CO2 outgassing off central Chile (31–30°S) and northern Chile (24–23°S) during austral summer...

Deep-Sea Research I 49 (2002) 1413–1429

CO2 outgassing off central Chile (31–301S) and northern Chile(24–231S) during austral summer 1997: the effect of wind

intensity on the upwelling and ventilation of CO2-rich waters

Rodrigo Torresa,*, David Turnera, Jos!e Rutllantb, Marcus Sobarzoc,Tarsicio Antezanad, Humberto E. Gonzaleze

aDepartment of Analytical and Marine Chemistry, G .oteborg University, SE-412 96, G .oteborg, SwedenbDepartamento de Geof!ısica, Universidad de Chile, Casilla 2777, Santiago, ChilecCentro EULA, Universidad de Concepci !on, Casilla 160-C, Concepci !on, Chile

dDepartamento de Oceanograf!ıa, Universidad de Concepci !on, Casilla 160-C, Concepcion, Chilee Instituto de Biolog!ıa Marina ‘‘Dr. J .urgen Winter’’, Universidad Austral de Chile, Casilla 567, Valdivia, Chile

Received 16 February 2001; received in revised form 6 May 2002; accepted 7 May 2002

Abstract

The distribution of pH and alkalinity has been used to calculate the distribution of total inorganic carbon (TC) and

fugacity of carbon dioxide (fCO2) in the upper 200m of the water column in coastal upwelling areas off northern Chile

(23–241S, near Antofagasta) and central Chile (30–311S, near Coquimbo) during austral summer 1997. In these

upwelling areas, colder surface waters were oxygen poor and strongly CO2 supersaturated (100% near Antofagasta and

200% near Coquimbo), although below the pycnocline the CO2 supersaturation invariably exceeded 200% in both

areas. The larger surface CO2 supersaturation and outgassing at 301S were associated with stronger winds that

promoted the upwelling of denser water (richer in CO2) as well as a higher air–sea CO2 transfer velocity. The consistent

decrease in intensity of the southerly winds (as derived from NSCAT scatterometer data) from 30–311S to 23–241S

suggests a corresponding decline in the intensity of the CO2 outgassing due to upwelling. Additionally, we suggest here

that the intensity of the local upwelling forcing (i.e. alongshore–equatorward winds) plays a role in determining the

water mass composition and phytoplankton biomass of the coastal waters. Thus, while deep upwelling of salty and cold

water resulted in high fCO2 (up to 1000 matm) and very low phytoplankton biomass (chlorophyll a concentration lower

than 0.5mgm�3), the shallow upwelling of less salty (e.g. salinity o34.5) and less CO2-supersaturated water resulted in

a higher phytoplankton biomass, which further reduced surface water fCO2 by photosynthesis.r 2002 Elsevier Science

Keywords: Carbon dioxide; Coastal upwelling; Air–sea exchange; Chile

1. Introduction

Wherever surface winds produce an oceanicdivergence (in mid-ocean or off the coast), subsur-face water is transported to the surface. In coastal

*Corresponding author. Department of Marine Chemistry

and Geochemistry, Woods Hole Oceanographic Institution,

Woods Hole, MA 02543-1543, USA. Fax: +1-508-457-2164.

E-mail address: rtorres@whoi.edu (R. Torres).

PII: S 0 9 6 7 - 0 6 3 7 ( 0 2 ) 0 0 0 3 4 - 1

areas this upwelled water is normally cold, rich innutrients and CO2, and poor in oxygen. So, itcould be expected that the initial effect of anupwelling pulse would be upward and downwardair–sea fluxes of CO2 and O2, respectively. How-ever, enhanced primary production due to thenutrient stimulus in the euphotic zone will tend tobalance or even reverse the CO2 and O2 fluxes.

Since the release and sequestering of CO2 in theocean depends on a range of complex processesand factors (e.g. photosynthesis, respiration, car-bonate system chemistry, stratification, mixing,etc.), the resulting balance at any particular placeand time can not be accurately predicted yet. Thus,the direct assessment of CO2 parameters isrequired to identify where, when and how thesefluxes occur and to understand their variability.This is especially true for coastal upwellingsystems, where observations of CO2 parametersare scarce and the spatio-temporal variability offCO2 distribution is typically high (Simpson andZirino, 1980; Copin-Mont!egut and Raimbault,1994; Torres et al., 1999; Geen et al., 2000).

Among the most typical characteristics of east-ern boundary current systems are coastal upwel-ling, a poleward undercurrent associated with theslope and a shallow oceanic oxygen minimumlayer (OML). The strongest OML is located off thewest coast of South America, extending commonlyfrom the lower portion of the thermocline to morethan 400m depth (Antezana, 1978). Near the coastthe OML coincides with the location of the Peru–Chile poleward undercurrent (Wooster and Gil-martin, 1961), whose large latitudinal range wasearly inferred by Brandhorst (1971).

The oxygen-deficient waters of the OML, whichare also strongly CO2 supersaturated, are thereforeadvected polewards and at the same timetransported towards the surface through deepupwelling off Peru (e.g. Copin-Mont!egut andRaimbault, 1994) and off Chile (e.g. Torres et al.,1999). However, the intensity of the upwellingvaries in space and time, sometimes producing aweak upwelling of ‘‘thermocline waters’’ and inother cases producing deeper upwelling of thecolder and saltier water (equatorial subsurfacewater, ESSW) that characterises the polewardundercurrent.

