Biogenic silica production and the diatom contribution to primary production and nitrate uptake in the eastern equatorial Pacific Ocean

Deep-Sea Research II 58 (2011) 434–448

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Biogenic silica production and the diatom contribution to primary production and nitrate uptake in the eastern equatorial Pacific Ocean Jeffrey W. Krause a,n, David M. Nelson a,1, Mark A. Brzezinski b a b

College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, OR 97331, USA Department of Ecology, Evolution & Marine Biology and the Marine Science Institute, University of California, Santa Barbara, CA 93106, USA

a r t i c l e in f o

a b s t r a c t

Article history: Received 9 August 2010 Accepted 9 August 2010 Available online 19 August 2010

To quantify biogenic silica production rates and estimate diatom contribution to primary production and nitrate uptake we examined the distribution of silicic acid, biogenic silica and silica production along zonal and meridional transects between 1101W and 1401W in the upwelling zone of the eastern equatorial Pacific during December, 2004, and September, 2005. Silicic acid concentrations [Si(OH)4] in the upper 75 m were consistently between 2.5 to 4.0 mM but were twice as high near the equator at 1101W, compared to 1401W, consistent with a strong eastward shoaling of the nutricline. Euphoticzone integrated biogenic silica concentrations [bSiO2] had a narrow range, averaging 9.8 7 1.4 mmol Si m  2 during 2004 and 10.4 7 4.0 mmol Si m  2 in 2005. Minima in [bSiO2] occurred within the equatorial undercurrent on all transects; creating atypical profiles where [bSiO2] decreased with depth from the surface and then increased below the equatorial undercurrent. Specific production rates of biogenic silica (Vb) were highest (  0.2-0.3 d  1) in the upper 50 m, decreased to o 0.05 d  1 at the base of euphotic zone, and were typically near zero at 150 m. Unlike other open ocean systems, there was a strong diel cycle in silica production in the euphotic zone as daytime rates were consistently 2-3 times higher than night rates. Vb was significantly correlated with ambient [Si(OH)4], but only when rate data were separated by the fraction of photosynthetically active radiation penetrating the sea surface that was present at each sampling depth (%I0). This interaction between [Si(OH)4] and light on Vb was highly systematic, with the extent of Si limitation of Vb being a linear function of %I0. Consistent with these effects, the average Vb at stations within one degree of the equator doubled from 0.11 to 0.21 d  1 between 1401W and 1101W paralleling the doubling of [Si(OH)4]. Vertically-integrated silica production rates (per day) tended to be higher within the zone of active upwelling (within 7 11 latitude of the equator), but high rates were also observed south of 21S. Variability in silica production rates along zonal transects showed spatial coherence with changes in the meridional velocity flow field forced by tropical instability waves. Despite such perturbations the mean integrated rate of silica production was very similar between cruises, 1.6 70.6 mmol Si m  2 d  1 in 2004 and 1.3 7 0.7 mmol Si m  2 d  1 in 2005; consistent with the idea that phytoplankton rate processes are much less variable in the equatorial Pacific than in other systems. Comparison of silica production rates to rates of C and N use indicates that diatoms were responsible for  18% and  13-18% of total primary production and nitrate uptake, respectively. Our results suggest that diatoms, despite their minor contribution to the autotrophic community biomass ( 6%), have a disproportionally large impact on organic matter cycling in the eastern equatorial Pacific. & 2010 Elsevier Ltd. All rights reserved.

Keywords: Silicon cycle Silica production Diatoms Equatorial Pacific

1. Introduction The equatorial Pacific upwelling zone plays an important role in the global carbon cycle. It is the largest oceanic source of CO2 to

n Corresponding author. Present address: Marine Science Institute, University of California, Santa Barbara, CA 93106, USA. Tel.: + 1 805 893 7061; fax: + 1 805 893 8062. E-mail address: [email protected] (J.W. Krause). 1 Present address: Institut Universitaire Europe´en de la Mer, Technopole BrestIroise, Place Nicolas Copernic, Plouzane´ 29280, France.

0967-0645/$ - see front matter & 2010 Elsevier Ltd. All rights reserved. doi:10.1016/j.dsr2.2010.08.010

the atmosphere (Takahashi et al., 2009), and has been estimated to support 410% of global marine new production (Chavez and Barber, 1987). Surface waters in this region have high macronutrient content compared to the Pacific subtropical gyres and, despite relatively high rates of new and export production (Murray et al., 1995), chlorophyll levels are relatively low and macronutrients are not completely utilized by phytoplankton. Thus the equatorial Pacific has been characterized as one of the world’s three main high-nutrient low-chlorophyll (HNLC) systems (Minas et al., 1986) where much of the potential for both primary and export production is unrealized.

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It has been hypothesized that diatoms account for much of the new production in the equatorial Pacific (Dugdale and Wilkerson, 1998; Leynaert et al., 2001). Ku et al. (1995) specifically noted the low ratio of Si(OH)4:NO–3 ( o1) in the source waters upwelling in this region, leading to the hypothesis that the rate of Si delivery to the euphotic zone controls primary and export production by limiting diatom growth. The estimated quantitative importance of diatoms to ecological and biogeochemical processes in the equatorial Pacific depends on the nature of the comparison. Diatoms occur in relatively low numerical abundance and have been estimated to comprise 5% of phytoplankton carbon in the system (see Landry et al., 1997 and references therein). However, extrapolation of measured rates of biogenic silica production in the western equatorial Pacific to new production values, using typical diatom cellular Si:N ratios (Brzezinski, 1985), implies that diatoms may be responsible for a significant proportion of the new production (Leynaert et al., 2001), consistent with some numerical simulations (Dugdale and Wilkerson, 1998). The idea that diatoms contribute significantly to primary and export production in the equatorial Pacific is bolstered by observation. A massive aggregation of diatoms, consisting of species from the genus Rhizosolenia, was observed north of the equator at a convergence zone (Yoder et al., 1994) and fresh diatoms were observed carpeting the sea floor beneath the equatorial Pacific at 1401W latitude (Smith et al., 1996). And when no major bloom is in progress, estimated diatom growth rates in the euphotic zone are high ( 41 d  1, Latasa et al., 1997). While many previous studies in the equatorial Pacific have measured diatom pigment concentrations and numerical abundances (e.g. Kaczmarska and Fryxell, 1995; Bidigare and Ondrusek, 1996; Kobayashi and Takahashi, 2002), few studies report silicification rates or estimate the diatoms’ contribution to primary production and nitrate uptake. Currently, there are only two published data sets reporting silicification rates in this region, with most of the data from the central and western equatorial Pacific (from 1701E to 1551W, Blain et al., 1997; Leynaert et al., 2001). These results show low biogenic silica concentrations (50-100 nmol Si L  1) and low integrated silicification rates ( 1 mmol Si m  2 d  1), compared to coastal systems and the Southern Ocean (typically 41 mmol Si L  1 and 410 mmol Si m  2 d  1 for standing stock and silicification, respectively; Brzezinski et al., 2001; Shipe and Brzezinski, 2001). Considering the dynamic physical environment of the equatorial Pacific and how it changes in response to atmospheric forcing (e.g. El ˜o, ENSO), our current understanding of silicification and how it Nin varies in space and time is very limited. Furthermore, measurements are sparse in the cold tongue of the eastern equatorial Pacific where there is higher nutrient delivery via upwelling and a stronger HNLC condition, compared to the western region. In this study, silicification rates were measured in the eastern equatorial Pacific cold tongue to quantify the magnitude and variability in silica production under HNLC conditions and to evaluate the contribution of diatoms to total primary production and nitrate uptake. We report rates of biogenic silica production measured by 32Si tracer incubations in the surface waters of the equatorial Pacific between 1101W and 1401W in December, 2004 and September, 2005, well to the east of previous studies (Blain et al., 1997; Leynaert et al., 2001). Silica production rates were consistent across the 43,300 km study area and between the two cruise years, indicating that diatom growth in the equatorial Pacific is highly regulated and much less variable than other systems studied to date. Comparison of silica-production rates, with concurrent measurements of C and N uptake rates and the Si/C and Si/N ratios of the most abundant diatoms in the system, shows that the contribution of diatoms to primary production and nitrate uptake exceeded their contribution to biomass by a factor

