Vertical fluxes of nutrients and carbon through the halocline in the western subarctic Gyre calculated by mass balance

Deep-Sea Research II 49 (2002) 5577–5593

Vertical fluxes of nutrients and carbon through the halocline in the western subarctic Gyre calculated by mass balance A. Andreev*, M. Kusakabe, M. Honda, A. Murata, C. Saito Japan Marine Science and Technology Center, 2-15 Natsushima-cho, Yokosuka, Kanagawa 237-0061, Japan Received 2 February 2001; received in revised form 18 February 2001; accepted 19 February 2001

Abstract Using data collected in the R/V Mirai (Japan Marine Science and Technology Center) cruises (November–December 1997, November–December 1998, May 1999, January–February 2000, and May–June 2000) we have computed the vertical fluxes of nutrients and carbon through the halocline, located at the base of the surface layer (B100 m) in the Western Subarctic Gyre (WSG). The vertical fluxes were estimated by mass balance using a one-dimensional model and taking the fresh-water flux to the surface of the WSG to be 0.40 m yr
1 at two sites: in the central part of the WSG (501N, 165–1701E) and close to its southwestern edge (44.0–44.51N, 1551E). Based on our calculations, the annually averaged vertical export fluxes of the silica, carbonate, total organic nitrogen, phosphorus, and carbon at the study sites were 1.070.2, 0.3470.06, 0.4070.10, 0.02470.006, and 3.270.6 mol m
2, respectively. Comparison with sedimenttrap data collected at 1000 and 3000 m at the study sites between 1997 and 2000 (Fluxes and chemical compositions of particulate matter in the northwestern North Pacific: results from sediment-trap experiments (1997–2000), this volume) show that B93% of particulate organic carbon and nitrogen exported from the surface layer of the WSG had undergone remineralization from 100 to 1000 m. The computed annual uptake of excess CO2 and the change in d13C due to the Suess effect in the study area were 0.7470.1 mol m
2 and
7.471.0 per mil m, respectively. The depth-integrated d13C change rate due to the Suess effect in the WSG was slightly lower than a global (Pacific, Indian, and Atlantic oceans) depth-integrated d13C change rate (–9.772.4 per mil m yr
1 (Global Biochemical Cycles 13 (1999) 857)). r 2002 Elsevier Science Ltd. All rights reserved.

1. Introduction The accurate determination of the fluxes of elements (mainly carbon) from the surface layer to the ocean interior related to the biological pump is very important for understanding the global carbon flux and its response to climate change. The northwest Pacific has been considered an *Corresponding author. Fax: +81-468-653-202. E-mail address: [email protected] (A. Andreev).

important region for ocean biochemical and climate change studies (Honjo, 1997). Also, it is important to monitor the regional uptake of excess CO2, defined as an increase in dissolved inorganic carbon (DIC) in the seawater due to absorption of anthropogenic carbon dioxide from the atmosphere (Brewer, 1978; Chen and Millero, 1979) for detection of change in the rate of carbon accumulation in the oceans (Wallace, 2001). In the North Pacific, the intensive, long-term surveys of the vertical fluxes of nutrients and

0967-0645/02/$ - see front matter r 2002 Elsevier Science Ltd. All rights reserved. PII: S 0 9 6 7 - 0 6 4 5 ( 0 2 ) 0 0 2 0 0 - X

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carbon have been conducted in the eastern and central zones (Wong et al., 1999; Karl et al., 1996; Emerson et al., 1997; Christian et al., 1997), but there is lack of such a systematic study in the western subarctic Pacific (Harrison et al., 1999; Honda et al., 2002). The circulation of the Western Subarctic Gyre (WSG) (Fig. 1) is strongly cyclonic and elongated along the Kamchatka Peninsula and Kuril Islands, and is formed by the East Kamchatka Current, Oyashio Current, Subarctic Current, and Alaskan Stream (Dodimead et al., 1963; Ohtani, 1970; Favorite et al., 1976). A principal feature of the WSG is the sharp halocline/pycnocline at the bottom of the upper layer at about 100 m depth (Figs. 2A and B). A distinct mesothermal layer (warm subarctic intermediate layer) is associated with this halocline (Fig. 2C). The halocline prohibits the deepening of the upper layer due to winter cooling because the surface water is too fresh (Reid, 1973). Climatological data show that low surface salinity in the northern North Pacific is due to low evaporation rate and a small rate of flow into and out of the near-surface layer (Warren, 1983). In the subarctic Pacific, the strong halocline/pycnocline limits diapycnal vertical exchange between the surface and deeper layers. The biological pump controls the vertical gradients of

the biogenic elements in the halocline by means of the vertical detritus flux (Wong et al., 1999). The halocline coincides with strong vertical gradients in DIC, nitrate, and silicate (Figs. 2D–F). Primary productivity measurements in the western subarctic Pacific have been conducted mainly in the coastal area of the Oyashio current and do not cover all seasons (Shiomoto, 2000). In the WSG and oceanic part of the Oyashio Current the integrated (from the surface to the 0.2% light depth) primary production in the spring of 1993, 1994, and 1995 ranged from 114 to 2046 mg C m
2 d
1 (Shiomoto, 2000). The magnitude of primary production in the western subarctic Pacific during spring bloom was comparable to that in the eastern subarctic Pacific (Harrison et al., 1999; Shiomoto, 2000). Sediment-trap experiments conducted in the northwestern Pacific (Noriki et al., 1999; Honda et al., 2000) provide information about the vertical flux of silica, carbonate, and particulate organic carbon (POC) in the deep layer (1000–5000 m) at two sites (Stn. KNOT (441N, 1551E) and 50N165E (501N, 1651E). The establishment of a time series station KNOT (441N, 1551E), located at the edge of the WSG was an important step towards understanding the seasonality of chemical parameters, primary productivity and flux in the NW Pacific. However, due to episodic

Fig. 1. Stations locations.