On a scale of weeks to months the thermohalinestructure of the coastal water column can bedetermined by passing coastal-trapped waves(CTW) of remote origin (e.g. Shaffer et al.,1997), but the intensity of the local upwelling isstrongly related to the alongshore wind stress onthe surface ocean. Therefore, large latitudinalvariations in the oceanic winds along the Chile–Peru coast (Bakun and Nelson, 1991; Shaffer et al.,1999) might be expected to produce an analogousvariability in the upwelling and ventilation of CO2-rich waters. A full test of this hypothesis wouldinvolve a synoptic, long-term comparison of theCO2 air–sea fluxes in upwelling centres scatteredalong the Chile–Peru coast, corrected for theremote forcing of the upwelling that propagatespolewards along the coast.

We present here a first exploratory test of thishypothesis based on the results of two surveys, onecarried out near Antofagasta (northern Chile,231S) and one near Coquimbo (central Chile,301S), during January and February 1997.

2. Methods

2.1. Oceanographic stations

Discrete seawater samples and CTD data werecollected onboard R.V. Abate Molina duringaustral summer 1997 off northern Chile (22.7–24.01S, January 1997) and central Chile (29.9–30.71S, February 1997) (Fig. 1A). The cruise offnorthern Chile, near Antofagasta, consistedmainly of the two grid surveys G1 and G2(Fig. 1B). The grid surveys comprised 31 stationseach arranged as five 140 km long east–westtransects covering an area of approximately21 000 km2 (Fig. 1B). The cruise off central Chile,near Coquimbo, was divided in two parts: a200 km ocean–coastal transect close to 301S, anda coastal grid survey (G3) with 42 stationsarranged as six east–west transects covering anarea of approximately 5000 km2. At each station(in both areas) Neil Brown Mark III CTD-oxygen-fluorescence rosette casts were carried out; ex-ceptionally on G3 (on 6 January 1997 at stations42–41 and 38–35; Fig. 1) salinity and temperature

R. Torres et al. / Deep-Sea Research I 49 (2002) 1413–14291414

only were recorded with a SBE25 CTD. The CTDresults are reported here only for the upper 200m.Discrete seawater samples for pH, total alkalinity(TA), and phytoplankton pigments were collectedon both cruises. On the Antofagasta survey and onthe westward transect off Coquimbo seawatersamples were collected for pH and TA at 9–11depths in the upper 200m. On the G3 surveyseawater samples were collected only at 2, 10 and30m depth.

2.2. Measurement of pH and TA, calculated fCO2

and air–sea CO2 fluxes

Seawater samples were collected in 5 l Niskinbottles mounted on the CTD rosette. Samples forpH measurement were taken in 50ml syringes witha Tygon tube and then transferred to a closed25ml cell thermostated at 2570.11C. pH wasmeasured with a Metrohm 713 pH meter (inputresistance >1013O, 0.1mV sensitivity and nominal

resolution 0.001 pH units) and a glass combineddouble junction Ag/AgCl electrode (Metrohmmodel 6.0219.100) calibrated with 8.089 tris and6.786 2-aminopyridine buffer (DOE, 1994) at2570.11C; pH values are therefore reported onthe total hydrogen ion scale (DOE, 1994). Theoverall uncertainty in the measured pH values wasestimated by Torres et al. (1999) as 0.006 pH forsurface waters (pH near 8) and o0.009 pH forvery acid waters (pHE7.2). Samples for TA werecollected at different depths in the 0–200m depthrange. Samples of about 250ml were drawn intohigh-density polypropylene bottles, fixed with20 ml of saturated mercuric chloride solution, andstored in darkness for 3 weeks before analysis byan automated potentiometric titration method(Haraldsson et al., 1997). The TA values fromthe 0–200m depth range were used to determine alinear regression between TA and salinity in orderto predict the alkalinity for the remaining stationsin the same depth range. The regression obtained

71.672.0

242220 21

P.Angamos

70.8 70.471.271.6

36 3937 38

403234

26 27 2825

111098 12

9 68 7 3 1

11 12 17

151719 18 16

282931 30

357 6 4

717374 72

Antofagasta

Coquimbo

T4T2T1

Longitude (˚W)

(A) (C)

Fig. 1. (A) Location of the study areas. (B) The Antofagasta study area showing stations in grids G1 (12–15 January 1997) and G2

(22–26 January 1997). (C) The Coquimbo study area showing stations in grid G3 (6–9 February 1997) and in the across-shore transect

(10–11 February 1997; outer stations shown in (A)). White dots mark the location of the coastal meteorological stations; solid lines

show the 200, 400 and 600m isobaths.

