome theirupward and westward transport rates during the cessation of vertical and horizontaltransport. In this situation, as has been discussed by Rodier et al. (2000) forsurface-layer nitrate concentration, surface-layer NTCO2 near the western end ofthe equatorial divergence zone could also decrease until nutrient depletes. Theconsequence of such a cessation of vertical and horizontal TCO2 transport is thatthe western and eastern boundary domains, or salinity front and boundary ofnitrate-depletion, separate from each other. Conversely, when the vertical and/orhorizontal transport is enhanced and overcomes the rates of biological uptake andCO2 emission to the atmosphere, NTCO2 and nitrate in surface layer wouldincreases in the western end of the equatorial divergence. In this situation, theeastern boundary domain approaches the western boundary domain and thesedomains are finally connected to each other. The consequence is that the gap in

Fig. 7. Property-property plots of salinity and NTCO2 in near-surface layer of the boundary zonebetween the warm pool and the equatorial divergence zone.

1930

1940

1950

1960

1970

1980

34 34.5 35 35.5

Jan 1999

Oct 1999

Jan 2001

Jan 2002

Jan 2003

NT

CO

2 /

m

ol k

g-1

Salinity

western boundary

eastern boundary

Variability of Surface Layer CO2 Parameters 77

surface NTCO2 as well as nitrate grows and the larger zonal change in pCO2swemerges at the boundary zone.

4. MAPPING OCEANIC CO2 PARAMETERS AND SEA-TO-AIR CO2 FLUX

4.1 Empirical relationship between NTCO2 and temperature in surface layer

A total of 1037 data of near-surface NTCO2 in the equatorial Pacific between135°E–95°W, 10°S–5°N that have been collected in 1990s and in early 2000s

1920

1940

1960

1980

2000

2020

2040

2060

2080

20 22 24 26 28 30

NT

CO

2 /

µm

ol k

g-1

Temperature / degC

(a) NTCO2 vs Temp

2240

2260

2280

2300

2320

2340

2360

2380

2400

20 22 24 26 28 30

NT

A /

µm

ol k

g-1

Temperature / degC

(b) NTA vs Temp

Fig. 8. Composite plots of (a) temperature vs. NTCO2 and (b) temperature vs. NTA in near-surfacewater in the equatorial Pacific between 135°E–95°W, 10°S–5°N in the period from April 1990 toJanuary 2003. Offset of +15.0 µmol kg–1 has been rather added to the data of titration alkalinity.

78 M. ISHII et al.

including those listed in Table 1 were plotted against temperature (Fig. 8(a)).Data from underway TCO2 measurements in 14 cruises were averaged into 1°longitude × 1° latitude pixels for each cruise. This averaging process reduced thetotal of 5069 underway TCO2 data to 835 pixel data (16.5%). Data from JGOFS/WOCE programs in early to mid-1990s (EqPac cruises in 1992, WOCE sectionsP9, P10, P13, P14, P16, P17, and National Oceanic and AtmosphericAdministration’s CGC90 and CGC96 cruises; http://www.aoml.noaa.gov/ocd/oaces/bottle_data.html) were also used for the analysis. In these TCO2measurements, the Certified Reference Material provided by Dr. A. G. Dicksonhas been used for quality control, and overall accuracy of TCO2 data taken byvarious laboratories in these programs has been estimated as ca. 3 µmol kg–1

(Lamb et al., 2002). In this analysis, data from the eastern equatorial Pacific closeto the coast of America (<95°W) were not used, because they showed a quitedifferent feature of surface NTCO2 vs. temperature, due probably to the effect ofthe Peru Current, which entrains the water from the Peruvian upwelling systemwith extremely high pCO2sw (>700 µatm; Feely et al., 2002).

In the divergence zone, near-surface NTCO2 can be expressed empirically asa function of temperature and the year of observation:

NTCO2/µmol kg–1 = 1945.5 – 20.85·(T/°C – 30) – 1.660·(T/°C – 30)2

–0.0909·(T/°C – 30)3 + 0.91·(Year – 1995). (3)

Standard error (±1σ) of the fitting was ±5.9 µmol kg–1. This is larger than theprecision of our underway TCO2 measurements (<±2.1 µmol kg–1) as well as thatin the Pacific CO2 survey programs. However, given the different processes thatalter TCO2 and temperature after the thermocline water upwelled to the surfacelayer, it is rather surprising that the plots of near-surface NTCO2 vs. temperaturein the divergence zone converged within the range of ±5.9 µmol kg–1 (±1σ)around the fitting curve expressed as Eq. (3), regardless of the sampling site andthe ENSO phase.