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Fig. 2. Vertical profiles of density anomaly (st =density—1000 kg m
4) (A), salinity (B), temperature (C), silicate (D), nitrate (E), and DIC (F) at the Stns. 50N, 165/170E (&—&) and KNOT/KNOT-N (m—m).

intrusions of the subtropical transformed water (warmer and saltier than WSG water) and Oyashio water (fresher and colder than WSG water) Stn. KNOT could not itself represent the biogeochemical features of the WSG. Based on data collected from 1997 to 2000 (R/V Mirai), we calculated the vertical fluxes of nutrients (silicate, nitrate, and phosphate) and DIC (including the excess CO2 flux) at the boundary (halocline) between the surface and intermediate layers in the WSG(B100 m). The vertical fluxes of nutrients and DIC were computed using a one-dimensional model in the central part of the WSG and at its south-western edge (Fig. 1). The sediment-trap experiment conducted at the study sites (Stns. 50N165E and KNOT) at depths of 1000 and 3000 m from 1997 to 2000 (Honda et al., 2002) allowed comparison of our computed upward fluxes of silicate, DIC, and nitrate, and the downward fluxes of silica, carbonate, POC, and particulate organic nitrogen (PON) obtained and referenced to 100 m.

Although our approach is limited in not considering lateral advection and diffusion, it provides information on the average nutrient and carbon fluxes in the WSG. In this paper, we first consider the difference in nutrients and DIC between the surface and intermediate layers due to the export fluxes of the opal, carbonate, organic carbon, organic nitrogen, and organic phosphorus from the euphotic layer. Second, we discuss the comparison of our computed fluxes of materials through the halocline with field data from the sediment-trap experiment. Finally, we analyze the coefficient of vertical diffusion, and change in vertical fluxes of nutrients in the study area.

2. Materials and methods 2.1. Sampling and analysis The data at four stations in the WSG (Fig. 1) were collected during five cruises of the R/V Mirai:

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MR97-02 (November–December 1997), MR98K02 (November–December 1998), MR99-K03 (April–May 1999), MR00-K01 (January–February 2000) and MR00-K03 (May–June 2000). Seawater sampling was supplied by a carousel Rosette (SBE32; Sea-Bird Electronics, Inc.) equipped with 30-l Niskin bottles and a SBE911 CTD. The salinity was measured with a Guildline Salinometer (Model 8400B). Dissolved oxygen (DO) was analyzed by potentiometric titration (Dickson, 1994). SiO4, PO4, NO3 (NO3 and NO2) in the water samples were determined using a Bran– Luebbe continuous flow analyzer (Gordon et al., 1992) with an accuracy (1s) of less than 1% for SiO4, and 1–2% for PO4 and NO3. DIC was measured by a coulometer (carbon dioxide coulometer Model 51012, UIC Inc.). Calibration of the coulometer was carried out using sodium carbonate solutions and Dickson seawater standard (DOE, 1994). The repeatability (1s) of DIC measurements was 7 3 mmol kg
1. Total alkalinity (TA) was determined by potentiometric titration in an open cell (Dickson et al., 2002). The repeatability (1s) of the TA measurements was 73 mmol kg
1. In our calculations we used isotopes of carbon (13C and 14C) data from the MR97-02 cruise. The sampling procedure and methods applied are described in (Kumamoto et al., 2002). The accuracy (1s) of the d13C and d14C measurements were 70.1 and 75 per mil, respectively. In this study sediment-trap experiments conducted in the WSG between December 1997 and May 1999 were used (Honda et al., 2002). Timeseries sediment traps with 21 collecting cups (Mark 7G-21) were installed at B1000, 3000, and 5000 m at Stns. KNOT and 50N 165E. The sediment-trap sample collection, treatment procedures, and methods are described in Honda et al. (2002). 2.2. Vertical fluxes calculation To calculate the fluxes of the dissolved material such as nutrients and DIC (C) between the surface and the intermediate layer, a one-dimensional (vertical) conservative equation was used. At steady state (constant annual average of nutrients and salt contents in the surface water), the annual

average fluxes (F ) of dissolved elements (nutrients and DIC) from the deep to surface layer should compensate for the export of silica, carbonate, organic phosphorus, organic nitrogen, and organic carbon fluxes from the surface layer Z  F ¼
w qC=qz dz þ Kz qC=qzz¼h ð1Þ and the flux of fresh water (P2E) should balance the salt flux Z ðP
EÞrS ¼
w qðrSÞ=qz dz þ Kz qðSrÞ=qzjz¼h ;

ð2Þ

where w and Kz are vertical velocity and the coefficient of vertical turbulent diffusion, respectively, r is density, S is salinity, and h is the depth of the surface layer. Excess CO2 uptake ðFCO2 ex Þ equals the increase in excess CO2 in the surface layer and the flux of excess CO2 from the surface to the deep layer: Z ðFCO2 ex Þ ¼ qCO2 ex =qt h
w qCO2 ex =qz dz þ Kz qCO2 ex =qzjz¼h :

ð3Þ

The overbar indicates the mean value over the upper 100 m, which approximates the depth of the winter mixed layer, h:

3. Results 3.1. The relation between the nutrients, DIC, DO, potential alkalinity, excess CO2, and the salinity Distributions of DIC, nutrients, and salinity in the surface layer of the WSG show seasonal variability. In early spring, the distributions of DIC, nutrients, and salinity were quite uniform. Average concentrations of silicate, nitrate, and salinity were 45–50, B25 mmol kg
1 and 33.0–33.1, respectively (Fig. 2). The significant decrease in nutrients and DIC concentrations took place during the spring bloom (April–June). In 1997– 1999 in the surface layer at Stns. KNOT and 50N165E, the reduction in the phosphate, nitrate, silicate, and DIC inventories during the bloom were 0.025–0.040, 0.4–0.6, 0.9–1.2,