R. Torres et al. / Deep-Sea Research I 49 (2002) 1413–1429 1415

was virtually the same for both study areas: TA(mmol kg�1)=2327� (salinity/35) in the salinityrange 34.2–35 (Fig. 2). In order to assess the errorin the other (calculated) carbonate system para-meters, we use the average of the absolutedifference between measured and estimated TAvalues (574 mmol kg�1 for Antofagasta and774 mmol kg�1 for Coquimbo) as the uncertaintyof each extrapolated surface TA value. This givesuncertainties of the order of 9 mmol kg�1 for totalinorganic carbon (TC) and 9 matm for fCO2

(Torres et al., 1999). The fCO2 was calculatedusing Roy et al., 1993 (K1 and K2) and Dicksonand Riley, 1979 (K0) constants, at in situtemperature and pressure. The net flux of CO2

was calculated from the equation given by Andri!eet al. (1986): FCO2 ¼ kaDfCO2; where k is theCO2 transfer velocity (Liss and Merlivat; 1986), ais the solubility of CO2 in seawater (Weiss, 1974)and DfCO2 is defined as the difference in fCO2

between the surface water and atmosphere. Amean atmospheric fCO2 of 352 matm (Lef"evre,personal communication, 1997) was used for thecomputation of DfCO2 in both areas. The physicalforcing of CO2 transfer is expressed as the productka; the transfer coefficient. The error in the FCO2

estimates are expected to be large mainly due tothe uncertainties in the bulk estimation of k

(Wanninkhof and McGillis; 1999). We used theformula given by Liss and Merlivat (1986) which isin good agreement (within the 0–8m s�1 wind

speed range) with the recent parameterisationsuggested by Wanninkhof and McGillis (1999).

2.3. Phytoplankton pigments

The concentration of chlorophyll a was deter-mined with Turner designs fluorometers Model 10-AU, for the Antofagasta cruise and Model 110 forthe Coquimbo cruise. Both fluorometers werecalibrated with pure chlorophyll a (extracted fromspinach; Sigma). Water samples were filteredthrough Whatman GF/F filters, the filters wereextracted in 90% acetone in the dark at 41C for24 h, and the fluorescence was then measured.Chlorophyll a was calculated following Parsonset al. (1992). During the Antofagasta surveys (G1and G2) the sampling depths were normally atapproximately 2, 10, 20, 40, and 100m. During theCoquimbo survey (G3) samples were collected at2, 10, and 30m depth.

2.4. In vivo fluorescence

A laser-stimulated fluorescence system mountedon the rosette recorded the in vivo fluorescence inevery CTD-rosette cast. The concentration ofchlorophyll a (measured in discrete samples asdescribed above) was found to be well correlatedwith the corresponding in situ/in vivo fluorescencerecord. A regression model

½chlorophyll a� ðmgm�3Þ

¼ cþ in vivo fluorescence ðVÞb

was fitted to both the Antofagasta and Coquimbodata sets, giving b ¼ 0:048 and c ¼ �0:12(n ¼ 310; r ¼ 0:9; p-value=0.000) for Antofagas-ta, and b ¼ 0:044 and c ¼ �0:44 (n ¼ 50; r ¼ 0:9;p-value=0.000) for Coquimbo.

2.5. Apparent oxygen utilisation (AOU)

AOU is defined as the difference between theoxygen concentration at 100% saturation and themeasured oxygen concentration. The oxygen con-centration at 100% saturation was calculated,following Murray and Riley (1969), from CTDmeasurements of temperature and salinity, and

Fig. 2. The relationship between TA and salinity in the upper

200m during austral summer 1997.

oxygen concentrations were obtained from theoxygen sensor (calibrated against Winkler method;Williams and Jenkinson, 1982) mounted on theCTD.

2.6. Wind speeds

Coastal winds were measured at two well-exposed coastal sites: Caleta Constitucion at 231S(Fig. 1B, at B7m above the sea surface) andPunta Lengua de Vaca at 301S (Fig. 1C, at B15mabove the sea surface). Measurements were per-formed with RM Young Wind-Sentry anemo-meters on 4m masts and averaged (mean andresultant wind speed and direction) every 30min ascomponents using Campbell automatic meteoro-logical stations. Satellite-derived ocean surfacewinds computed at a height of 10m above thesea surface (from the NSCAT-ADEOS scattero-meter) were obtained from the D !epartment d’Ocea-

nographie Spatiale, IFREMER, France (http://www.ifremer.fr/droos). These data are distributedon a WOCE CD-ROM at 0.51� 0.51 resolution.The validation of the NSCAT-ADEOS windmeasurements is described by Atlas et al. (1999).

2.7. Pathfinder sea surface temperature

Sea surface temperature (SST) data, derivedfrom the advanced very high resolution radiometer(AVHRR) on board the NOAA Polar Orbiters,were obtained from the Jet Propulsion Laboratory(JPL) by ftp (http://podaac.jpl.nasa.gov/sst/). Adescription of the measurements and of thevalidation of this data set can be found in Smithet al. (1996).

2.8. Sea level

Daily values of ‘‘research quality’’ data from thetide gauges at Antofagasta (23.61S) and Caldera(27.01S) were obtained by ftp from the Sea LevelCentre at the University of Hawaii (http://uhslc.soest.hawaii.edu). Quality control and data pro-cessing details can be found at http://www.soest.hawaii.edu/kilonsky/uhslc.html.

2.9. Subsurface temperature at 65m depth

These data were recorded off Mejillones penin-sula at a mooring placed near station 14 (seeFig. 1B).