On the other hand, in the warm pool and in the zone of the North EquatorialCounter Current (NECC), near-surface NTCO2 can also be expressed as afunction of temperature and the year of observation:

NTCO2/µmol kg–1 = 1938.3 – 8.80·(T/°C – 30) + 0.74·(Year – 1995). (4)

Standard error (±1σ) of the fitting was ±5.2 µmol kg–1. As described in Subsection3.1, a lower near-surface NTCO2 was observed above the barrier layer wheresurface temperature tended to increase. Such an effect of the barrier layerformation is one of the factors that contribute to the negative slope in Eq. (4). Inaddition, turbulent mixing in the western equatorial Pacific to the west of 140°Ecaused an elevation of near-surface NTCO2 and a decrease in sea surfacetemperature. A trend that the near-surface NTCO2 increases with the decrease oftemperature is also clear in the zone of the NECC to the north of the divergencezone.

Variability of Surface Layer CO2 Parameters 79

These relationships between near-surface NTCO2 and temperature agreedfairly well with those reported by Loukos et al. (2000) who used total of 435 dataobtained in the early 1990s between 10°N–10°S, 170°W–96°W. The meandifference between Eq. (3) for the equatorial divergence and the relationshipgiven by Loukos et al. (2000) is +7.3 µmol kg–1. However, the uncertainty hasbeen significantly improved from ±9 µmol kg–1 in Loukos et al. (2000) to < ±6µmol kg–1 in Eqs. (3) and (4) in the western and central equatorial Pacific byapplying different equations to the warm pool and to the divergence zone.

It should also be noted that the standard error of the regression analyses ofnear-surface NTCO2 with temperature is improved by approximately ±1 µmolkg–1 when the long-term change in the near-surface NTCO2 is taken into accountby incorporating the term for the year of observation into Eqs. (3) and (4). Thecoefficient of the long-term NTCO2 change, +0.91 ± 0.09 µmol kg–1 year–1 for thedivergence zone and +0.74 ± 0.07 µmol kg–1 year–1 for the warm pool, correspondsto the pCO2sw increase rate of 2.0 ± 0.2 µatm year–1 and 1.2 ± 0.1 µatm year–1,respectively. These values are comparable to the globally-averaged atmosphericCO2 increase rate of 1.5 ppm year–1 in the period from 1979 to 2001 (NOAA,2002) and a long-term pCO2sw increase rate of 1.27 ± 0.18 µatm year–1 in theperiod from 1961 to 1996 (Feely et al., 1999) in the divergence zone of theequatorial Pacific. These coefficients may possibly be indicative of the long-termincrease in NTCO2 due to anthropogenic CO2 input. This will be discussed inmore detail elsewhere.

The relationship between near-surface NTA and temperature was alsoexamined (Fig. 8(b)). The mean near-surface NTA calculated from pCO2sw,TCO2, T, and S in the western and central equatorial Pacific (T ≥ 25.0°C) usingdissociation constants given by Roy et al. (1993) (2317.7 ± 3.2 µmol kg–1, n =730) was +15.0 µmol kg–1 larger than the surface NTA measured by titration(2302.7 ± 4.6 µmol kg–1, n = 195). Therefore, the offset of +15.0 µmol kg–1 wasadded to the data of titration alkalinity taken in the equatorial Pacific includingthe eastern region where T ≤ 25°C, so that the pCO2sw calculated from TCO2 andTA became consistent with the measured pCO2sw. Consequently, near-surfaceNTA between 135°E–90°W, 10°S–5°N was expressed as

NTA/µmol kg–1 = 2317.4 – 0.08·(T/°C – 30) + 0.081·(T/°C – 30)2. (5)

Standard error (±1σ) of the fitting was ±3.6 µmol kg–1. The divergence zone andthe warm pool were not discriminated since the difference in near-surface NTAbetween these domains was not significant. A small temperature dependence wasobtained from the data of titration alkalinity obtained in the eastern divergencezone where T < 25°C.

4.2 Mapping of surface NTCO2, pCO2sw, and sea-to-air CO2 flux in the equatorialPacific in 1990–2000

In order to obtain monthly fields of surface NTCO2 in the equatorial Pacific

80 M. ISHII et al.

between 135°E–95°W, 10°S–5°N for the period 1990–2000, the empiricalrelationships between near-surface NTCO2 and temperature given as Eqs. (3) and(4) were combined with the monthly global sea surface temperature dataset(JMA, 1991). The longitude of the boundary that delimitates the warm pool andthe equatorial divergence zone was approximated by the empirical formula of thelongitude of pCO2sw boundary that is expressed as a function of a 5-monthsrunning mean of SOI:

Fig. 9. Time-longitude distributions of (a) surface NTCO2 and (b) ∆pCO2 (=pCO2sw – pCO2air) inthe equatorial Pacific (Mean of 0°–5°S).