A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

and 2.5–3.5 mol m
2, respectively. From spring to fall, the salinity decreased mainly because precipitation exceeded evaporation (the difference is >2 mm d
1; NOAA-CIRES Climate Diagnostics Center). In November–December, due to cooling and strong mixing, the seasonal pycnocline was destroyed, and distributions of nutrients and salinity in the upper water layer were uniform. In late winter and early spring the concentration of nutrients and salinity in the surface water increased, due to decreases in the export fluxes of elements from the euphotic layer (Honda et al., 2002), the difference between precipitation and evaporation, and the increase in nutrients and vertical salinity fluxes through the halocline. In the halocline of the WSG (depth E100– 150 m, salinity E33.2–33.6), the concentrations of nutrients, DIC and excess CO2, were linearly correlated with salinity (Figs. 3–5 (A, E)). An analysis of historical data (1953–1974, data of JODC) (Figs. 3D–F), and data from the MR9702–MR00-K03 cruises (Figs. 3A–C) did not show

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any significant seasonal changes in slopes between the silicate, nitrate, phosphate, and the salinity unlike their surface variability. Using the linear relationships between the nutrients, DIC, excess CO2, and salinity, we can eliminate the turbulent diffusion and advection terms from the right-hand side of the conservative equations (Eqs. (1)–(3)). Therefore, the quantification of the excess CO2 transported downward from the surface layer can be derived from the upward flux of silicate, nitrate, phosphate, and DIC to the surface, or from the annual average of fresh-water flux (P2E): ðFCO2 ex
qCO2 ex =qt hÞ=k1 ¼ FSiO4 =k2 ¼ FDIN =k3 ¼ FDIP =k4 ¼ FDIC =k5 ¼ ðP
EÞS;

ð4Þ

where k1 ; k2 ; k3 ; k4 ; and k5 are slopes in the linear relationships between the excess CO2 and salinity, the silicate and salinity, the DIN and salinity, the phosphate and salinity, and the DIC and salinity, respectively.

Fig. 3. (A) Silicate, (B) nitrate, and (C) phosphate versus salinity at the Stns. 50N,165/170E (&—&) and KNOT/KNOT-N (m—m) (MR97-02-MR00-K04); (D) silicate, (E) nitrate, and (F) phosphate versus salinity in the WSG (1953-1974, JODC data).

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Fig. 4. (A) Silicate, (B) nitrate, and (C) phosphate, normalized to salinity 35, versus temperature from 50 to 100 m of the WSG and the subarctic frontal zone (MR97-02-MR00-K03); (D) silicate, (E) nitrate, and (F) phosphate, salinity-normalized and temperature corrected by Eqs. (5a)–(5c), versus salinity at the Stns. 50N, 165/170E (&—&) and KNOT/KNOT-N (m—m) (MR97-02-MR00K03).

Using data from the MR97-02–MR00-K03 cruises, we computed the slopes of the correlations between silicate, nitrate, phosphate, DIC, DO, excess CO2, and salinity at Stns. KNOT/KNOT-N and 50N-165E/50N-170E (Table 1). Excess CO2 was calculated from DIC, TA and DO data (Appendix A). The excess precipitation over evaporation in the WSG decreases the concentration of the nutrients and DIC in the surface-mixed layer. To take this into account the relations between nutrients and DIC normalized to salinity 35 (NSiO4, NNO3, NPO4 and NDIC) and salinity were considered (Table 1). In our calculations we assumed the effects of lateral mixing and advection to be negligible. However, to keep salinity at the stationary state, the fresh-water flux to the WSG should be compensated by the lateral flux of salt to the WSG from the subtropical area. The study layer (100–150 m) is located between a colder, fresher surface layer and a warmer, saltier intermediate

layer (Figs. 2B and C), caused by the lateral mixing between the WSG and Subarctic Frontal Zone waters. The difference in the nutrients, DIC and DO between the surface and intermediate layer in the WSG also can be due to differences in the preformed values of silicate, phosphate, nitrate, and DIC between these layers. The lateral flux of the subtropical water to the WSG varied on a decadal scale (Andreev and Kusakabe, 2001). In 1997–2000 (MR97-02–MR00-K03 cruises), the supply of subtropical water to the WSG was not significant (Andreev and Kusakabe, 2001). However, to find the difference in the nutrients and DIC between the surface and intermediate layers of the WSG related to the biological pump, we should take into account the difference in the preformed values between these layers. To quantify the difference in DO and DIC through the halocline in the WSG related to the biological activity (Table 1), we used apparent oxygen utilization (AOU) and NDICt (NDIC+11t (1C)). The temperature coefficient of NDICt

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Fig. 5. (A) DIC, (B) potential alkalinity (PTA), (C) excess CO2, (D) DIC, salinity normalized and temperature corrected (NDIC+11  t (1C)), and (E) PTAt=PTA+4t (1C) versus salinity at the Stns. 50N, 165/170E (&—&) and KNOT/KNOT-N (m—m) (MR97-02MR00-K03); (F) B—d13C (per mil) and E—(d14C/50+1.2) (per mil) versus salinity at the Stns. 50N, 165/170E and KNOT (MR97– 02). Table 1 Differences in the nutrients, DIC, DO, AOU, PTA, excess CO2, d13C, d14C between the surface and the intermediate layer of the WSG (B100–150 m) normalized by difference in salinity DSiO4 DS 7073

DNO3 DS 2572

DPO4 DS 1.570.1

DDIC DS 22075

DNSiO4 DS 6673

DNNO3 DS 2472

DNPO4 DS 1.470.1

DNDIC DS 17575

DNSiO4t a DS 7573

DNNO3t a DS 3072

DNPO4t a DS 1.870.1

DNDICt a DS 21275

DDO DS 36575

DAOU DS 32575

DCO2 ex DS
5075 DPTA DS 4075

Dd13 C DS
1.870.2

DPTAt a DS 5075

Dd14 C DS
100710

a NSiO4, NNO3 and NPO4 are calculated by Eqs. (5a)-(c); NDICt=NDIC+11t (1C) and PTAt=PTA+4t (1C). Dd13 C=DS and Dd14 C=DS in per mill and all others in mmol kg
1.