2.10. Water column stratification

The intensity of stratification of the watercolumn was assessed from the geopotential energyanomaly (F50), defined as (Bowden, 1983)

ðrm � rÞgz dz;

where rm is the mean density of the water columnfor each station, r is the density at depth z; g is theacceleration due gravity and H ¼ 50m. Thisparameter quantifies the depth-averaged deficit ofpotential energy due to stratification, comparedwith the potential energy of a completely mixedcolumn. Small values of F50 indicate low stratifi-cation.

Water column stability was also assessed fromthe Brunt–Vaisala frequency N; which is ameasure of water column stability defined by

N2 ¼ gr�1 qr=qz:

Calculated values of N were smoothed by meansof a polynomial fit.

3. Results

3.1. Time series of winds and hydrographic

variables

In the two study areas (231S and 301S) both thecoastal and oceanic winds were upwelling favour-able during the study period (Table 1 and Fig. 3Aand B), but stronger at 301S than at 231S (Table 1).The offshore intensification of the meridional windcomponents also promoted upwelling in bothareas (Table 1). The coastal equatorward windsfollow a prominent diurnal cycle, particularly inthe austral summer, because they are partiallyforced by land–sea temperature differences. Asexpected, the coastal meridional wind componentswere better correlated with the oceanic ones

(NSCAT) near the coast than at more offshorelocations (Table 1).

The upwelling favourable winds along thecoast from 23–301S (near shore NSCAT winds)followed a synoptic intensification–relaxationcycle (Fig. 3A), where the major peaks in thesurface pseudo-stress coincide in the two studyareas (Fig. 3B). The wind-driven upwelling pulsescan be expected to be more effective at 301S wherethe variability of the offshore equatorward windswas substantially greater (Table 1).

The nearshore (o10 km from the coast) satelliteSST observations were in general poorly corre-lated with the meridional winds at each latitude(0.51 resolution analysis in the range 20–301S;results not shown here). However, the correlationwas statistically significant and consistently nega-tive from 281S to 301S, confirming that the moreenergetic variations in the upwelling wind forcingwithin the 28–301S latitude range cause larger SSTperturbations, in spite of the large temporal andspatial gaps in the coastal SST data (Fig. 3C).

Variations in sea level (SL) were not correlatedwith the local winds (coastal or oceanic) or withSST during the study period (Fig. 3A–D). Verysimilar SL fluctuations were observed at tide gaugestations at Arica, Antofagasta and Caldera(Fig. 3D, Arica at B18.51S is not shown here).However, the SL oscillations were strongly corre-lated with the subsurface (65m depth) temperatureover the slope off the Mejillones Peninsula(r ¼ 0:8; p-value=0.000, n ¼ 22), suggesting asignificant variation in the thermohaline structurein the water column over the slope during thestudy period.

3.2. Differences in CO2 transfer coefficient between

231S and 301S

Because of the consistent heading of the off-shore winds, the average alongshore equatorwardwinds (Fig. 3A) are strongly correlated with theaverage wind speed. Therefore, fluctuations in theCO2 transfer velocity (a function of wind speed)

Table 1

Averages of daily measurements of the coastal and oceanic equatorward component of wind speed (S wind), total wind speed, and

calculated CO2 transfer coefficient (ka) in the study areas during the first 45 days of 1997. The degree of correlation between coastal

and oceanic S are showed in the last column, calculated on the basis of corresponding coastal and oceanic daily observations

S wind (m s�1) Total wind (m s�1) ka (mmolm�2 yr�1 matm�1) n Pearson coefficient

Coastal winds 1.771.4 3.871.1 676 45

NSCAT winds

43–99 km offshore 2.471.7 2.671.5 375 20 0.7

99–154 km offshore 3.371.3 3.471.3 576 31 0.5

154–210km offshore 3.571.5 3.671.4 677 30 0.3

210–266km offshore 3.771.7 3.971.5 878 29 0.3

266–321km offshore 3.971.8 4.271.7 10710 31 0.2

321–377km offshore 4.271.8 4.671.7 12710 32 �0.1

Coastal winds 2.271.6 4.371.1 977 39

NSCAT winds

56–111 km offshore 5.572.0 5.872.0 21714 20 0.5

111–167km offshore 5.672.2 5.872.1 21714 23 0.2

167–222km offshore 5.872.7 6.172.5 24717 24 0.0

222–278km offshore 6.172.7 6.372.5 26718 26 �0.1

278–333km offshore 5.873.0 6.172.8 24720 24 �0.0

333–389km offshore 5.873.6 6.373.0 25723 26 �0.1

will be correspondingly large at the upwelling-event scales (Fig. 3E). The CO2 transfer coefficientka between about 50 and 100 km offshore is about5 times higher at Coquimbo (301S) than atAntofagasta (231S; Table 1). Although the differ-ence is smaller at the coast (Table 1), it isstatistically significant (the Wilcoxon signed rankstest, p-value=0.03, n ¼ 39).