435

of 2-3, indicating a disproportionate role for diatoms in elemental cycling across an extensive region of the eastern equatorial Pacific.

2. Methods Sampling was conducted along one meridional and one zonal transect on each of two cruises (Fig. 1). Profiles of silicic acid concentration, [Si(OH)4], and the concentration of biogenic silica, [bSiO2], were obtained in the upper 300 m at 50 stations (Fig. 1). For measurement of [Si(OH)4], 50-mL samples were drawn and analyzed colorimetrically at sea as described by Brzezinski and Nelson (1995). Nirate+ nitrite [NO–3 +NO–2] concentrations were measured at sea on a Bran and Luebee Autoanalyzer II (Dugdale et al., 2007). For measurement of [bSiO2], a 1.0 or 2.0-L sample was filtered using a 47-mm diameter, 0.6-mm pore, polycarbonate filter and analyzed via NaOH digestion, as described by Nelson et al. (2001). Vertical profiles of silicification rates were measured at 30 of the 50 stations (shaded symbols, Fig. 1A, B). Between 9 and 29 December 2004 (EB04), eight profiles were obtained on a northsouth transect from 41N to 31S on 1101W and eight profiles on an east-west transect from 1101 to 1401W on the equator (Fig. 1A). Between 8 and 24 September 2005 (EB05), eight profiles were obtained on a north-south transect from 41N to 2.51S on 1401W and five profiles sampled along a west-east transect from 13215 to 1251W on 0.51N (Fig. 1B). The zonal transect in EB05 was offset 0.51 to the north of the equator in order to sample through the greatest possible range in sea-surface temperature (SST) associated with westward-propagating tropical instability waves (TIWs). One additional station at 1.751N, 1251W sampled the cold-area of a TIW, as determined from satellite SST data. TIWs are common features in the equatorial Pacific and are generated by the shear between the westward propagating south equatorial

Fig. 1. Station locations for the EB04 (A) and EB05 (B) cruises. The meridional track was from north to south in both years, while zonal track was in opposite directions in each year (westward in EB04, eastward in EB05). Shaded symbols indicate stations where nutrients, [bSiO2], and silicification rates were measured. Open symbols indicate stations where only nutrients and [bSiO2] were measured.

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current, and countercurrents to the north and south (Philander, 1978). TIWs are physically dynamic and can impact local biogeochemistry by perturbation of the three-dimensional flow regime. For a more exhaustive discussion of the impact of TIWs on the hydrographic and chemical environment during EB04 and EB05 see Strutton et al. (2011). Seawater for Si cycling studies was collected from a tracemetal clean CTD rosette system during EB04, as described in Brzezinski et al. (2008). In EB05 samples were collected with a standard CTD rosette. During both years, sampling was done at depths where 100, 31, 13, 8, 5, 0.8, and 0.1% of the photosynthetically available radiation measured just below the sea surface, I0, was available. An additional sample was collected at 150 m and incubated in the dark during EB05 as previous studies in oligotrophic environments had shown silica production extending to depths greater than the 0.1% light depth (Nelson and Brzezinski, 1997; Brzezinski et al., 1998). Seven profiles in EB05 had triplicate subsamples per depth, whereas all other profiles (in both years) had one sample per depth. Seawater was directly subsampled into acid-cleaned  250 mL polycarbonate incubation bottles from the CTD rosette using acid-washed silicone tubing, or drained into an acid-cleaned 6-L polypropylene container when using the trace-metal clean rosette (EB04) or when performing replicate subsampling (EB05). Water from each polypropylene container was used to fill one or more polycarbonate incubation bottles. The container was mixed between subsampling to keep the particle assemblage homogenized. The rate of bSiO2 production (r, in nmol Si L  1 h  1) was measured in incubation experiments using high specific activity ( 450 kBq/mmol Si) 32Si. Prior to use, trace metals were removed from the 32Si stocks by passage through Chelex resin. The 32Si activity added to samples ranged from 330 to 830 Bq (  20,000–50,000 DPM) per sample, depending on the ambient [Si(OH)4]. After the tracer was added, all bottles were incubated on deck in acrylic incubators that maintained sea-surface temperature by continuous flow of surface seawater. Seven separate incubators were used, each screened with neutral density filters to simulate in situ irradiance at one of the sampling depths. Samples from 150 m were incubated in the 0.1% I0 incubator in bottles darkened with black electrical tape. All samples were incubated for 8-24 h (see below for details regarding incubation times). After incubation, samples were filtered under vacuum ( o25 cm Hg) onto 25-mm diameter, 0.6-mm pore, polycarbonate membrane filters, and rinsed with 0.6-mm filtered seawater to remove excess tracer not incorporated into particles. Filters were then placed in 20 mL plastic scintillation vials, loosely capped and allowed to dry. After drying, all vials were capped until analyzed for 32Si activity on land. The 32Si activity collected from the incubated samples was measured by liquid scintillation counting. Samples were aged for at least 100 days after filtration allowing the decay of any daughter isotope (32P) which was taken up or scavenged during the incubation, and to allow the 32P generated from the decay of 32Si to reach secular equilibrium with the parent isotope. Then the total combined activity of 32Si and 32P was determined by adding 2.0 mL of 2.5 N HF to each vial, dissolving the bSiO2 for 2 h, then adding scintillation cocktail (10 mL of HP Ultima Gold XR), and determining the total radioactivity using a Wallac Model 1409 scintillation counter. The calculation of r and additional specifics of the method can be found in Nelson et al. (2001). On the 1101W transect, r was measured separately in daylight and night-period hours (rDAY and rNIGHT, respectively) at each station; one vertical profile was obtained from a cast initiated at 04:00 and the other from a cast initiated at 18:00 (local times). Samples from the  04:00 cast were incubated for 8 h in daylight and those from the  18:00 cast for  8 h at night.