140E 160E 180 160W 140W 120W 100W

140E 160E 180 160W 140W 120W 100W

Longitude

(a) NTCO2(umol/kg)

(b) del_pCO2(uatm)

2001

2000

1999

1998

1997

1996

1995

1994

1993

1992

1991

1990

Year

2001

2000

1999

1998

1997

1996

1995

1994

1993

1992

1991

1990

Year

Longitude

Variability of Surface Layer CO2 Parameters 81

Longitude of pCO2sw boundary (°E) = 171.3 – 17.4·SOI + 1.63·SOI2

(1σ = 5.4° in longitude, n = 14). (6)

The latitude of the boundary between the equatorial divergence zone (SouthEquatorial Current) and the NECC to the north also varies in time and longitudewith the passage of the Tropical Instability Wave. However, it was assumedconstant at 4°N in this estimate. Similarly, the boundary between the equatorialdivergence zone and the subtropical current system to the south was also assumedconstant at 8°S. Another formula from a regression analysis: NTCO2/µmol kg–1

= 1927.6 – 11.84·(T/°C – 30) + 0.69·(Year – 1995) was applied for the southernsubtropics.

Time-longitude distribution of surface NTCO2 averaged between 5°S andthe equator (Fig. 9) demonstrates that the high-NTCO2 surface water (>1980µmol kg–1) of the divergence zone retreated to the east of 170°W during El Niñoperiods. During the strongest 1997/1998 El Niño, such a high-NTCO2 water isinferred to have disappeared from the equatorial Pacific. In contrast, during thenon-El Niño period 1998–2001, the high-NTCO2 surface water progressed to thewest across the Dateline, reaching 160°E.

Similarly, monthly fields of surface NTA were estimated, and then monthlyfields of temperature, NTCO2 and NTA were combined to provide the monthlyfields of pCO2sw. Mean surface salinity observed in the divergence zone (35.16)and in the warm pool (34.39) was also used for the calculation. Uncertainties inthe calculated pCO2sw of ±21 to ±30 µatm (±2σ) depending on the temperaturein the divergence zone and ±19 µatm (±2σ) in the warm pool were derived fromthe uncertainty of surface NTCO2. In reality, an additional uncertainty of ca. ±10µatm due to the variability in salinity should be also taken into account.

Monthly net sea-to-air CO2 flux in the equatorial Pacific was next estimatedwith Eq. (7) from the monthly fields of pCO2sw and pCO2air, CO2 solubility inseawater (s) (Weiss, 1974), and CO2 transfer velocity (k) parameterized withlong-term wind speed given by Wanninkhof (1992):

F = k·s·(pCO2sw – pCO2air) = k·s·∆pCO2k/cm h–1 = 0.39·U

av2·(Sc

20/Sc)1/2 (7)

where Uav is one-month averaged wind speed (m sec–1), Sc20 and Sc are theSchmidt number for CO2 at 20°C and at sea surface temperature (Jahne et al.,1984). For the calculation of pCO2air, CO2 concentrations at Christmas Island,Kiribati (2°00′ N, 157°18′ W) (http://www.cmdl.noaa.gov/info/ftpdata.html)were interpolated with respect to time, and values were applied to the entireequatorial Pacific. Data of one-month averaged wind speed were acquired fromthe National Centers for Environmental Prediction/National Center forAtmospheric Research reanalysis wind fields (Kalnay et al., 1996).

The time-longitude distribution of sea-to-air CO2 flux averaged between 5°Sand the equator from the year 1990 to 2000 (Fig. 10) demonstrates that the regionof large sea-to-air CO2 flux (>6 mmol m–2day–1) had been confined to the east of

82 M. ISHII et al.

160°W during the El Niño period in 1991–1994 and even disappeared during thestrongest 1997/1998 El Niño period. By contrast, the region of the large sea-to-air CO2 flux extended westward to around 165°E during the La Niña period in1998–2000 with the westward progress of high-pCO2sw domain. In this period,the flux was intensified (>12 mmol m–2day–1) at around 150°W due to theenhanced wind speed in this region.