3.3. Hydrography

The T–S and T–fCO2 diagrams (Fig. 4) for bothcruises were consistent with previous observations

(Torres et al., 1999). In both areas maximum fCO2

values occur within the core of the salty and coldESSW, which is advected poleward by the Peru–Chile undercurrent (Wooster and Gilmartin, 1961;Shaffer et al., 1995). Above the ESSW, amoderately cold and low salinity water mass(Subantartic water; SAAW) is often present. Thepresence of the SAAW can be traced by a shallowsalinity minimum, because this water mass(SAAW) sinks under the less dense, saltier andwarmer Subtropical water (SSW) when movingequatorward from the rainy higher latitudes (Reid,1973). Since ESSW is located below SAAW, a

Fig. 3. (A) Nearshore meridional winds (m s�1) derived from NSCAT scatterometer data. (B) Afternoon (1200–2400) alongshore

pseudo-windstress at coastal meteorological stations. (C) SST (1C) within 10 km of the coast, derived from Pathfinder satellite data;

dots mark the location of 9� 9 km2 resolution observations. (D) Normalised SL at the Antofagasta (23.71S) and Caldera (27.01S) tide

gauges. Also shown is the sea temperature recorded at 65m depth off Antofagasta. (E) CO2 transfer coefficient ka estimated from

nearshore NSCAT wind speed data (mmolm2 yr�1 matm�1).

deep upwelling involves both water masses(Brandhorst, 1971).

3.4. Transects

A synoptic east–west transect off the MejillonesPeninsula (23.31S) on January 14 (Fig. 5, G1)shows a weak upward deflection of the isopycnalsnear the coast in the upper 50m. However, belowthis depth the isopycnals tended to slope downfrom stations 16–14 (i.e. 26.2 and 26.4 isopycnals).The water column was highly stratified in theupper 20m. The layer of >0.5mgm�3 chlorophylla (Fig. 5 shaded area) was located within thethermocline although at oceanic stations theuppermost warm layers (10–20m depth) had aparticularly low in vivo fluorescence and chlo-rophyll a (B0.2mgm�3). At these stations thehighest chlorophyll and fluorescence occurredbetween 40 and 70m depth (Fig. 6A) associated

with the lower temperature gradient (16–131C) justabove the shallow salinity minimum. The subsur-face maximum of chlorophyll a and in vivofluorescence tended to rise toward the coast, wherea sharp fluorescence peak covering only a fewmetres depth range was observed (Fig. 6A).A shallow subsurface dissolved oxygenmaximum (negative AOU) at about 20m depthwas associated with a relatively high pH(7.9–8.0) and low fCO2 (near equilibrium withthe atmosphere), above the fluorescence max-imum. In the uppermost layer, the saltier andwarmer water showed a small but persistent CO2

supersaturation.During the second survey (G2), the rise of the

shallow and deep isopycnals intensified. Colderand denser water seems to migrate offshore (e.g.the 24.4 isopycnal in G1 and G2). However,although near the coast (0–60 km offshore) theupper 50m was colder, offshore waters (80–140 km

352 704 1056

34.0 34.2 34.4 34.6 34.8 35.0

81012141618202224

ospheric fCO

Salinity

f CO2(µatm)

ESSW : Equatorial Subsurface Water.SAAW : Subantartic Water.

SSW: Subtropical Water.

Fig. 4. T–S and T–fCO2 relationships. Circles and dots depict observations at 231S and 301S, respectively.

offshore) increased in temperature, especiallybetween 15 and 60m depth. In general, the surfacecoastal cooling was associated with a salinitydecrease and conversely the surface oceanicwarming (in the upper 40m) was associated withsaltier waters. The oceanic deepening of theshallow salinity minimum during G2 appears tobe associated with the deepening of the fluores-cence maximum (Figs. 5 and 6A) and of the AOU/fCO2 isolines (Fig. 5). Similarly, the upwelling ofless salty water (B34.4) near the coast in G2 wasassociated with an increase in surface fluorescenceand chlorophyll a, and with the occurrence ofparticularly low surface AOU and fCO2 values atstations 16 and 17. However, below the veryshallow coastal thermocline (i.e. below 20m depthat station 14), the very salty (>34.6) and coldwater was extremely supersaturated in CO2

(>200%, Fig. 5) and very poor in fluorescence(Fig. 6A) and in oxygen (Fig. 5).

The east–west transect carried out off Coquim-bo shows the deflection of isopycnals toward thecoast, which was particularly intense betweenstation 7 and the coast (Fig. 5). The denser surfacecoastal water (st > 26:0; station 9) was cold(o131C), salty (>34.5), poorly oxygenated, lowin pH, strongly CO2 supersaturated (up 200%)and particularly low in chlorophyll a and fluores-cence (Figs. 5 and 6). The maximum values ofchlorophyll a and fluorescence (e.g. Fig. 6, station34) occur near the upwelling divergence (densercoastal water), and the most oxygenated(AOUp0) and CO2-poor surface waters(fCO2p352 matm) were observed only at the morestratified offshore stations (>50 km offshore;stations 4 and 2; Fig. 5). Note that in these two

Fig. 5. Across-shore sections from grids G1 and G2 and from the long transect at Coquimbo (T stations, Fig. 1). Station numbers are

marked at the top of the upper panels. The shaded area in each section depicts the location of the in vivo fluorescence maximum, which

corresponds approximately to concentrations of chlorophyll a >0.5mgm�3. Solid dots show the discrete sampling depths.

study areas the location of the offshore chloro-phyll a and fluorescence maxima occurred belowthe surface, associated with the thermocline(13–171C) and above the shallow salinity minimum.