To calculate a daily rate of biogenic silica production (r24-H), day and night-period rates (per hour) were adjusted to a 12-hour period (i.e. 12:12 light:dark cycle) as follows:

r24-H ðnmolSi L1 d1 Þ ¼ ½rDAY ðnmol SiL1 h1 Þ  12 1 þ ½rNIGHT ðnmol SiL1 h Þ  12

ð1Þ

On the equatorial transect during EB04, and at all stations in EB05, the ship did not remain on-station long enough to permit pre-dawn and late-afternoon casts at each site. Thus rDAY and rNIGHT were measured by taking two samples at each depth from a morning cast then incubating one for  8 h in daylight and the other for  24 h. The 8-h daylight incubation provided a measure of rDAY (as on the 1101W transect) while the  24-h incubation provided a direct measure of r24-H. rNIGHT was calculated by subtracting rDAY from the r24-H while correcting for the 12 hour light and dark periods during a full day:

rNIGHT ¼ f½r24-H ðnmol i L1 h1 Þ  24 1

½rDAY ðnmolSi L1 h

Þ  12gC12

ð2Þ

During both cruises, profiles of 14C primary productivity (Balch et al., 2011) and 15N uptake (Parker et al., 2011) were obtained at the same stations where r was measured. During EB04, tracer studies with 14C, 15N and 32Si were all done together using water from the same CTD casts. In EB05, rDAY was measured at  04:00 with the 14C and 15N rate measurements. We did not measure r24-H or rNIGHT on the  04:00 cast due to constraints on water availability, but measured both rDAY and r24-H by incubating samples from casts that began at 07:00. Thus in EB05, we have two measurements of rDAY and one measurement of rNIGHT at each station. We used trapezoidally-integrated rates of r, photosynthesis (Balch et al., 2011) and N uptake (Parker et al., 2011) from the surface to 0.1% I0 (i.e. euphotic zone) to compare silica production rates with those of primary production and nitrate uptake. Silicification rates from 150 m were not used in integrals of r unless this depth was in the euphotic zone. The specific rate of biogenic silica production (Vb, in d  1) at each sampling depth was calculated as: Vb ¼ r24-H C½bSiO2 

ð3Þ

The mean Vb for the euphotic zone (VAVE) in a given profile was calculated by trapezoidally integrating Vb measurements within the euphotic zone, and dividing by the euphotic zone depth.

3. Results 3.1. Hydrography, nutrients, and biogenic silica distributions The equatorial Pacific is a region where many different water masses mix, creating significant meridional and zonal physical variability (Fiedler and Talley, 2006). We discuss physical data here to give context to our biogeochemical data. For a more thorough examination of the hydrography and nutrient structure from the EB04 and EB05 cruises see Strutton et al. (2011). Sections of current velocity (zonal and meridional components), salinity and potential density (sy) in the upper 300 m are shown in Figs. 2–5 (see panels A, B, D). During the EB04 cruise, the physical structure of the system was consistent with previous observations (Fiedler and Talley, 2006 and references therein). Along the 1101W transect, zonal

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Fig. 2. Physical, standing stock, and silicification rate data in the upper 300 m on the 1101W meridional transect (EB04 cruise). (A) Meridional (v, North¼ positive) and (B) zonal (u, East ¼positive) current velocity; velocities were measured with shipboard acoustic Doppler current profiler (ADCP, for data processing see Strutton et al., 2011). (C) Nitrate plus nitrite concentrations [NO–3 + NO–2]. [NO–3 +NO–2] data, for Figs. 2–5, from Parker et al. (2011). (D) Salinity profiles with an overlay of potential density (sy) contours, (E) silicic acid concentrations [Si(OH)4], and (F) biogenic silica concentrations [bSiO2]. (G) Daily silicification rate r24-H and the (H) specific rate of silicification Vb. Contouring and smoothing (for Figs. 2–5) was done with a built-in algorithm in the Ocean Data View software package (Schlitzer, R. 2006, http://odv.awi.de).

velocity was generally symmetric about the equator (Fig. 2B), with a sub-surface maximum marking the core of the Equatorial Undercurrent (EUC) observed within one degree of the equator. The meridional flow at stations north of 11S was predominantly southward in the upper 300 m, with minor northward flow in the core of the EUC. South of 11S, there was a two-layer character, where surface waters had southward flow but northward flow was generally dominant below  120 m (Fig. 2A). There was an overwash of low-salinity water from north of the equator (Fig. 2D) and a shoaling of isopycnals to the south of the equator (Fig. 2D) due to TIW activity (Strutton et al., 2011). The zonal transect on the equator also showed influences from TIW activity, most noticeable in meridional velocity oscillations (Fig. 3A). In the upper 100 m, meridional velocity was predominantly northward between  136-1271W and to the east of  1181W, while southward velocity was observed between  118-1271W and to the west of 1361W (Fig. 3A). Zonal velocity clearly showed the location of the EUC and its shoaling from west to east (Fig. 3B). Along the equatorial transect, salinity showed little change below 180-200 m; however, salinity showed zonal oscillations in the upper 150 m, with lower salinity water ( o35 psu) east of 1241W and higher salinity water ( 435 psu) to the west of 1241W. The salinity section suggests the advection and mixing of low salinity water (from north of the equator) and higher salinity water (from south of the equator), likely forced by TIW

activity. Isopycnals generally mirrored the west-east shoaling of the EUC; however, the 23.5 kg m  3 sy isopycnal shoaled to near the surface at  1251W. During the EB05 cruise, the hydrography was similar to that observed in EB04. On 1401W, strong eastward velocity beneath the equator marked the location of the EUC (Fig. 4B). The meridional velocity at the EUC core was southward (Fig. 4A), opposite from that associated with the EUC at 1101W. Isopycnals shoaled south of the equator, centered at 11S (Fig. 4D). Another difference, compared to the 1101W transect in EB04, was the higher salinity water in the upper 200 m at all stations south of equator, 35.2-35.6 (Fig. 4D). We sampled through the cusp of a TIW on the 0.51N zonal transect. This was indicated by strong northward meridional velocity in the upper 100 m, between  132-1281W, bounded by southward flow on either side (Fig. 5A). The 23.5 kg m  3 isopycnal broke the surface between  1241 and 1301W, and higher salinity water was observed in the upper 100 m between  1281 and 1241W, likely due to mixing by the passing TIW (Fig. 5D). The zonal velocity high, associated with the EUC core, was present between 1341 and 1251W but absent further to the east (Fig. 5B). Concentrations of macronutrients ([NO–3 +NO–2] and [Si(OH)4]) in the upper 300 m were consistent with previous data from this region (Archer et al., 1996; Wilkerson and Dugdale, 1996 and references therein). Typically we observed a range of o5-30 and

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Fig. 3. Physical, standing stock and silicification rate data in the upper 300 m on the equator zonal transect (EB04 cruise). Contoured data and layout is the same as Fig. 2.