(11 mmol kg
1 1C
1) is close to that applied in the calculation of the excess CO2 (Appendix A). This coefficient shows the decrease in DIC with the increase of temperature in the seawater, which is in CO2 equilibrium with the atmosphere. It is well known that nutrients and temperature are highly correlated in the surface layer of the

ocean (Kamykowski and Zentara, 1986). In the surface layer of the northwestern Pacific there is also a tendency for salinity-normalized silicate, nitrate, and phosphate to decrease with increasing temperature (Chen et al., 1986) (Figs. 4D–F). To find the difference in nutrients between the surface and intermediate layer of the WSG related to the

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biological pump, NSiO4t NNO3t ; and NPO4t computed by Eqs. (5a)–(5c) were used NSiO4t ¼ NSiO4 þ 2:0t ð1CÞ;

ð5aÞ

NNO3t ¼ NNO3 þ 1:7t ð1CÞ;

ð5bÞ

NPO4t ¼ NPO4 þ 0:1t ð1CÞ:

ð5cÞ

The absolute values of the temperature coefficients of Eqs. (5a)–(5c) were derived by the linear fit of the data shown in Fig. 4. The correction of the nutrient concentrations by Eqs. (5a)–(5c) increased the slopes between the salinity-normalized silicate, nitrate, phosphate and salinity by 20– 30% (Table 1). To quantify the export flux of the carbonate from the euphotic layer the difference in the TA between surface and intermediate layers was considered. To eliminate the change in TA due to protonation and deprotonation during organic matter synthesis and oxidation (Brewer et al., 1975), and due to the difference between precipitation and evaporation, the TA together with nitrate, normalized to salinity 35 (PTA), was used (Fig. 5B, Table 1). The salinity-normalized TA decreased with temperature in the surface North Pacific at B3 mmol kg
1 1C
1 (Chen et al., 1986). In the surface WSG and Subarctic Frontal Zone the PTA decreased with temperature at B4 mmol kg
1 1C
1. The temperature correction of PTA (PTAt=PTA+4t (1C)) increased the slope between PTA and salinity from 40 to 50 mmol kg
1 (Table 1). In order to assure the validity of our approach, we examined the stoichiometric ratios between P, N, C and DO related to the formation/decomposition of particulate organic matter (P:N:Corg: O2) in the WSG. For comparison, we used stoichiometric ratios determined by Anderson and Sarmiento (1994). They found that for the global ocean these ratios are P:N:Corg: O2=1:1671:117714:
170710 and approximately independent of depth and ocean basin. Using NPO4t ; NNO3t ; NDICt ; and AOU (Table 1) we found that stoichiometric ratios in the WSG were P:N:Corg:O2B0.95:16:125:
171. The difference in the DIC between the surface and intermediate layer of the WSG due to particulate

organic matter flux (DCorg) was computed as (DNDICt+DCO2 ex–0.5DPTAt(DDICcarb.dis) (see Appendix A)) using data of Table 1. The calculated stoichiometric ratios are in good agreement with those of Anderson and Sarmiento (1994). 3.2. Isotopes of carbon To provide additional information about the fluxes of carbon in the study area we did consider the difference in isotopes of DIC (d13C and d14C) between the surface and the intermediate layers. d13C in seawater is affected by the anthropogenic CO2 supply from the atmosphere (d13Cantr) (Suess effect) and the biological activity (d13Cbiol) (Broecker and Peng, 1982). Changes in 14C came from nuclear weapons testing (Broecker et al., 1985). Using the rate of change of the depth-integrated d13Cantr, in ocean basins, the global oceanic uptake of excess CO2 uptake is estimated (Quay et al., 1992). Due to the Suess effect the d13C value of the DIC in the surface waters of world ocean has been decreasing at a rate of B0.02 per mil yr
1 (Quay et al., 1992). The largest decreases in both mixedlayer and depth-integrated d13C values occurred in the subtropical gyres; smaller decreases occurred in the equatorial and subpolar oceans (Quay et al., 1992; Gruber et al., 1999). These d13C trends were consistent with the increase of d14C in the mixed layer, because both the 13C and 14C were brought into the ocean by air-sea CO2 exchange and affected by the same physical processes (advection and water mixing) (Quay et al., 1992). Quay et al. (1992, Fig. 4) showed that the ratio of d14C increase to d13C decrease in the surface North Pacific from 1970 (1973) to 1991 was B
200. In our calculations we assumed that in the surface and upper intermediate layers of the WSG, the ratio between the decrease in d13C due to the Suess effect and the increase in d14C is
200 (see Appendix B). Thus, in the 100–150 m of the study region the slope between d13Cantr and salinity (Dd13 Cantr =DS) was 0.50 per mil and the slope between d13Cbiol and salinity (Dd13 Cbiol =DS) was
2.30 per mil. Dd13 Cantr =DS and Dd13 Cbiol =DS represent, respectively, the difference in

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Table 2 Calculated vertical fluxes between the surface and the intermediate layer, and the increase in excess CO2 and decrease in d13Cant in the WSG FSiO4 ; FSiO2 (mol m
2 yr
1)

FNO3 ; FTON (mol m
2 yr
1)

FPO4 ; FTOP (mol m
2 yr
1)

FTOC a (mol m
2 yr
1)

0.5  FPTA b, Fcarb (mol m
2 yr
1)

FCO2 ex (mol m
2 yr
1)

d13Cant (per mil m yr
1)

1.070.2

0.4070.10

0.02470.006

3.270.6

0.3470.06

0.7470.1


7.471.0

a b

FTOC is computed by Eq. (6). The coefficient 0.5 shows that the flux of carbonate is half of those for the potential alkalinity.