3.5. Grid surveys

The first grid survey (G1, Fig. 7) was carried outunder relatively weak upwelling favourable winds(Fig. 3A and B). The SST gradient was about 4–51C associated with a cold filament centred at ca.23.6–23.71S (Sobarzo and Figueroa, 2001). Thecold coastal waters were less salty (o34.5) than the

oceanic waters (>34.7), consistent with theupwelling of Subantartic water (Fig. 5). In general,the oceanic surface waters (warm and saltysubtropical water) were weakly undersaturated inO2 and supersaturated in CO2. Near the coastminimum AOU and fCO2 values where observed.

During the second grid survey (G2), coldercoastal water was observed, which coincided withthe intensification of the upwelling favourablewinds (Fig. 3B). However, it is probable that thecoastal cooling of the entire 0–200m depth range(Fig. 5) was not solely due to intensification of thecoastal meridional winds (Fig. 3B); see discussion

Sta. 19 (~140km offshore) Sta. 17 (~47km offshore) Sta. 14 (~3 km offshore)

Sta. T5(~54km offshore)

Sta. 34(~30km offshore)

Sta. 42(~6km offshore)

Chlorophyll a (mg Chla m-3)0 1 2 3 4 5 6 7 8 9

Chlorophyll a (mg Chla0 1 2 3 4 5 6 7 8 9

m-3) m-3)

m-3)m-3)m-3)

Fig. 6. Vertical profiles of chlorophyll a (symbols) and in vivo fluorescence at selected stations at 231S (A) and at 301S (B).

in Section 4.3 below. The coastal cooling wasassociated with a decrease in surface salinity(Fig. 7). The coldest and least stratified coastalwaters off Mejillones Peninsula were undersatu-rated in O2 and supersaturated in CO2. However, afew kilometres offshore, in more stratified water,the surface water showed minima in AOU andfCO2. Similarly low AOU and fCO2 were foundnorth of Mejillones Bay with a similar degree of

vertical stratification and chlorophyll a concentra-tion (Fig. 7). In general, less salty surface water(o34.5) was characterised by high fluorescenceand chlorophyll a values. The warm and saltysurface water (normally located at oceanicstations) remained CO2 supersaturated andundersaturated in oxygen as in G1, withcharacteristically low levels of surface chlorophyll(o0.5mgm�3).

Temperature (˚C)

Salinity

Chlorophyll a (mg m )-3

f CO2 (µatm)

25˚CpH (pH units)

AOU (µmol Kg )-1

71.8Longitude (˚W)

71.4 71.0 70.6

71.8Longitude (˚W)

71.4 71.0 70.6

Temperature (˚C)

Salinity

Chlorophyll a (mg m )-3 f CO2 (µatm)

25˚CpH (pH units)

AOU (µmol Kg )-1

71.8Longitude (˚W)

71.4 71.0 70.6

71.8Longitude (˚W)

71.4 71.0 70.6

(J m )-3φ50

G2G1 G2

>400<400

(J m )-3φ50

Fig. 7. Surface distributions of temperature, salinity, chlorophyll a, AOU, pH at 251C, fCO2 and F50 for grids G1, G2 and G3. Filled

circles show the sampling stations. Open circles in the G3 chlorophyll a plot indicate that chlorophyll a was not measured, but

estimated from in vivo fluorescence.

The grid survey off Coquimbo (G3) was carriedout under relatively strong upwelling favourableoceanic winds (Fig. 3A and B). The shape of thesurface gradient in SST was similar to those ofsalinity, AOU, pH25, fCO2 and stratification(Fig. 7). The coldest and saltiest surface water(which was essentially ESSW, see Section 3.3) atpoorly stratified locations was characterised bymaximum fCO2 (exceeding 1000 matm) and mini-mum oxygen and chlorophyll a concentrations

(Fig. 7). In general, the highest chlorophyll a

concentrations are associated with moderate stra-tification (F50E40 Jm�3), and low and highstratification levels are associated with low surfacechlorophyll a (Fig. 7B). A different situation isobserved for fCO2 and AOU, which reach mini-mum values (i.e. undersaturation in CO2 andsupersaturation in O2) only in more stratifiedoffshore water. In general the coastal stratificationat 301S was weaker than at 231S (Fig. 7), which is

Temperature (˚C)

72.4 72.0 71.6

Longitude (˚W)

72.4 72.0 71.6

Longitude (˚W)

Salinity

Chlorophyll a (mg m )-3 f CO2 (µatm)

25˚CpH (pH units)

AOU (µmol Kg )-1

-20-20

(J m )-3φ50

Fig. 7 (continued).

also evident in profiles of the Brunt Vaisalafrequency (Fig. 8): note that the upper 10m weresubstantially more stratified at Antofagasta thanat Coquimbo.