o3-20 mM for [NO–3 +NO–2] and [Si(OH)4], respectively, along the transects (Figs. 2–5). Concentrations of silicic acid within the upper 75 m ranged from 2.5 to 4.0 mM (1st and 3rd quartiles) and, on average, varied by only 1.2-1.3 mM between the surface and  75 m at all stations. [NO–3 +NO–2] was higher than [Si(OH)4] in surface waters, consistent with previous observations that [Si(OH)4] is less than [NO–3] throughout the upper water column of the equatorial Pacific (Ku et al., 1995; Dugdale and Wilkerson, 1998; Fiedler and Talley, 2006). In EB05, nutrient isopleths did not exhibit a west to east shoaling along the 0.51N zonal section (Fig. 5C, E). The range in [bSiO2] within the euphotic zone was similar during both cruises. [bSiO2] was 75-100 nmol Si L  1 in the upper 150 m along most of the 1101W transect (Fig. 2F), and a similar range in [bSiO2] was observed on the 1401W transect within about 2 degrees of the equator (Fig. 4F). [bSiO2] was generally 4100 nmol Si L  1 in the upper 100 m along both zonal transects (Figs. 3F and 5F). The highest [bSiO2] were observed on the 0.51N transect, where the maximum surface [bSiO2] was 200 nmol Si L  1 and [bSiO2]475 nmol Si L  1 persisted to 300 m at all the transect stations (Fig. 5F). Overall, the mean and range of [bSiO2] in the euphotic zone were similar between years. The average [bSiO2] in the upper 150 m was 93 nmol Si L  1 (range 18-224) during EB04, and 99 nmol Si L  1 (range 24-268) in EB05. Additionally, the average euphotic-zone integrated [bSiO2], R bSiO2, was similar in the two years, 9.871.4 mmol Si m  2 in EB04 and 10.474.0 mmol Si m  2 in EB05 (Table 1). Closer R inspection of the data shows that while the mean bSiO2 was

nearly identical along 1101W (9.671.4 mmol Si m  2) and the equator in EB04 (9.971.4 mmol Si m  2), somewhat lower values were observed along 1401W in EB05 (7.771.2 mmol Si m  2). The R highest bSiO2 values were found on the 0.51N zonal transect (EB05), where the mean was 14.172.9 mmol Si m  2 (Table 1). The EUC was a local minimum in [bSiO2] during both cruises. This was seen most clearly in the distribution of bSiO2 along the equator (EB04, Fig. 3F) where the minimum in [bSiO2] coincided with the zonal band of strong eastward flow, denoting the location of the EUC (Fig. 3B, F). Low [bSiO2] within the EUC was apparent in the meridional transects as well (compare Fig. panels 2B with 2F and 4B with 4F). The zonal minimum in [bSiO2] shoaled from west to east along the equator, paralleling the shoaling of the high-velocity core of the EUC (Fig. 3B, F). This resulted in atypical [bSiO2] profiles along the equator, with a decrease between the surface and the depth of strong eastward flow and then an increase below (compare Figs. 2B, 3B, 4B with Figs. 2F, 3F, 4F). The resulting subsurface maxima in [bSiO2] at 150 - 300 m reached 60-90 nmol Si L  1, similar to surface values and notably higher than the o50 nmol Si L  1 concentrations in the EUC (Figs. 2F, 3F, and 4F).

3.2. Biogenic silica production Biogenic silica production rates showed a strong diurnal R pattern during both cruises. Integrated daytime rates ( rDAY)

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Fig. 4. Physical, standing stock and silicification rate data in the upper 300 m on the 1401W meridional transect (EB05 cruise). Contoured data and layout is the same as Fig. 2.

R were more than twice the integrated night-period rates ( rNIGHT) at all but two stations (Table 1). The exceptions occurred at the southernmost stations on both meridional transects; at these R stations there was dramatically higher rNIGHT, compared to the R other stations, and little change in rDAY. For both cruises, the R R mean rDAY and rNIGHT were 88.5745.6 and 39.3 729.4 mmol Si m  2 h  1, respectively (error terms are standard deviations). Both r24-H and Vb showed consistent patterns with depth throughout both cruises (Figs. 2–5). Generally, the maximum r24-H (420 nmol Si L  1 d  1) and Vb (40.2-0.3 d  1) were found in the upper 50 m, usually within the mixed layer. Both r24-H and Vb then decreased with depth to very low values near 150 m (r2 nmol Si L  1 d  1 and o0.05 d  1, respectively). Out of 30 total rate stations (both years) three were anomalous in that Vb remained elevated with depth (Figs. 2H and 4H). Two of these stations were the southernmost on each of the meridional transects (i.e. stations R R were rNIGHT 4 rDAY), the other station was located at 11N on the 1101W transect (Fig. 2H). Biogenic silica production measurements at sea are rarely replicated; this is primarily due to the high cost of the radioisotope 32 Si and the large sample volumes and labor requirements of methods employing the stable isotope 30Si. Prior to this project, only one field study reported replicate measurements of bSiO2 production rates (Brown et al., 2003). In EB05, r24-H was measured in triplicate subsamples per depth at seven stations (Fig. 6), in addition to duplicated rDAY samples at every station. Within the euphotic zone, the standard deviation averaged 16% of the mean r24-H (range

2-39%; n¼49). The 32Si activity recovered from 150 m was uniformly low, and often difficult to distinguish from background activity; thus, at 150 m the relative variance increased and the standard deviation was  74% of the mean r24-H (n¼7). We R estimated variance for an integrated rate (i.e. r24-H) in a manner that paralleled the trapezoidal integration method used to compute integrated rates. First, the standard deviation between adjacent sampling depths (e.g. z0 and z1) was computed as:

spooled:1 ¼ OfðStdevZ0 2 Þ þ ðStdevZ1 2 Þg

ð4Þ

This mean pooled standard deviation was then multiplied by the depth interval:

sinterval:1 ¼ spooled:1  ðz1 -z0 Þ

ð5Þ

Then pooling all interval standard deviations we estimated the standard deviation of the integral as:

sintegral ¼ Ofðs2interval:1 Þ þ ðs2interval:2 Þ þ   þ ðs2interval:n Þg

ð6Þ

The resulting standard deviations of integrated production R rates (i.e. r24-H, Fig. 6B) averaged  11% of the mean integral R R r24-H (range 5–16%). For rDAY measurements the mean coefficient of variation between duplicated samples was 16%, indicating slightly higher variability between different casts versus sample-to-sample variability on the same cast. Overall,

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Fig. 5. Physical, standing stock and silicification rate data in the upper 300 m on the 0.51N zonal transect (EB05 cruise). Contoured data and layout is the same as Fig. 2. Data from 1341W station is from the equator (see cruise track, Fig. 1B).

these data demonstrate that uncertainty on both a single measurements of bSiO2 production and on integrated rates is generally on the order of 20-32% (2 standard deviations). R Variation in r24-H was spatially coherent with apparent TIW activity, as outlined by Strutton et al. (2011). The low production in the northern part of the 1101W transect coincided with an overwash of low salinity subtropical waters from the north (Fig. 2D), consistent with the strong observed southward current induced by a TIW (Fig. 2A, Strutton et al., 2011). On the 1101W R transect, both r24-H and VAVE showed the lowest values at the northernmost station (gyre-like conditions). Higher values were observed in the upwelling zone (i.e. within711 from the equator) and at 31S (likely under the influence of TIW advection of upwelled water southward, Table 1). On 1401W, the lowest silicification rate was observed at the northernmost station (41N) R similar to 1101W. r24-H increased by 42 times at 2.51N, where TIW activity shoaled nutrient isopleths (Fig. 4C, Table 1). Similar to the 31S, 1101W station, the southernmost station on 1401W (2.51S) was also under the influence of strong southward velocity. R Here, both r24-H and VAVE were  2  greater than at any other 1401W-transect station. On the equatorial transect, the highest R values in both r24-H and VAVE were found at the 125.51W station, the location where TIW influence resulted in the shoaling of the 23.5 kg m  3 sy isopycnal and strong southward advection R (Fig. 3). On the 0.51N transect, r24-H and VAVE at the 129.91W and 125.61W stations were both high; at these stations the

23.5 kg m  3 sy broke the surface (Fig. 5D). Additionally, high R r24-H and VAVE were observed in the cold area of a TIW (1.751N, 1251W, Table 1). Rates of silica production were higher along the 1101W transect than at 1401W, especially at stations within 11 of the equator (i.e. in R the zone of maximum upwelling). At those stations the mean r24-H 2 1 and VAVE were 2.170.1 mmol Si m d and 0.1970.02 d  1, respectively, on 1101W and 0.870.2 mmol Si m  2 d  1 and 0.0970.02 d  1, respectively, on 1401W (Table 1). Despite variability induced by spatial differences (e.g. 1101W vs. 1401W) and TIW R activity, both r24-H and VAVE were similar in EB04 and EB05 when R averaged across all stations. The mean r24-H was 1.670.6 mmol 2 1 Si m d in EB04 and 1.370.7 mmol Si m  2 d  1 in EB05, with VAVE being 0.1570.05 d  1 in EB04 and 0.1270.05 d  1 in EB05 (Table 1).