salinity-normalized d13C due to the Suess effect and biological activity between the surface (100 m) and the upper intermediate layer (150 m). The ratio between the increase in excess CO2 and the decrease in d13Cantr in the seawater varied from B
50 (subtropics) to B
100 mmol kg
1 mil
1 (subarctic, subantarctic and equator) (Quay, 2000). Using DCO2 ex =DS (
50 mmol kg
1) and Dd13 Cantr =DS (0.5 %) (Table 1) we obtained the same ratio (
100 mmol kg
1 per mil) for the study area (subarctic region). 3.3. Fluxes of nutrients and carbon To determine the fluxes of nutrients and carbon (including excess CO2 flux) between the surface and intermediate layers by Eq. (4), we used the annual average of excess precipitation over evaporation (1.1 mm d
1 or 0.40 m yr
1) (data of the NOAA-CIRES Climate Diagnostics Center) in the WSG region. The uncertainty in the (P
E) can be as high as 30–40% with additional interannual variations. The accepted fresh-water flux is very close to the 0.36 m yr
1 (as excess precipitation over evaporation) of Warren (1983) in the northern North Pacific (451–601N). The additional source of the fresh water for the surface of the northern North Pacific related to river run-off is B60% of the excess precipitation over evaporation (Warren, 1983). The input of the river run-off to the water balance is more important for the eastern subarctic Pacific than for the WSG. The mixed surface layer in the Alaska Gyre was fresher (SB32:5232:6) than that for the WSG (SB33:0233:1), although there was no difference in annually averaged excess precipitation over evaporation between these regions (data of the NOAA-CIRES Climate Diagnostics Center). The

over/under estimation of (P2E) could lead to the over/under estimation of the absolute values of fluxes of elements computed by us (Table 2), but it does not change the ratios between the fluxes. In the mass balance of DIC in the surface layer, the flux of the carbon dioxide between atmosphere and ocean ðFCO2 atm2ocean Þ should be taken into account. The distribution of FCO2 atm2ocean in the world oceans shown by Takahashi et al. (1999) and data of Nojiri et al. (1999) suggest that the WSG was a weak sink for the atmospheric CO2. The net sea-air CO2 flux was B–171 mol m
2 yr
1 (Takahashi et al., 1999). Within this net flux of CO2 the flux of excess CO2 was B–0.7 mol m
2 yr
1 (Table 2). Thus the residual sea-air CO2 flux was B–0.371 mol m
2 yr
1, several times smaller than its uncertainty, hence negligible in our balance calculations. We assumed that in the study area FCO2 atm2ocean was equal to FCO2 ex : The export flux of the total (in particulate and dissolved forms) organic carbon (FTOC ) (Table 2) was determined as FTOC ¼ FDIC þ FCO2 ex
Fcarb ;

ð6Þ

where FDIC calculated by Eq. (4) using the slope between NDICt and salinity for k5 (Table 1).

4. Discussion 4.1. The comparison with the Sediment-trap data To compare the calculated export fluxes of the organic carbon and nitrogen through the halocline (Table 2) with the sediment-trap data, we have to separate the fluxes of POC and PON from the fluxes of dissolved organic carbon (DOC) and dissolved organic nitrogen (DON). In the eastern

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Pacific (in the area north of the Subarctic front) the concentration of DOC in the surface layer was B60 mmol l
1 (Abell et al., 2000). The difference in DOC and DON through the main halocline were B15 and 2 mmol l
1, respectively (Abell et al., 2000). At Stn. KNOT in 1998/99 the concentration of DOC in the surface layer and the difference in DOC through the halocline were B60 and 15 mmol l
1, respectively (Yamaguchi et al., 1999, 2000). The DIC and DIN gradients across the halocline were 10 times higher than those for the DOC and DON and therefore, the fluxes of DOC and DON from the surface to intermediate layer were 10% of those for, respectively, DIC and dissolved inorganic nitrogen. Based on mass balance, the fluxes of organic carbon and organic nitrogen in particulate form were estimated to be 90% of total organic carbon and nitrogen export fluxes (Table 2). Thus, in the halocline of the WSG FPOC and FPON were B2.9 and 0.36 mol m
2 yr
1. Based on sediment-trap data (Honda et al., 2002), the carbonate flux at Stns. KNOT and 50N, 165E between December 1997 and December 1999 was 0.14 mol m
2 yr–1 at 1000 m and 0.14 mol m
2 yr–1 at 3000 m, with the average value of 0.1470.02 (1s) mol m
2 yr–1. The silica flux was 0.39 mol m
2 yr–1 at 1000 m and 0.44 mol m
2 yr–1 at 3000 m with the average value of 0.4270.08 (1s) mol m
2 yr–1. At 1000/3000 m of the WSG in 1997/ 99 the annual average silica to carbonate ratio (mole) ðFSiO2 =FCaCO3 Þ was 3.070.4 (1s) and the ratio (mole) of carbon to nitrogen in the particulate organic matter (FCorg =FNorg ) was 8.070.4 (1s). The ratio between FSiO2 and FCaCO3 and between FCorg and FNorg determined by sediment traps in the deep layer is the same as our estimates (FSiO2 =FCaCO3 B3 and FCorg =FNorg B7:7; respectively) in the halocline (B100 m) of the WSG. A similar value for the FCorg =FNorg ratio (B8) in the particulate organic matter was obtained in the eastern subarctic Pacific (Wong et al., 1999) using sediment-trap data. To compare the calculated carbonate export flux with the sediment-trap data, the carbonate flux obtained from sediment traps at 1000 and 3000 m was referenced to 100 m using the equations of Martin et al. (1993) and Yamanaka and Tajika (1996), hereafter (M) and (YT). As suggested by