3.6. CO2–density relationship

The surface density gradient near the coast is anindicator of the upwelling. We show here therelationship between density and CO2 in order tocompare the two upwelling areas (G1–G2 vs. G3)in the context of phytoplankton biomass andwater mass composition. Fig. 9 shows that bothupwelling areas (Antofagasta and Coquimbo)share a common surface density range. In thisrange the TC–density relationship and the occur-rence of high phytoplankton biomass are analo-gous, if we assume that the particularly low TC inthe densest water is an effect of local primaryproduction. The largest differences arise from theabsence of high density surface water at 231S(ESSW, see Section 3.3) and the absence of lighter,salty water at 301S (SSW). In general, the largerscatter in the TC–density relationship at 231Ssuggests a more effective reduction of surface fCO2

by biology, which was particularly evident in theless dense and less salty surface waters in thehighly stratified coastal upwelling front (Fig. 9A).The lowest scatter in the TC–density and fCO2–density relationships was observed in dense, poorlystratified water at 301S, which was characterisedby a very low phytoplankton biomass.

3.7. The mean CO2 air–sea flux

In general, during active upwelling indicated bystrong coastal SST gradients (G2 and G3; Fig. 7)the surface fCO2 (integrated and normalised byarea) increased towards the coast within the first35 km (Fig. 10A). In the coastal 25 km the DfCO2

was about 3 times higher at 301S compared to231S. In addition, the CO2 transfer coefficient kain this area (the coastal 25 km) was approximatelytwice as large at 301S than at 231S (Fig. 10B;interpolated from coastal and offshore windrecords, Table 1); thus the resulting FCO2

(DfCO2 � ka) was about 5.5 times higher at 301Scompared to 231S (Fig. 10C). Further offshore(35–85 km from the coast) at 231S, the DfCO2

tends to increase towards the open ocean (as thesurface water becomes warmer and saltier), and at301S the surface waters tend to be CO2 under-saturated. Since the integrated FCO2 during activeupwelling is a function of the distance offshore, acomparison between the two upwelling areas willbe highly dependent on the offshore limit ofcoastal waters. In the case of Antofagasta (231S)a sharp surface salinity front separates colder andchlorophyll-rich coastal surface waters from war-mer and oligotrophic oceanic waters (subtropicalwaters), providing a clear coastal boundary. In thecase of Coquimbo (301S) this difference is lessevident because gradients in surface salinity wereless intense (with the exception of the inner-shelffront; see Smith, 1995). We will consider the

0 5 10 15 20 25 30 35

00 5 10 15 20 25 30 35

00 10 20 30

-1N (h )

G3(Coquimbo)

G2(Antofagasta)

G1(Antofagasta)

0 10 20 30

-1N (h )

0 10 20 30

-1N (h )

Fig. 8. Vertical distribution of the Brunt V.ais.al.a frequency in grids G1–G3 for stations located o40 km from the shore.

coastal area enclosed by the baroclinic Rossbyradius of deformation (30–40 km at these latitudes)as the area directly affected by upwelling, in orderto compare the two areas. Thus we suggest that aCO2 outgassing is the direct consequence ofupwelling in both areas, although part of theCO2 released to the atmosphere was reabsorbed inlow salinity, stratified surface waters.

4. Discussion

4.1. Latitudinal variation in the CO2 outgassing

caused by coastal upwelling

Along north-central Chile, oxygen-poor andCO2-rich waters are located below the pycnocline.Thus the erosion of the pycnocline near the coastby upwelling and mixing cause a drastic CO2

outgassing near the coast. The average alongshoreequatorward winds, which promote upwelling,have previously been found to increase fromminimum values at about 201S to maximum valuesat about 301S (Bakun and Nelson, 1991; Shafferet al., 1999), in agreement with our observations.This wind pattern explains the fact that the coastalsurface water was denser (Fig. 9) and morestrongly CO2 supersaturated at 301S compared to231S (Fig. 10A). The combination of a larger air–sea CO2 disequilibrium (DfCO2) and higher windspeeds (larger CO2 transfer coefficient) at 301Scompared to 231S (Fig. 10) produces an evenlarger FCO2 difference within the coastal 35 km(Fig. 10C). Thus we suggest that latitudinal varia-tions in the nearshore wind regime producelatitudinal differences in the CO2 outgassing bymodifying both the fCO2 of the recently upwelledwater and its ventilation rate. This statement

23.6 24.0 24.4 24.8 25.2 25.6 26.0 23.6 24.0 24.4 24.8 25.2 25.6 26.0 23.6 24.0 24.4 24.8 25.2 25.6 26.0

(kg m )tσ -3 (kg m )tσ -3 (kg m )tσ -3

(kg m )tσ -3 (kg m )tσ -3 (kg m )tσ -3(A)

Fig. 9. TC, fCO2 and salinity vs. density (st) relationships for surface waters at 231S (A) and 301S (B). Crosses within the circles depict

high phytoplankton biomass (>2mg chlorophyll a m�3).

should consider also differences in the high short-term variability caused by pulsed upwellingfavourable wind events in both areas (Fig. 3A),biological activity (see Gonz!alez et al., 1998;Daneri et al., 2000) and a complex coastalcirculation as we will discuss later.