4. Discussion 4.1. Silicic acid and biogenic silica distributions In coastal regions, with high diatom biomass, Si(OH)4 can be drawn down to very low levels such that the [bSiO2] can significantly exceed [Si(OH)4] in surface waters (e.g. Brzezinski et al., 2001; Shipe and Brzezinski, 2001). In contrast, our observations in the eastern equatorial Pacific show a Si(OH)4

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Table 1 R R R R Euphotic-zone average [Si(OH)4], integrated bSiO2 ( bSiO2), integrated bSiO2 production ( r24-H), ratio of day to night bSiO2 production ( rDAY: rNIGHT) and average R specific production rate (VAVE). #Indicates stations where triplicate r24-H samples were taken (EB05 only). + Indicates the station off the 0.51N transect but within a TIW (see Methods). Error term is the standard deviation. Station

Ave. [Si(OH)4] (mM)

R

4.01 N, 110.01 W 3.01 N, 110.01 W 2.01 N, 110.01 W 1.01 N, 110.01 W 01, 110.01 W 1.01 S, 110.01 W 2.01 S, 110.01 W 3.01 S, 110.01 W 01, 116.71 W 01, 120.001 W 01, 125.51 W 01, 128.31 W 01, 131.71 W 01, 135.21 W 01, 138.71 W 01, 140.01 W

3.74 5.89 6.10 5.82 5.87 5.98 7.13 7.45 4.73 3.55 4.06 3.39 3.17 3.62 3.21 3.02

8.51 12.22 10.32 8.74 9.34 10.42 9.93 7.70 8.25 10.47 11.93 9.23 9.22 9.96 11.73 8.11

bSiO2 (mmol Si m  2)

9.76 7 1.38

R

r24-H (mmol Si m  2 d  1)

R

R

rDAY: rNIGHT

VAVE (d  1)

0.62 1.40 1.89 1.98 2.16 2.11 0.93 1.74 1.06 1.39 2.59 1.95 1.22 1.42 2.06 0.79

1.4 1.9 1.7 1.6 1.5 1.6 2.4 0.4 1.5 3.0 1.2 2.4 5.7 2.9 3.5 8.8

0.080 0.094 0.161 0.216 0.156 0.183 0.102 0.256 0.133 0.113 0.200 0.164 0.118 0.127 0.152 0.081

EB04 mean

4.80 7 1.50

1.58 70.56

2.6 7 2.0

0.1467 0.050

#

4.01 N, 140.01 W 2.51 N, 140.01 W # 1.01 N, 140.01 W 0.51 N, 140.01 W # 01, 140.01 W 0.51 S, 140.01 W # 1.01 S, 140.01 W 2.51 S, 140.01 W # 0.51 N, 132.51 W 0.51 N, 129.91 W # 0.51 N, 127.91 W 0.51 N, 125.61 W # 0.51 N, 123.41 W + 1.751 N, 125.01 W

1.78 4.55 3.41 3.91 3.19 2.82 2.72 3.26 4.13 5.57 5.38 5.03 4.82 4.21

3.77 10.16 9.48 8.41 7.70 6.71 8.29 7.30 18.01 15.20 9.47 14.41 13.10 14.19

0.31 1.00 0.71 1.13 0.67 0.86 0.71 1.98 1.82 2.49 0.98 1.79 1.18 2.29

6.9 4.6 2.6 3.9 3.3 8.3 0.5 4.3 3.9 1.9 4.0 3.0

0.091 0.072 0.082 0.101 0.082 0.122 0.084 0.245 0.105 0.164 0.096 0.154 0.106 0.162

EB05 mean Grand mean

3.91 7 1.10 4.38 7 1.38

10.44 7 3.96 10.08 7 2.85

1.28 70.67 1.44 70.62

3.9 7 2.1 3.2 7 2.1

0.1227 0.051 0.1437 0.055

Fig. 6. Variability in r measurements from seven stations in EB05 with replication. (A) r24-H with triplicate replication at each depth; station location at 0.51 N, 132.51 W. R (B) r24-H for each profile (n¼ 7) with triplicate replication at each depth (note: integrations in (B) are to 0.1% I0). Dotted line separates profiles on 1401W and 0.51N transects. Error bars are 71 standard deviation.

pool that was always many times larger than that of bSiO2. The ratio of [Si(OH)4]:[bSiO2] between 100%I0 and 5%I0 was 32.5 714.6, similar to conditions found in the subtropical gyres (Brzezinski et al., 1998; Krause et al., 2009b). The relatively low [bSiO2] and the high [Si(OH)4]:[bSiO2] are similar to those previously reported from the equatorial Pacific (Blain et al., 1997; Leynaert et al., 2001). The westward deepening of the pycnocline in the equatorial Pacific (e.g. Kessler et al., 2006; Pennington et al., 2006) has a pronounced influence on nutrient and biomass distributions. The westward deepening of the pycnocline gradually displaces the nutricline to depths below that where Ekman pumping is effective, creating increasing oligotrophy in surface waters towards the west. This trend is seen in a comparison of our data

from between 1101W and 1401W with those of Blain et al. (1997) and Leynaert et al. (2001), who measured [Si(OH)4] and [bSiO2] further to the east between 1701E and 1551W. Blain et al. (1997) observed zonal gradients in [Si(OH)4] and [bSiO2] between 1701E and 1551W; with [Si(OH)4] in the upper 100 m increasing from o2.0 mM at 1701E to 2-3 mM at 1551W. That increase in dissolved Si was matched by increases in [bSiO2] from 10 - 30 nmol Si L  1 at 1701E to 490 nmol Si L  1 at 1551W. Leynaert et al. (2001) reported [Si(OH)4] from a meridional transect on 1801. Near the equator [Si(OH)4] was similar to our 1401W values; however, at  100 m [Si(OH)4] was only 3 mM at 1801 versus 5 mM at 1401W (Fig. 4E). Leynaert et al. (2001) reported [bSiO2] of  40-50 nmol Si L  1 in the upper 100 m at 1801 which are lower than those observed further east by us and Blain et al. (1997).