Net sea-to-air CO2 flux integrated over the equatorial Pacific between135°E–95°W, 10°S–5°N is summarized in Fig. 11. When the uncertainty in theflux derived from ±2σ of surface NTCO2 is taken into account, the sea-to-air CO2flux in the period 1990–2000 was +0.38 ± 0.16 PgC year–1 on average, varyingfrom +0.30 ± 0.15 PgC year–1 in 1997 to +0.46 ± 0.16 PgC year–1 in 1999 and2000. The monthly flux exhibited greater variability from +0.17 ± 0.11 PgCyear–1 in April 1998 during the strongest El Niño period to +0.57 ± 0.23 PgCyear–1 in August 1995 during the non-El Niño period and +0.56 ± 0.21 PgCyear–1 in December 1999 and August 2000 during the La Niña period. Given theerror in these estimates, the temporal variation in the integrated net sea-to-air fluxwould not exceed the range of +0.04 PgC year–1 in the strong El Niño period to+0.80 PgC year–1 in the La Niña period during the year 1990–2000.

These results are in good agreement with the estimates from the shipboardpCO2sw surveys: between +0.3 ± 0.2 PgC year–1 in 1992 and +0.9 ± 0.6 PgCyear–1 in 1996 (Feely et al., 1995, 1999), that derived from differentparameterizations of TCO2 and TA: +0.5 ± 0.3 PgC year–1 with anomalies of0.4 ± 0.2 PgC year–1 during 1982–1993 (Loukos et al., 2000), and that from theinterpolation of fCO2sw-temperature relationship combined with satellite

Fig. 10. Time-longitude distributions of net sea-to-air CO2 flux in the equatorial Pacific (mean of0°–5°S).

Variability of Surface Layer CO2 Parameters 83

temperature data: +0.14 ± 0.1 PgC year–1 in the strong El Niño periods and+0.43 ± 0.1 PgC year–1 in the La Niña periods (Cosca et al., 2003). Meanintegrated CO2 flux from the 5°S–0° zonal band in the period 1990–2000(+0.16 ± 0.06 PgC year–1) was also comparable to that estimated for the period1985–1997 from the empirical relationship of pCO2sw vs. temperature (+0.18 ±0.06 PgC year–1) (Boutin et al., 1999). The somewhat smaller flux in our estimatecould be attributed to our exclusion of the effect from the Peruvian upwellingsystem in the eastern equatorial Pacific adjacent to the coast of America. A lessadequate parameterization of surface NTCO2 in the eastern region due to smallerdataset in the colder region (<25°C) (see Fig. 8(a)) used for the estimate ofpCO2sw (see Fig. 9(b)) and different dataset of wind used in the calculation offlux as well as the difference in the period may also be responsible.

Net sea-to-air CO2 flux integrated over the warm pool was usually very small(+0.02 ± 0.03 PgC year–1 on the average). However, it had a tendency to increaseto +0.06 ± 0.01 PgC year–1 during El Niño periods. Although the observedelevation of pCO2sw in the western equatorial Pacific during the El Niño periods(see Figs. 2 and 4, and Inoue et al., 2001) are not properly reproduced in Fig. 11,the stronger westerly, often exceeding 8 m sec–1, resulted in the increase of theflux estimate in this region.

5. ROLE OF BIOLOGICAL ACTIVITIES IN CHANGINGTHE SURFACE LAYER CO2 PARAMETERS

The variability in the surface layer CO2 parameters in the equatorial Pacific,as shown in the previous sections, is controlled by a variety of processes: the winddriven circulation that supplies TCO2-rich thermocline water vertically and

Fig. 11. Integrated sea-to-air CO2 flux from the equatorial Pacific between 135°E–95°W, 5°N–10°S(thick line), and the integrated sea-to-air CO2 flux from the warm pool (dotted line). Thin lines showthe ranges of uncertainty that was derived from the residuals (±2σ) of NTCO2 from Eqs. (3) and (4).

-0.1

0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.9

90 91 92 93 94 95 96 97 98 99 00

CO

2 f

lux /

P

gC

y-1

Year

84 M. ISHII et al.

horizontally, net biological uptake of CO2 and net formation of sinking particulatecarbonate in surface layer, and evasion of CO2 to the atmosphere. In thedivergence zone, these processes change the surface layer CO2 parameters morevigorously than in the warm pool.