4.2. Biological-physical control of surface fCO2

At Coquimbo, the strong correlation betweendensity and surface fCO2 (Fig. 9B) can be inter-preted as the dominance of the mixing vs. the localprimary production in the area most directlyaffected by upwelling. However, the occurrenceof CO2 undersaturation in more stratified waters,as observed off Antofagasta, suggests that vertical

stratification is the key factor promoting biologicalreduction of the surface fCO2 (see Section 3.6).Near the coast off Mejillones peninsula (23–241S),the biological activity was most probably en-hanced by the export of phytoplankton frominshore waters (bays) and the upwelling of‘‘thermocline-phytoplankton rich’’ waters. There-fore, a moderate surface stratification and arelatively high phytoplankton biomass couldeventually produce a rapid CO2 uptake, assuggested by the occurrence of low fCO2 in densewater (Fig. 9A). Conversely, a small CO2 uptake(see Daneri et al., 2000) in phytoplankton-poordenser upwelled waters off Coquimbo can explainthe smaller deviations from the fCO2–density trendcaused by mixing (Fig. 9B).

4.3. Origin of the CO2 outgassing variability: local

and remote forcing

Off Coquimbo a high short-term variability (onscale of days) of the SST and FCO2 has beenassociated with drastic changes in the alongshorecoastal winds (Torres et al., 1999) in connectionwith the cycles of active upwelling and relaxation(sensu Send et al., 1987). A similar explanation(short-term variability in the wind forcing) holdsfor the cooling and concomitant increase in coastalfCO2 between G1 and G2 off northern Chile(Fig. 7A).

The connection between winds and SST varia-bility is expected since the wind stress on thesurface ocean is a major forcing function of theupwelling. Others studies of the wind regime atAntofagasta have shown that the wind variabilityaffects the thermohaline structure of the coastalwater at near synoptic scales (Pizarro et al. 1994;based on SL and wind time series). However, onlonger time scales (weeks–months) the SL varia-bility in this region has been associated with thepoleward propagation of CTW (Pizarro et al.,1994; Shaffer et al., 1997). During the long surveyoff Antofagasta we noticed that the temperaturevariation at 65m depth (T65) was stronglycorrelated with the major SL fluctuations observedalong the central-north Chile coast (Fig. 3D).However, no connection was found between SLor T65 and the local winds (Fig. 3).

O(µatm

µatm)

oefficientα

80 60 40 20 0

Distance offshore (km)

G2 (23˚S)

G3 (30˚S)

Fig. 10. Mean across-shore distributions of DfCO2 (A),

transfer coefficient ka (B) and FCO2 (C) for active upwelling

conditions at 23–241S (G2) and 31–301S (G3). The mean

DfCO2 values were derived from the surface DfCO2 distribu-

tions shown in Fig. 7 assuming a constant atmospheric DfCO2

of 352matm (Lef"evre, personal communication, 1997). The

mean transfer coefficient ka was interpolated from the coastal

and NSCAT wind speed data for the first 45 days of 1997 (see

Table 1).

This leads to the question: Do the intraseasonaloscillations of the vertical thermohaline structuremodulate the entrainment of CO2-rich water to thesurface? Although the Antofagasta survey was notaimed at studying the effect of CTW on the surfaceproperties, a first approximation to evaluating thepotential effect of those oscillations on surfacefCO2 can be made from the relationship betweendensity and fCO2 (Fig. 9B). These data show thatfCO2 rises more than 300 matm as density increasesfrom 25.8 kgm�3 to 26.2 kgm�3 (which corre-sponds approximately to the observed range ofvariability of T65). This last observation implythat the fCO2 at 65m depth can be affected bypassing CTW, thus it can be expected that theCTW activity provoke variations in the CO2

outgassing as the subsurface water (e.g. from65m depth) is transported to surface by upwelling.

However, the CO2 ventilation of subsurfacewater is largely controlled by mixing and upwel-ling, so that the surface manifestation of thoseoscillations (CTW) could be effectively masked bythe local wind regime. For the specific case of G1and G2, we suggest that the combination of bothfactors, i.e. a non-locally forced cooling of theentire coastal water column (Fig. 5), together withan intensification of the local wind-forced upwel-ling (Fig. 3A and B), may better explain the sharpdrop in coastal SST and the increase of coastalfCO2 observed from G1 to G2 (Fig. 7A).

Acknowledgements

We thank Madeleine Hamam!e, Carolina Paradaand Alvaro Sotomayor for chlorophyll a analysesfrom the Coquimbo survey. We thank OscarPizarro and Nathalie Lef"evre for comments andsuggestions. We thank the Department of Inor-ganic Chemistry and EULA-chemistry Laboratoryat the University of Concepci !on for provision oflaboratory facilities. We thank the D!epartment

d’Oceanographie Spatiale IFREMER France, SeaLevel Center at the University of Hawaii, and theJet Propulsion Laboratory (JPL, NASA) for thesupply of hydrographic data. This work wasfunded mainly by SAREC (Sweden, through theJGOFS-Chile project) and CONICYT (Chile,

through FONDECYT 5960020 and 5960002projects). Additional financial support was pro-vided by G .oteborg University, the Program forAtmosphere and Climate Dynamics (PRODAC)at the University of Chile, and the FONDAPHumboldt Program.

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