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On our cruises the mean vertically integrated [Si(OH)4] in the R euphotic zone ( Si(OH)4), for stations located within one degree of the equator, at 1101W was more than twice that for corresponding stations on 1401W (808 762 mmol Si m  2 versus 3547104 mmol Si m  2). Siliceous biomass and silica production rates were also higher along 1101W compared to 1401W. These trends of increasing [Si(OH)4] and [bSiO2] from the west to the east were not readily apparent along the equator between 1401W and 1101W in EB04 due to high variability imposed by TIW activity (Strutton et al., 2011). Our work and that of previous studies (Blain et al., 1997; Leynaert et al., 2001) indicate that zonal gradients in near surface [Si(OH)4] and [bSiO2] occur across a broad region of the equatorial Pacific; however, the magnitude of the gradient between 1101W and 1401W is of similar magnitude to the variability imposed by TIW activity. A unique aspect of the bSiO2 distribution was the distinct minimum in [bSiO2] associated with the EUC. We can only speculate about the origin of this pattern. The EUC originates in the extreme western portion of the equatorial Pacific basin and thus traverses the more oligotrophic western equatorial region before reaching our study area in the eastern cold tongue. The lower [bSiO2] in surface waters reported to the west (Blain et al., 1997) may mean that sedimentation of bSiO2 into the EUC is low in the western region, imparting lower [bSiO2] to the current by the time it reaches 1401W. Thus, the minimum in [bSiO2] within the EUC, between 1401W and 1101W, may represent the advection of low [bSiO2] waters into a region of generally higher siliceous biomass. This would mean that the relative high [bSiO2] observed beneath the EUC in our study area did not arise by simple one dimensional vertical sinking. The relatively high concentration of siliceous particles found beneath the EUC would have been transported both laterally and vertically around the core of the current. Considering the known complexity of the equatorial flow-field in that it is three dimensional and includes meridional recirculation pathways (Lukas 1986), this scenario seem reasonable but requires confirmation by more detailed measurements.

4.2. Biogenic silica production rates 4.2.1. Biogenic silica production rates at 1101W versus 1401W Comparison of the two meridional transects reveals higher silica production rates at 1101W than at 1401W. The observed spatial difference in silicification does not appear to be due to temporal variation. The final station on the equatorial transect in R EB04 was at 1401W (Fig. 1A), and the values for r24-H and VAVE at that station were similar to those observed in EB05 at the same location (1401W transect, see Table 1). A more likely explanation for the difference between the silicification rates near the equator on 1101W and 1401W is changes in nutrient supply resulting from zonal gradients in physical forcing mentioned above. Brzezinski et al. (2008) and Leynaert et al. (2001) both observed significant Si limitation of rates of silica production in the equatorial Pacific, such that greater supply rates of Si(OH)4 would drive higher rates of silica production. The doubling of the Si(OH)4 inventory within the euphotic zone between 1401W and 1101W was paralleled by a R 2.6 fold increase in mean r24-H and a 2.0 fold increase in VAVE within the zone of maximum equatorial upwelling (between 11N and 11S). Considering these differences, and the west-east shoaling of the nutricline, one would expect that silicification rates would increase progressively eastward along the equator. But similar to the lack of zonal gradients in [Si(OH)4] and [bSiO2] in the euphotic zone along the equator, no such gradient was observed in silicification rates (Fig. 3G, H, Table 1). This may have

been due to forcing by TIW activity being the dominant factor influencing variability along the equator rather than simple one dimensional vertical upwelling, such that zonal gradients are obscured without sampling multiple stations at a given longitude as was done at both 1101 and 1401W.

4.2.2. Tropical instability waves and biogenic silica production While TIWs appear as waves in satellite images of SST and chlorophyll, they are westward-propagating trains of tropical instability vortices offset from the equator (Strutton et al., 2011). Their passage induces an oscillatory meridional flow along the equator which displaces surface waters to the north and then to the south as the vortex passes, creating wave-like patterns in chlorophyll distributions along the equator. However, given the lack of detailed 3-dimensional circulation data, we cannot constrain the exact position of these vortices. Instead, we relate processes to TIW crests (e.g. regions of northward transport of upwelled water from the equator) and troughs (e.g. regions of localized southward transport) as discussed in Strutton et al. (2011). We see evidence of both the influence of the wave-like characteristics of TIWs along the equatorial zonal section in EB04 and possible influences of localized vertical motion along the meridional and 0.51N zonal sections. On both meridional transects, TIW activity induced southward flow of lower-nutrient and less-saline water from the north, R resulting in lower availability of Si(OH)4 and NO3 . r24-H at the northern station on each transect was the lowest observed for each respective cruise (Figs. 2 and 4). At 1401W, both the winddriven upwelling and the amplitude of TIWs were weaker than at 1101W. On the 1401W transect, the western edges of two TIW were passed near 21N and 21S, with the northern feature being more intense (Strutton et al., 2011). Alteration of the water column by TIW activity was evident as we did not observe a R maximum in r24-H near the equator. Instead we observed R enhancements in r24-H at 21N and 2.51S; these locations were closest to the locations where the our transect intersected the TIW crests (Strutton et al., 2011). During the 1101W transect the TIW advection shifted the equatorial upwelling southward, with the maximum centered on 11S (see sy in Fig. 2D). No discernable R difference in r24-H or VAVE was observed at stations between 11S and 11N. This could indicate that upwelling enhancements, or north-to-south shifting of the upwelling maximum, by TIWs do not induce bSiO2 production enhancements at this longitude. The influence of three-dimensional flows within a TIW was observed to be greatest on the 0.51N transect in EB05. Strong northward velocity was associated with the uplift of isopycnals such that the 23.5 kg m  3 sy surface broke the surface, indicating R localized upwelling, between 1271W and  1321W. r24-H reached a maximum value in this same area and was accomR panied by the highest bSiO2 observed on any transect. In contrast, production was low in the regions of southward advection at 1341W and 1251W. At 1.751N, 1251W station, determined by satellite SST to have the coolest surface waters R within the TIW, high rates of r24-H and VAVE were also observed R along with high bSiO2 (Table 1). Displacement of the EUC by TIW activity added further complexity to the observed spatial patterns in silica production. The elevated silica production rates at 125.61W may have been influenced by strong eastward zonal velocity (Fig. 5B). Within the region of strong southward advection, there was a zonal velocity maximum associated with the EUC at the 125.61W station (high silica production) but not at the 123.41W station (low silica production); this suggests that bSiO2 production may be influenced not only by TIW perturbations to the meridional flow field

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but also by changes in zonal flow by north-south displacement of the EUC. The influences of the wave-like properties of the TIWs were apparent by the spatial variability in silica production along the equator in EB04. During the EB04 zonal transect, portions of at least two TIWs were traversed (Fig. 3A). Just prior to our occupation of the zonal transect the zonal band of high chlorophyll associated with equatorial upwelling had become displaced to the north and south of the equator by TIW activity, forming a wave-like pattern in images of ocean color (Strutton et al., 2011). During the equatorial section, the westward propagation of TIWs caused both the northern and southern peaks in the wave-like distribution of high chlorophyll to advect back towards the equator. Stations where we measured high rates of silica production coincided with positions and times when we intersected the areas of high chlorophyll. This occurred at the 125.51W and again at 138.71W. At each of these sites, high chlorophyll that was formerly displaced to the north was advected southward, and intersected the equator coincident with our sampling (Strutton et al., 2011). Lows in production were associated with northerly flow and coincided with stations where chlorophyll that had been displaced to the south was returning to, but had not reached, the equator when we sampled these areas.