Honda et al. (1997), these equations can be used as upper and lower limits of the changes in the vertical flux of carbonate with depth in the northwestern Pacific. Thus at 100 m of the WSG the carbonate flux was 0.19–0.32 (YT) and 0.42–0.70 mol m
2 yr
1 (M) with the average value of 0.40 mol m
2 yr
1. The average value of the carbonate flux obtained by sediment traps and referenced to 100 m agrees well with our estimate of 0.34 mol m
2 yr
1 (Table 2). At the upper boundary of the intermediate layer of the WSG (B100 m) the computed FPOC and FPON were, respectively, B2.9 mol C m
2 yr
1 and 0.36 mol N m
2 yr
1. At the lower boundary of the intermediate layer (B1000 m) of the WSG the FPOC and FPON obtained by sediment trap were, respectively, B0.2 mol C m
2 yr
1 and 0.025 mol N m
2 yr
1 (Honda et al., 2002). Therefore, B93% of POC and PON exported from the surface to the deep layers in the WSG remineralized in the intermediate layer. The value of the computed export carbon flux at B100 m of the WSG supports the suggestion of Wong et al. (1999). Based on primary production of 140 g C m
2 and a mean annual f -ratio of about 0.25, obtained in the eastern subarctic Pacific, they have suggested that the export flux of POC from the surface to the intermediate water in the subpolar oceans could potentially yield 35 g C m
2 (B3 mol m
2 yr
1). For comparison in the central North Pacific (Stn. Aloha) the export flux of the TOC from the surface layer estimated by Emerson et al. (1997) was B2 mol m
2 yr
1. The important characteristic of the study area is the rain ratio (the ratio of organic carbon to inorganic carbon export flux). This ratio shows the ability of the different regions of the world ocean to decrease the partial pressure of the carbon dioxide in the surface layer by means of the biological pump to enhance the absorption of atmospheric CO2 by the surface waters. In the halocline of WSG (B100 m) the rain ratio was E9 from 3.2 mol m
2 yr
1 (Forgcarb )/0.34 mol m
2 yr
1 ðFCaCO3 Þ: It was higher than the averaged rain ratio of 4 of the world’s oceans (Broecker and Peng, 1982; Sarmiento and Toggweiler, 1984). The calculated rain ratio shows that the increase in biological activity, and thus export fluxes of the

A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

particulate materials from the surface to the deep layers of the WSG, enhances atmospheric CO2 sink. 4.2. The seasonal change in the nutrients and DIC in the surface WSG Because biological processes in the ocean are largely seasonal, particularly in subpolar regions, the spring–summer nutrient drawdown from the surface waters by biological activity provides an unambiguous lower bound estimate of export production (Minas and Codispoti, 1993; Louanchi and Najjar, 2000). The carbon production estimated by the change in the nitrate, phosphate, and DO concentrations in the surface WSG during spring bloom (from April to June–July) varied from B10 mol C m
2 in the coastal zone (near Kamchatka Peninsula) to 3–4 mol C m
2 in its pelagic zone (the central part of the WSG) (Smetanin, 1959). Accordingly, the silicate content of the surface layer decreased by B1.5 mol Si m
2 in the coastal zone and by B 0.7 mol Si m
2 in the pelagic zone of the WSG during the bloom (Smetanin, 1959). Louanchi and Najjar (2000) have estimated new production (NP) in the ocean by vertically integrating the seasonal nutrient between 0 and 100 m from winter to summer using the global monthly climatology of nutrients (Conkight, 1998). Their value of NP for the WSG (B4 mol m
2) is in agreement with the NP estimated by Smetanin (1959). In 1997–1999, at Stns. 50N, 165E and KNOT, the decrease in nutrients and DIC in the surface layer from the spring to the fall were 0.025– 0.040 mol m
2 for phosphate, 0.4–0.6 mol m
2 for DIN (nitrate and nitrite), 0.9–1.2 mol m
2 for silicate and 2.5–3.5 mol m
2 for DIC, respectively. These seasonal changes were similar to the annual average vertical fluxes determined in this study (Table 2). 4.3. The excess CO2 uptake The annual average of increase in the DIC in the surface North Pacific was 0.770.2 mmol m
3 yr
1 (Winn et al., 1998). Taking qCO02 ex=qt to be

5587

0.7 mmol m
3 yr
1 we determined by Eq. (7) that the uptake of excess CO2 in the WSG ðFCO2 ex Þ was 0.7470.2 mol C m
2 yr
1. The annual average increase in the DIC content from 1973 to 1991–1993 was 0.6370.15 mol C m
2 yr
1 in the northeastern Pacific (Slansky et al., 1997) and 0.770.2 mol C m
2 yr
1 in the western subarctic Pacific (Ono et al, 1999). According to our calculation (Andreev et al., 2001) the net transport of excess CO2 from the Okhotsk Sea to the western subarctic Pacific was almost in balance with the flux of the excess CO2 from the western subarctic Pacific to the subtropical area. According to our simple calculations, we concluded that the excess CO2 exchange between the surface and deep layers was an important factor in the excess CO2 accumulation in the subarctic Pacific (Andreev et al., 2001). The agreement between the computed excess CO2 uptake in the WSG (Table 2) and the increase in the excess CO2 in the subarctic Pacific (Slansky et al., 1997; Ono et al., 1999) supports this conclusion. Based on CO2 ex/d13Cantr ratio (B–100 mmol m
3 mil
1) the depth-integrated change rate in the d13Cantr at the study sites was 7.471.0% m yr
1 (Table 2). Over the time period 1970–1990, the rate of change of the global (Pacific, Indian, and Atlantic oceans) depthintegrated d13C was estimated to 9.772.4% yr
1 (Sonnerup et al., 1999), with higher absolute values in the subtropical regions and lower values in the equator and high-latitude areas. The computed depth-integrated d13C change rate in the WSG was slightly lower than the global mean. 4.4. Vertical eddy diffusivity To quantify the coefficient of the vertical diffusion in the study area, it was assumed that the flux of the dissolved material through the halocline (Eq. (2)) was determined mainly by the vertical turbulent diffusion  F ¼
Kz qC=qz : ð7Þ z
h

In general, the coefficient of the vertical diffusion (Kz ) depends both on the external source of the turbulence and on the dynamical stability of the flow. The location of Stns. KNOT and

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A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