On the assumption that the rate of advective transport is much larger than theeddy diffusion in the divergence zone, the change in NTCO2 with time (∂C/∂t) inthe surface layer of this zone is approximated as

∂C/∂t = –u(∂C/∂x) – v(∂C/∂y) – w(∂C/∂z) – (FNCP

+ FPIC

+ Fgas

)·zML

–1 (8)

where FNCP represents the rate of net biological CO2 uptake, FPIC the flux ofsinking particulate carbonate, Fgas the net sea-to-air CO2 flux, and zML thethickness of the mixed layer. The net biological CO2 uptake rate, FNCP, isequivalent to the rate of net community production (NCP), which is defined asprimary production minus respiration by all the autotrophic and heterotrophicorganisms present in the community (Codispoti et al., 1986; Minas et al., 1986).

In principle, with the input of the temporal and spatial variation of NTCO2and sea-to-air flux of CO2 (see Figs. 10(a) and 11), the sum of net biological CO2uptake and particulate carbonate flux, FNCP + FPIC, is calculated from Eq. (9):

FNCP

+ FPIC

= –{(∂C/∂t) + u(∂C/∂x) + v(∂C/∂y) + w(∂C/∂z)}·zML

– Fgas

. (9)

Data of horizontal velocity, u and v, are also available from the TAO buoysdeployed over the equatorial Pacific (http://www.pmel.noaa.gov/tao/).

However, FNCP + FPIC thus estimated strongly depends on the choice of thevelocity of advection and subject to a large uncertainty. In the top 100 m at 170°Win 1990s, for example, even monthly average values of zonal velocity u exhibitedlarge variability with depth and time, ranging from –70 to +130 cm sec–1. In thewestern region, within the divergence zone, the zonal gradient of near-surfaceNTCO2, (∂C/∂x), ranged between +1.0 and +3.7 µmol kg–1degree–1 (see Table 2).Therefore, the uncertainty of 5% in u, viz. ±10 cm sec–1, that might be bestattained in estimating the monthly mean of u, results in an uncertainty of ±8 to ±30mmol m–2day–1 in –u(∂C/∂x)·zML. These values are comparable to or larger thanthe –(∂C/∂t)·zML of (15 mmol m–2day–1 and Fgas of less than 15 mmol m–2day–1.Accordingly, it is unlikely that FNCP + FPIC could be estimated within a satisfactoryuncertainty with this method.

The role of biological activity in changing the CO2 parameters in the surfacelayer could also be evaluated with the concentration of a surface-layer nutrienttaken into account. Concentrations of nutrients such as nitrate and phosphate alsodecrease on moving from the east to the west with the decrease of NTCO2, andthere is a linear relationship between concentrations of nutrients and NTCO2 inthe surface layer of the equatorial divergence zone (Fig. 12). On the assumptionthat ∂C/∂t = 0 on the average and the longitudinal transport dominates over themeridional and vertical transport in the western region of the divergence zonewhere the thermocline is deeper, the zonal gradient of NTCO2 (∂C/∂x), nitrate

Variability of Surface Layer CO2 Parameters 85

(∂N/∂x), and phosphate (∂P/∂x) in the surface layer is expressed as Eqs. (10) to(12), respectively:

∂ ∂ = + +{ } ⋅ ⋅ ( )− −C x F F F u z/ NCP PIC gas ML1 1 10

∂ ∂ = ⋅( ) ⋅ ⋅ ( )− − −N x F u z/ NCP NCP MLC / N 1 1 1 11

Fig. 12. Relationships between NTCO2 corrected for the particulate carbonate formation and (a)salinity-normalized nitrate + nitrite, and (b) salinity-normalized phosphate, in the surface layer ofthe equatorial divergence zone to the west of 170°W in Jan. 1999 (�) and in Dec. 1999 (�), and tothe west of 160°W in Jan. 2001 (+).

1950

1960

1970

1980

1990

2000

2010

2020

2030

0 1 2 3 4 5 6 7 8

NT

CO

2-0

.5(p

otN

TA

-23

17

) /

mo

l kg

-1

salinity-normalized NO3+NO

2 / mol kg

-1

1950

1960

1970

1980

1990

2000

2010

2020

2030

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7

NT

CO

2-0

.5(p

otN

TA

-23

17

) /

mo

l kg

-1

salinity-normalized PO4 / mol kg

-1

86 M. ISHII et al.

∂ ∂ = ⋅( ) ⋅ ⋅ ( )− − −P x F u z/ NCP NCP MLC / P 1 1 1 12

where (C/N)NCP and (C/P)NCP denote the molar stoichiometric ratio of the netbiological uptake of CO2 to nitrate, and CO2 to phosphate, respectively, and udenotes the mean zonal advection velocity. On the other hand, the change in thepotential alkalinity, pot.NTA = (TA + [NO3