4.3. Limitation of silica production by light and Si(OH)4 Unlike other open-ocean systems where silicification has been studied, silica production rates in the equatorial Pacific show a R strong diurnal rhythm. We observed that rDAY was generally 2-3 R times rNIGHT (Table 1), similar to the observations at 1801 along the equator by Leynaert et al. (2001). Other studies in open-ocean regions (e.g. Gulf Stream warm-core rings, Brzezinski and Nelson, 1989; Sargasso Sea, Nelson and Brzezinski, 1997) report day and night rates of biogenic silica production that are statistically indistinguishable. Thus, biogenic silica production in the equatorial Pacific is more tightly coupled to light availability than it is in other open-ocean systems. During both years, r decreased with depth in both the day and night profiles and at some point deep in the euphotic zone the day and night rates became low and approximately equal (Fig. 7A, B). When normalizing rDAY and rNIGHT to r24-H (expressing all three rates in nmol Si L  1 h  1), the ratio ((rDAY – rNIGHT)/r24-H) was essentially constant in the euphotic zone above the 5% I0 isolume (Fig. 7C). Over that depth range rDAY was consistently 2-3 times rNIGHT even though both rDAY and rNIGHT decreased with depth. Below the 5% I0 isolume rDAY and rNIGHT were either equal or rDAY was less than rNIGHT (Fig. 7C, see reference levels). The decrease in both rDAY and rNIGHT with depth in the upper euphotic zone does not appear to be driven by differences in siliceous biomass with depth, as [bSiO2] did not decrease dramatically over that depth interval. Rather, the pattern is driven mainly by the decline in specific rates, as the trends in Vb for vertical profiles and the normalized change in Vb with light depth (i.e. (VDAY – VNIGHT)/V24-H) are nearly identical to those for the corresponding r values (data not shown). We can also rule out responses of r to changes in the concentration of macronutrients such as Si(OH)4, NO–3 or PO34  in this pattern. The concentrations of these nutrients, while not changing dramatically with depth within the upper euphotic zone, increased with depth; which would tend to increase production rates. Thus, both r and Vb responded to changes in the light field independently of significant changes in [Si(OH)4] and [bSiO2]. Light also modulated the response of silica production to changes in Si(OH)4 availability. We observed statistically

Fig. 7. The coupling of r with light depth. The average rDAY (open circle) and rNIGHT (filled circle) with depth during EB04 (A) and EB05 (B), error bars are7 the standard error. (C) rNIGHT was subtracted from rDAY and the difference was normalized to r24-H (i.e. the hourly r24-H rate is the average of the hourly rDAY and rNIGHT rates). The normalized difference was plotted against light depth for EB04 (filled square) and EB05 (open square); error bars are7 the standard error. Reference lines are provided to show where rDAY and rNIGHT are equal (solid line), rDAY is 2 times greater than rNIGHT (dotted line), and rDAY is 3 times greater than rNIGHT (dashed line). The ‘‘No Light’’ depth is any depth deeper than the 0.1% isolume.

significant increases in Vb with increasing [Si(OH)4] in the euphotic zone, but only when data were grouped according to light levels (Fig. 8, compare panel a with panels b-e). This pattern occurred when examining both day-period (Fig. 8B–E) and 24-hr (data not shown) incubations. The slope of the regression of Vb on [Si(OH)4] was greatest at the surface, and declined with decreasing light level (Fig. 8F). This light dependence caused the overall relationship between Vb and [Si(OH)4] for all stations and depths pooled across light levels to be poor (Fig. 8A). However, when euphoticzone average [Si(OH)4] and VAVE are compared between stations, where both quantities average through the vertical light gradient, there was a statistically significant correlation between average Si(OH)4 and VAVE (r¼0.39, two-tailed test, po0.05; Table 1, see also Fig. 8B–E). These results are consistent with the direct experimental observation that Vb was limited by the ambient [Si(OH)4] on these cruises at the  50% I0 isolume (Brzezinski et al., 2008). But they reveal that the strength of substrate limitation of silica production was strongly light dependent, decreasing with depth as the ambient irradiance diminished. Our results, together with direct observations of limitation of silica production by ambient [Si(OH)4] in the equatorial Pacific (Leynaert et al., 2001; Brzezinski et al., 2008), suggest that [Si(OH)4] and light intensity have separate but interactive effects on silica production rates in the EEP. Si uptake kinetics show that

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Fig. 8. Dependence of specific biogenic silica production rates (Vb, h  1) during the daytime incubations on ambient [Si(OH)4] (mM) for the entire dataset (A). Data from (A) was broken into light depths as follows: 100% I0 (B), 31% I0 (C), 13% I0 (D), 8% I0 (E). The lines in panels (B) through (E) are the least squares regression to each light-depth subset; the reported p-value is probability that the slope of the regression is not significantly different from zero. The upper four light depths are the only depths where there is a statistically significant regression; at all depths below 8% I0 the slope of the regression is not significantly different from zero. (F) A plot of the Vb-verses-[Si(OH)4] slopes (slopes from panels (B) through (E) and including depths below 8%I0) with light depth, error bars are the standard error of the slopes from the regression analysis.

silica production, at a fixed light level, increases somewhat as ambient [Si(OH)4] increases (Leynaert et al., 2001; Brzezinski et al., 2008). A direct response of r to light alone is suggested by the decline in r with depth in the mixed layer where silicic acid concentrations were nearly uniform, but light was decreasing exponentially. An interaction between the effects of light and [Si(OH)4] is revealed by strong evidence that the sensitivity of Vb to [Si(OH)4] is a direct function of light (Fig. 8F). Together these data suggest that silica production in the EEP responds to both light and [Si(OH)4], with the result that the substrate dependence of Vb essentially disappears at low light. These findings are reminiscent of the two-layered model of oceanic euphotic zones (Dugdale, 1967), where nutrient limitation of phytoplankton rate processes is strong in well lit surface waters and light limitation ensues at greater depths. The physiological mechanisms causing light to have such a strong effect on silica production in the EEP are not clear. The diurnal pattern in silica production could arise through the effects of the daily photocycle on diatom cell division. Azam and Chisholm (1976) demonstrated that photocycles affect both cellular energy supply and diurnal patterns of cell division, which drive parallel patterns in r and Vb because of the close ties between silicon deposition and the diatom cell cycle (Chisholm et al., 1978). However, entrainment of diatom division rates to photocycles often results in multiple preferred periods of cell division which are not all timed to the light period (Chisholm et al., 1980). And sufficient interspecific variability exist in such responses (Chisholm et al., 1978; Nelson and Brand, 1979) that it would be difficult for this mechanism to yield a strong and persistent diurnal pattern in silica production over a significant fraction of the euphotic zone (Fig. 7). Another, perhaps more likely, explanation is the interaction between light and dissolved iron shown in culture studies (Sunda and Huntsman, 1997). The importance of Fe limitation in the equatorial Pacific is well known (e.g. Coale et al., 1996a, b). Irondepleted diatoms have compromised photosystems and an increase in light would therefore allow for increased primary production, growth, and cell division, thereby increasing bSiO2

production. If this hypothesis is correct then the relationship between bSiO2 production and light should also be strong in other HNLC regions; however, none of the studies that have examined silica production in other HNLC areas have examined diurnal variations in silica production (Franck et al., 2000; Brzezinski et al., 2001; Mosseri et al., 2008). Regardless of the cause, it is clear that we do not yet adequately understand the underlying mechanisms driving the observed coupling among biogenic silica production, ambient light, and depth within the euphotic zone.