50N165E is far away from the strong currents like West Kamchatka Current and Oyashio Current. In the ocean, away from the sources and sinks of turbulence, and for negligible shear, Kz depends inversely on the vertical stability (N2Eg(qr/qz)/r) (Welander, 1968; Sarmiento et al., 1976). As can be seen from the sediment-trap data (Broecker and Peng, 1982; Wong et al., 1999; Honda et al., 2002) there is no marked decrease in the amount of opal caught as function of depth. Assuming that the dissolution of silica takes place mainly in the deep and bottom layers (Broecker and Peng, 1982), and just below the surface layer the vertical flux of the silicate does not change with depth (FSiO4 is constant) we could get the analytical expression for the dependence between Kz and the vertical stability. In the 100–500 m the relation between qðCSiO4 Þ=qz and N
2 was linear (Fig. 6A), thus Kz ¼ aN
2 :

ð8Þ

Taking the annual flux of the silicate in the WSG to be 1.070.2 mol m
2, we get that a (Eq. (8)) was 2.770.6  10
9 m2 s
3. In the halocline of the WSG the N 2 was B7.5  10
5 s
2 and thus the annual average Kz was 3.571  10
5 m2 s
1. The computed Kz is about two times higher than the coefficient of the vertical diffusion estimated by Matear and Wong (1997) (Kz B1:5  10
5 m2 s
1) in the northeastern Pacific (using one-dimensional advection–diffusion model to analyze the vertical distribution of the chlorofluorcarbons) at the base of the mixed layer (N 2 B12  10
5 s
2). 4.5. The climate change in the flux of nutrients Model simulations have shown that anthropogenic climate change should lead to the increase of rainfall and thus result in the freshening of the surface layer, and the increase in stratification in high latitudes (Sarmiento et al., 1998). Increased stratification and enhanced river run-off may affect the concentration of silicate and phosphate in the surface waters of high latitudes. The main uncertainty is how the change in nutrient concentrations will influence the biological productivity in the surface layer. In the model presented by

Fig. 6. (A)Vertical gradient of silicate (q½SIO4 =qz) versus vertical stability (N2) in the 100–500 m of the WSG (MR9702-MR00-K03); (B) the change in DIC due to the carbonate dissolution calculated by total alkalinity (0.5DPTAt) versus the change in DIC due to the carbonate dissolution calculated by DIC and d13C (DNDICt+100d13C ) from 100 to 1500 m of the WSG (Stns. KNOT and 50N, 170E; MR97-02).

Sarmiento et al. (1998) the influence of change in nutrient concentration on biological productivity was not considered due to its high unpredictability. Matear and Hirst (1999) in their model simulations of the future oceanic CO2 uptake assumed that the decrease in the nutrients in the surface layer should decrease the biological activity and thus decrease the export flux of carbon. The decrease in the salinity and the nutrient concentration in the surface layer over the past 50 years have been reported in the eastern subarctic Pacific (Freeland et al., 1997). Such changes in the stratification and the concentration of nutrients also could take place in the western subarctic

A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

Pacific. Where do the changes in the surface nutrient concentration influence the export fluxes from the surface to the deep layer? To detect changes in the nutrient vertical fluxes between surface and intermediate layer of the WSG, we compared data at 5 stations in the WSG in 1953 with the data collected at Stn. 50N165E in 1998–2000 (Fig. 7). The vertical distributions of temperature at three stations (49.51N, 1651E; 51.11N, 1641E and 53.51N, 1711E) in 1953 were similar to those observed in 1998–2000 at Stn. 50N165E, with the temperature minimum centered at 100 m. But at two stations located, respectively, near the Kamchatka Peninsula (53.41N, 1611E) and the Aleutian Islands (53.31N, 1661E), the temperature minimum was deeper (B150 m) (Fig. 7). At 100 m the salinity, density and the concentration of phosphate and silicate in 1953 were higher than those in 1998–2000 by B0.23, 0.18 kg m
3, 0.6, and 18 mmol kg
1, respectively. The intercepts in the phosphate-salinity (Fig. 7E) and the silicate-salinity linear relationships com-

5589

puted by 1953 data were higher than those based on 1998–2000 data by B0.2 (B7% in PO4) and 5 mmol kg
1 (B6% in SiO4), respectively. In 1953, the nutrients were measured manually using spectrophotometric methods such as those described by Strickland and Parsons (1972). In 1998– 2000 (R/V Mirai cruises) SiO4 and PO4 in the water samples were determined using a continuous flow analyzer (Gordon et al., 1992). Comparison studies between automated and manual methods show that results from both methods are within experimental deviations (o1–3%) (Berberian and Barcelona, 1979), except at low concentrations (o0.5 mmol l
1). The differences in intercepts in the phosphate-salinity and the silicate-salinity linear relationships computed by 1953 data and those based on 1998–2000 data may be due to systematic errors in the nutrient concentrations of these two data sets. But there was no difference in the values of qCPO4 =qst and qCSiO4 =qst in the halocline of the WSG in 1998–2000 and 1953. Hence, the computed fluxes of the silicate and

Fig. 7. Vertical profiles of density anomaly (st ) (A), temperature (B), salinity (C), silicate (D), phosphate (E), and the phosphate versus the salinity (F) in the WSG. Data of 1998/ 2000 (Stn. 50N, 165E) (2) and data of 1953 (E—E).

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A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

phosphate between the intermediate and surface layers in 1998–2000 by Eqs. (7) and (8) were equal to those in 1953. Because of large interannual variations and sparse data, it is difficult to say whether the difference in the mixed layer salinity and nutrients between 1953 and 1998–2000 was a climate trend or the result of interdecadal oscillations.

We wish to thank MWJ for their efforts in the chemical parameters analysis and the officers and crew members of the R/V Mirai for their assistance with this study. We also thank anonymous referees for valuable comments which helped to improve the paper.