–])35/S (Brewer and Goldman, 1976),accounts for the change in the mass balance of acid and base other than biologicaluptake of nitric acid, and is thus indicative of the formation and dissolution ofparticulate carbonate. Like Eqs. (10) to (12), its zonal gradient (∂A/∂x) isexpressed as Eq. (13):

∂ ∂ = ⋅ ⋅ ⋅ ( )− −A x F u z/ .2 131 1PIC ML

From Eqs. (10), (11) and (13), the ratio of zonal NTCO2 gradient to zonal nitrategradient is given as

∂C/∂x·(∂N/∂x)–1 = 0.5·∂A/∂x·(∂N/∂x)–1 + {1 + Fgas

/FNCP

}·(C/N)NCP

. (14)

The ratio of CO2 evasion to NCP, Fgas/FNCP, is thus expressed as

Fgas

/FNCP

= [{∂C/∂x – 0.5·∂A/∂x}·(∂N/∂x)–1·(C/N)NCP

–1] – 1. (15)

The C/N stoichiometry of assimilation in the net community production, (C/N)NCP, is one of the critical factors in calculating the FGAS/FNCP. Since organicmatter assimilated from CO2 and nitrate is transformed into either sinkingparticulate organic matter or suspended and dissolved organic matter, (C/N)NCPis given as

(C/N)NCP = (FPOCsink + FTOCacm)·(FPONsink + FTONacm)–1

= FPOCsink

(1 + FTOCacm

/FPOCsink

)·{FPONsink

·(1 + FTONacm

/FPONsink

)}–1 (16)

where FPOCsink and FPONsink are the vertical flux of organic carbon and organicnitrogen, respectively, in the sinking particulate organic matter, and FTOCacm andFTONacm is the accumulation rate of organic carbon and organic nitrogen,respectively, in either dissolved and suspended organic matter in the surfacelayer. A term in Eq. (16), FTONacm/FPONsink, is further expanded to

FTONacm

/FPONsink

= (FTONacm

/FTOCacm

)·(FTOCacm

/FPOCsink

)·(FPOCsink

/FPONsink

). (17)

Hansell et al. (1997) quantified the importance of FTOCacm from the property-property plots of NTCO2 vs. concentration of organic carbon in dissolved andsuspended organic matter in the equatorial divergence zone of the central Pacific(150°W) in November 1994. They concluded that the formation of dissolved andsuspended organic carbon that is resident in the surface layer accounted for ca.

Variability of Surface Layer CO2 Parameters 87

Tab

le 3

. R

atio

s of

zon

al g

radi

ents

of

NT

CO

2 –

0.5p

ot.N

TA

to n

itra

te +

nit

rite

and

NT

CO

2 to

NT

A,

rati

os a

nd e

stim

ates

of

flux

es i

n th

e w

este

rn d

iver

genc

e zo

ne i

n th

e eq

uato

rial

Pac

ific

obs

erve

d in

La

Niñ

a pe

riod

199

9 –20

01.

(a) D

ate

and

long

itud

e w

hen

and

whe

re t

he e

quat

oria

l di

verg

ence

zon

e w

as o

bser

ved.

(b) C

alcu

late

d fr

om ∆

pCO

2 an

d in

sit

u w

ind

spee

d w

ith

CO

2 tr

ansf

er v

eloc

ity

para

met

eriz

ed w

ith

shor

t-te

rm w

ind

spee

d gi

ven

by W

anni

nkho

f (1

992)

for

the

z one

be t

we e

n th

e bo

unda

ry o

f w

a rm

poo

l a n

d th

e 20

deg

ree s

ea s

t of

the

bou

nda r

y.(c

) Unc

e rta

inty

wa s

ca l

c ula

ted

wit

h th

e st

a nda

rd e

rror

of

[∂C

/∂x

– 0.

5·∂A

/∂x ]

·(∂N

/∂x )

–1,

unc e

rta i

nty

of (

C/N

) NC

P r

a tio

(6.

6 ±

0.4)

, a n

d a n

pre

sum

ed 2

5%un

c ert

a int

y of

Fga

s.(d

) Unc

e rta

inty

wa s

ca l

c ula

ted

wit

h th

e st

a nda

rd e

rror

of

∂C/∂

x ·(∂

A/∂

x )–1

and

an

unc e

rta i

nty

of F

gas/

FN

CP.

Cru

ise

Dat

e(a)

Lon

gitu

de(a

)[∂

C/∂

x −

0.5∂

A/∂

x]⋅(∂

N/∂

x)−1

Fga

s/FN

CP

∂A/∂

x⋅(∂

C/∂

x)−1

FP

IC/F

NC

PF

gas(b

)F

NC

P(c

)F

PIC

(d)

MR

98K

2Ja

n. 7

−15,

199

916

0°E

−170

°W10

.1 ±

0.1

0.52

± 0

.03

0.13

± 0

.02

0.10

± 0

.02

4.0

7.6

± 2.

00.

8 ±

0.3

MR

99K

7D

ec. 1

−9, 1

999

160°

E−1

70°W

9.2

± 0.

20.

40 ±

0.0

30.

21 ±

0.0

20.

17 ±

0.0

29.

925

.0 ±

6.5

4.3

± 1.

2M

R00

K8

Jan.

12−

23, 2

001

160°

W−1

60°W

10.0

± 0

.10.

51 ±

0.0

30.

03 ±

0.0

10.

02 ±

0.0

38.

917

.3 ±

4.5

0.4

± 0.

2

88 M. ISHII et al.

20% of the NCP. Their result is consistent with that evaluated by Zhang and Quay(1997) on the basis of the budget of TCO2 and its 13C/12C in boreal spring and fallof 1992 in the eastern equatorial Pacific. According to these results, we assumethat FTOCacm is 20% of NCP, hence FTOCacm/FPOCsink = 0.2/0.8 = 0.25. Then, takingthe C/N stoichiometric ratio in total organic matter in surface layer (FTOCacm/FTONacm = 12; Hansell and Waterhouse, 1997) and the C/N stoichiometric ratio insinking particulate matter at the bottom of the euphotic layer (FPOCsink/FPONsink =5.9 ± 0.4; Rodier and LeBorgne, 1997), (C/N)NCP was estimated as 6.6 ± 0.4. Theresults of FGAS/FNCP obtained in three R/V Mirai cruises in the western andcentral equatorial Pacific in the boreal winters during the non-El Niño periods1999–2001 are listed in Table 3.

Due to the interannual variation in [∂C/∂x – 0.5∂A/∂x]·(∂N/∂x)–1 ratiobetween 9.2 ± 0.2 and 10.1 ± 0.1, Fgas/FNCP changed between 0.40 to 0.52.Assuming that (C/P)NCP ratio is 106, Fgas/FNCP was also calculated from a similarequation to Eq. (15) as 0.39 to 0.45. These results indicate that the net biologicalCO2 uptake rate in the surface layer is about twice as fast as the net CO2 evasionrate to the atmosphere on average. Mean Fgas in the western region in thedivergence zone within 20 degrees from the boundary to the warm pool that wasobserved during these three cruises ranged from +4.0 to +9.9 mmol m–2day–1, andFNCP thus calculated was 7.6 to 25 mmol m–2day–1. On the other hand, whencalculated with the mean Fgas in the western divergence zone in October–Januaryin the years 1990–2000 (4.2 ± 1.8 mmol m–2day–1; see Fig. 11), the mean FNCP inOctober–January is estimated as 9.3 ± 4.3 mmol m–2day–1.

These estimates are consistent with the FPOCsink out of the euphotic layerevaluated by drifting sediment trap experiments in August–September 1992 at140°W (2 to 30 mmol m–2day–1) and that in October 1994 in the westerndivergence zone at 150°W (FPOCsink = 21.2 ± 2.4 mmol m–2day–1; Rodier and LeBorgne, 1997) with 20% accumulation of organic carbon in the surface layer(Hansell et al., 1997). They are also in fair agreement with the NCP that wasestimated from the budgets of TCO2 and its 13C/13C isotopic ratio (11.5 ± 6.3mmol m–2day–1; Zhang and Quay, 1997) and a TCO2 budget calculated with itszonal transport (13 mmol m–2day–1; Wanninkhof et al., 1995) in August–September 1992 in the divergence zone in the eastern equatorial Pacific, as wellas new production evaluated along 140°W in 1992 (18.5 mmol m–2day–1;McCarthy et al., 1996) and that evaluated for 90°W–180° with an ecosystemmodel embedded within an ocean general circulation model (16 mmol m–2day–1;Chai et al., 1996).