5. Summary

Appendix A. The calculation of the excess CO2

Based on data collected in 1997–2000 (R/V Mirai), we have computed the vertical fluxes of nutrients and carbon (including the flux of excess CO2) between the surface and intermediate layers in the central part of the WSG and close to its southwestern edge. The annual average vertical fluxes estimated by mass balance at the study sites were 1.070.2 mol m
2 for silica, 0.3470.06 mol m
2 for carbonate, 0.4070.10 mol m
2 for TON, and 0.02470.006 mol m
2 for TOP, and 3.270.6 mol m
2 for TOC, respectively. Based on our calculations and sediment-trap data (Honda et al., 2002) we have estimated that B93% of POC and PON exported from the surface layer in the WSG undergo remineralization in the intermediate layer. We have calculated that in the study area the annual uptake of excess CO2 and the change in d13C due to the Suess effect (d13Cantr) were 0.7470.1 mol m
2 and
7.471.0 per mil m, respectively. The rate of change of depth-integrated d13C was slightly lower in the WSG than the estimate for the global ocean (–9.772.4 per mil yr
1 (Sonnerup et al., 1999)). To detect the changes in the vertical fluxes of nutrients (phosphate and silicate) between the surface and intermediate layers of the WSG, we compared the data collected in 1953 with those obtained in 1998–2000. Although in 1953 the salinity, density and the concentration of nutrients at 100 m in the WSG were higher than in 1998– 2000 (at B0.23 in salinity, 0.6 mmol kg
1 in phosphate and 18 mmol kg
1 in silicate), we did not find a significant difference in the computed vertical fluxes of nutrients between the intermediate and the surface layers.

The excess CO2 concentrations in seawater was computed from TA and DIC data (Chen and Millero, 1979):

Acknowledgements

CO2 ex ¼ NDIC
DIC0
ðDDICorg
0:5 DTAorg Þ
0:5ðNTA
TA0 Þ;

ðA:1Þ

0

where DIC is preformed, pre-industrial DIC, TA0 is preformed TA, DDICorg and DTAorg are the change in DIC and TA due to the organic matter oxidation. In the calculation of excess CO2 concentrations in the North Pacific (Chen et al., 1986; Tsunogai et al., 1993), it was accepted that DDICorg
0:5 DTAorg ¼ 0:78AOU; TA0 ¼ a1
3:0y and DIC0 ¼ a2
11:5yq; where AOU is apparent oxygen utilization (the difference between the equilibrium and the measured concentration of DO), a1 and a2 are constants, y is potential temperature. Using the above equations for (DDICorg– 0.5 DTAorg), TA0 and DIC0, Eq. (A.1) can be written as CO2 ex ¼ NDIC
0:5NTA
0:78AOU þ 10yq þ a3;

ðA:2Þ

where a3 is a constant. Since there is no deep-water formation in the North Pacific, penetration of gases of anthropogenic origin like freons and excess CO2 are limited by the lower boundary of the intermediate layer (sy ¼ 27:5227:6) (Chen et al., 1986; Warner et al.,

A. Andreev et al. / Deep-Sea Research II 49 (2002) 5577–5593

1996). The constant a3 in Eq. (A.2) was determined by assuming that excess CO2 in the deep layer (sy X27:6) of the northwestern Pacific was zero.

Appendix B. Application for the calculation of the excess CO2 and 13 dCantr Assuming that in the study area the ratio of the decrease in d13C due to the Suess effect to the increase in d14C was 200, we have estimated that the ratio of the increase in excess CO2 to the d13Cantr decrease (CO2 ex/d13Cantr) was –100 mmol kg
1 per mil (3.2). To check CO2 ex/ d13Cantr ratio and thus verify our assumption (Dd14C/Dd13Cantr was B200), we conducted the independent estimation of the changes in the DIC due to the carbonate dissolution (DDICcarb) in 100–1500 m of the WSG by TA data (Chen et al., 1986) DDICcarb ¼ 0:5 DPTAt ;

ðB:1Þ

where PTAt is (NTA+NNO3+4y); and by DIC and d13C data using Eq. (A.1) DDICcarb ¼ DDIC
DIC0
DDICorg
CO2 ex :

ðB:2Þ

The ratio between the change in DIC and d13C of DIC due to the organic matter flux from the euphotic layer (DDICorg/d13Corg) can be determined by following equation: DDICorg =d13 Cbio ¼ ðDIC=er Þ;

ðB:3Þ

where er is typical value of d13C in organic matter (Broecker and Peng, 1982). According to Mino et al. (2000), er in the suspended POC in the euphotic zone and in sinking POC at 60 m at Stn. KNOT varied seasonally from –20.4 to –27.6 per mil, averaging to B24 per mil. Taking DIC and er to be, respectively, 2400 mmol kg–1 and
24 per mil, by Eq. (B.1) we find that in the WSG the DDICorg/ d13Cbio ratio was B–100 mmol kg–1 mil
1. By DDICorg /d13Cbio and CO2 ex/d13Cantr ratios we get that
DDICorg
CO2 ex ¼ 100 Dd13 C:

ðB:4Þ

5591

Hence the Eq. (B.2) can be rewritten as DDICcarb ¼ NDIC
DIC0
DDICorg
CO2 ex ¼ DNDICt þ 100 Dd13 C;

ðB:5Þ

where DNDICt is (NDIC+11.5y
const) (Appendix A). The total uncertainty (1s) of the DDICcarb calculation by TA data (Eq. (B.1)) (DDICcarb*) and by DIC and d13C data (Eq. (B.5)) (DDICcarb**) was B717 mmol kg
1 (73 mmol kg
1 in 0.5PTAt, 74 mmol kg
1 in NDICt and 710 mmol kg
1 in 100d13C). Fig. 6B shows that DDICcarb* versus DDICcarb** 1s scatter about 1:1 correlation line was less than 710 mmol kg
1. The good agreement between DDICcarb calculated by DIC and d13C with those estimated by TA partly support our conclusion that in the study area the (CO2 ex/d13Cantr) ratio was B
100 mmol kg
1 mil
1.

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