The ratio of sinking particulate carbonate flux to NCP, FPIC/FNCP, in thewestern divergence zone could also be estimated from the property-property plotsof pot.NTA vs. NTCO2 (Fig. 13). On the same assumption as in Eqs. (10) to (13),the ratio of zonal pot.NTA gradient to zonal NTCO2 gradient is expressed as

∂A/∂x (∂C/∂x)–1 = 2·FPIC·{FNCP + Fgas + FPIC}–1

= 2·FPIC

·{FNCP

·(1 + Fgas

/FNCP

) + FPIC

)}–1 (18)

Variability of Surface Layer CO2 Parameters 89

and then the ratio of FPIC/FNCP is given as

FPIC

/FNCP

= (1 + Fgas

/FNCP

)·{2∂C/∂x·(∂A/∂x)–1 – 1}–1. (19)

The ratio of zonal pot.NTA gradient to zonal NTCO2 gradient ∂A/∂x·(∂C/∂x)–1,and the ratio of FPIC/FNCP in the divergence zone in the western and centralequatorial Pacific observed in three R/V Mirai cruises in the boreal winters 1999–2001 are also listed in Table 3. Although the zonal gradient of near-surface NTAis small (see Table 2), the ratio ∂A/∂x(∂C/∂x)–1 was determined to range between0.03 ± 0.01 and 0.21 ± 0.02. Consequently, FPIC/FNCP also changed year by yearfrom 0.02 ± 0.01 to 0.17 ± 0.02. These results indicate that the flux of sinkingparticulate carbonate changed in the range between 2% and 17% of NCP, andbetween 4% to 44% of sea-to-air CO2 flux. Taking the NCP calculated for eachcruise, FPIC is then calculated as ranging from 0.4 ± 0.2 mmol m–2day–1 in January2001 to 4.3 ± 1.2 mmol m–2day–1 in December 1999. If we take the mean NCP of9.3 ± 4.3 mmol m–2day–1 in the western divergence zone for the period 1990–2000, mean FPIC is estimated as 0.9 ± 0.8 mmol m–2day–1.

These estimates are in fair agreement with the calcification rate in the surfacelayer (ca. 3 mmol m–2day–1) in August/September 1992 in the divergence zone inthe eastern equatorial Pacific (Balch and Kilpatrick, 1996) and particulatecarbonate export out of the euphotic layer of 2.3 ± 0.3 mmol m–2day–1 asdetermined by a flowing sediment trap experiment in October 1994 at 150°W inthe western equatorial divergence zone (Rodier and Le Borgne, 1997).

Fig. 13. Relationships between potential NTA and NTCO2 in surface layer of the equatorialdivergence zone to the west of 170°W in Jan. 1999 (�) and in Dec. 1999 (�), and to the west of160°W in Jan. 2001 (+).

2305

2310

2315

2320

2325

2330

2335

1950 1970 1990 2010 2030

po

t. N

TA

/ µ

mo

l k

g-1

NTCO2 / µ mol kg

-1

90 M. ISHII et al.

In conclusion, these results demonstrate that an investigation of the spatialand temporal variability in the oceanic CO2 parameters, including TCO2, pH and/or TA as well as pCO2sw and other oceanographic properties, is important inunderstanding the fluxes of carbon and their variability within the ocean andbetween the ocean and the atmosphere. The result of quantification in this studyshows that the production of organic matters as well as the circulation of wateris playing a major role in the variability of the oceanic CO2 system in the surfacelayer and is thus significantly affecting the sea-to-air CO2 flux in the divergencezone of the equatorial Pacific. Sea-to-air CO2 evasion is also important inchanging the oceanic CO2 system in this zone. Formation of sinking particulatecarbonate is also perturbing, but its effect is relatively smaller than the netproduction of organic carbon and net CO2 evasion to the atmosphere. It is alsostrongly suggested that the net community production and the flux of particulatecarbonate as well as the net sea-to-air CO2 flux are changing year by year in thewestern part of the equatorial divergence zone.

Acknowledgements—This study was supported by a grant from the “Global Carbon Cycleand Related Mapping based on Satellite Imagery Program (GCMAPS)” of the Ministry ofEducation, Culture, Sports, Science and Technology (MEXT). We gratefully acknowledgethe officers and crew of the R/V Kaiyo and R/V Mirai of Japan Marine Science andTechnology Center (JAMSTEC), and skillful technicians from Marine Work Japan Ltd.,Nippon Marine Enterprise Ltd., Global Ocean Development Inc., and Kansai EnvironmentalEngineering Center Co. for their skillful assistance on board and providing the hydrographicand meteorological data sets. Data obtained on board R/V Mirai are to be available fromData Management Office of JAMSTEC (http://www.jamstec.go.jp/mirai/).

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