4.4. Comparing biogenic silica production in the equatorial Pacific with that in other systems The equatorial Pacific differs from other open-ocean systems in that biogenic silica production decreases to very low rates at the base of the euphotic zone (Fig. 7). In other open-ocean areas such as the Sargasso Sea (Brzezinski and Nelson, 1995; Nelson and Brzezinski, 1997, Krause et al., 2009a) and the North Pacific gyre (Brzezinski et al., 1998), silicification rates between 100 and 200 m are often similar to rates in the upper 100 m. In a somewhat unusual case, Brzezinski and Nelson (1989) found that positive net production of bSiO2 (i.e. bSiO2 production 4bSiO2 dissolution) extended to a depth of 300 m in a Gulf Stream Warm Core Ring during a period of deep convective mixing. Silicification rates declined rapidly below 100 m at 90% of the stations in our study, which was typically within 10 m of the 0.1% light level (Figs. 2–5 and 7). At stations where the 150 m sampling depth was deeper than the 0.1% light level, we observed that 91 75% of integrated biogenic silica production was confined to the euphotic zone (i.e. surface to 0.1% I0), suggesting that the vertical extent of silicification is strongly tied to the depth of the euphotic zone in this region. These patterns indicate that euphotic-zone profiles capture nearly the full vertical extent of silica production. This situation is not commonly observed in other open-ocean systems as bSiO2 production rates at  160 m (  1.6  deeper than euphotic zone) can be similar to surface rates.

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Silica production rates in the equatorial Pacific are similar to, but less variable than, rates observed in subtropical gyres. R Brzezinski et al. (1998) reported a mean r24-H in the North Pacific (25-301N) of 1.270.7 mmol Si m  2 d  1 (range 0.5-2.9) which is not statistically different from the mean rate that we observed in the equatorial Pacific. However, Brzezinski et al. (1998) observed an average Vb of 0.2470.15 d  1, 450% higher than the overall mean VAVE we observed (Table 1). Moreover, at stations with diatom blooms in the North Pacific, Vb was 41.0 d  1 at multiple depths (Brzezinski et al., 1998) whereas the single highest Vb observed on either of our cruises was 0.57 d  1 (see Figs. 2–5). During a multi-year study in the Sargasso Sea, Nelson and Brzezinski (1997) reported a much lower mean R r24-H of 0.4 mmol m  2 d  1 (range 0.1-0.9) in the upper 160 m, and a mean Vb of 0.1570.15 d  1. Recently it has been reported R that r24-H in the Sargasso Sea near the Bermuda can increase to 41 mmol Si m  2 d  1 with Vb 40.5 d  1 during winter-storm and subsequent stratification events (Krause et al., 2009a). While biogenic silica production rates in the equatorial Pacific are similar to those in oligotrophic regions, they pale in comparison to rates observed in the Southern Ocean and in coastal systems. In a review of global silicification and dissolution, Nelson et al. (1995) R reported a mean r24-H of 20 mmol Si m  2 d  1 in the Southern Ocean in spring and summer, while Brzezinski et al. (2001) reported mean values of 28 and 23 mmol Si m  2 d  1 for a developed diatom bloom in the Southern Ocean during December 1997 and January 1998, respectively. Thus mean rates of biogenic silica production in the eastern equatorial Pacific are only  5% of those in the Southern Ocean in summer. The ranges of biogenic silica production rates are higher and more dynamic in coastal systems. For example, in the Santa Barbara Basin (USA) Shipe and Brzezinski (2001) reported a mean of 17 mmol Si m  2 d  1, based on a  2-year time series (range of o5-60 mmol Si m  2 d  1). Brzezinski et al. (2003), presenting data from Monterey Bay (USA, a site of seasonal R upwelling), reported a mean r24-H of 43 mmol Si m  2 d  1 during active upwelling periods. Some biogeochemical processes in the equatorial Pacific upwelling region have been described as being at or near steady state (see Frost and Franzen, 1992; Dugdale and Wilkerson, 1998); however, it has not been clear whether this is true for silicification. During our cruises there was noticeable spatial R variability in both r24-H and VAVE but the range was narrower than in other open-ocean (Nelson and Brzezinski, 1997; Brzezinski et al., 1998; Brzezinski et al., 2001) and coastal (Shipe and Brzezinski, 2001; Brzezinski et al., 2003) systems. This was true despite a very large sampling area (  2.6  106 km2) in the study reported here. Using a time-series of remote sensing data from the equatorial Pacific, Strutton et al. (2011) determined that two periods of elevated chlorophyll are generally observed during each year in this region. The EB04 and EB05 cruises were quite different from one another in terms of regional chlorophyll anomalies; EB04 had negative chlorophyll anomalies, while EB05 was strongly positive and during the secondary chlorophyll peak for the year. Strutton et al. (2011) determined this inter-cruise difference to be a seasonal effect. In addition, TIW activity is highest during the fall and winter seasons, maximizing the potential for variability in biological rates. Despite these R factors, the average r24-H and VAVE each varied by only 20% between years (Table 1). Thus, our data suggest that biogenic silica production rates in the equatorial Pacific are much less variable than in other systems. As discussed above, silica production did respond to TIW influence so the system is clearly not in steady state, but responses to perturbations are muted, especially compared to other systems, possibly due to geographically extensive Fe limitation (Coale et al, 1996a, b; Brzezinski et al., 2011) and/or high grazing pressure (Landry et al., 2011).

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4.5. Contribution of diatoms to primary production and nitrate uptake Our data on silica production, combined with size-fractionated rates of photosynthetic carbon uptake, nitrogen utilization, and cell-specific measurements of diatom stoichiometry obtained by others on these same cruises, allows us to test whether diatoms supported most of the nitrate uptake (e.g. Dugdale and Wilkerson, 1998; Leynaert et al., 2001) during the EB04 and EB05 cruises. In the eastern equatorial Pacific during our cruises, diatoms accounted for  6 % of the total phytoplankton carbon and were vastly outnumbered by prokaryotes and picoeukaryotes (Taylor et al., 2011), so their dominance of new production would indicate a disproportionately important role in carbon export. Previous studies inferring high diatom new production in the equatorial Pacific have extrapolated the contribution of diatoms to organic matter cycling using spatially and temporally inconsistent data sets. For instance, Leynaert et al. (2001) converted bSiO2 production measurements to diatom new production using Si:C and Si:N ratios derived from cultures (from Brzezinski, 1985; Takeda, 1998), and used previously published rates of primary and new production in the same vicinity. While diatoms under nutrient-replete conditions generally have a mean Si:C of 0.13 and Si:N of 1.1 (Brzezinski, 1985), these ratios increase (by greater than two fold for Si:C) in diatoms grown under low dissolved iron [Fe] conditions (Takeda, 1998). Other studies have estimated diatom new production by nutrient drawdown ratios (Dugdale and Wilkerson, 1998; Dunne et al., 1999), which is a more temporally integrated measurement than that used by Leynaert et al. (2001). During the EB04 cruise, there was indirect evidence that diatoms and larger-sized phytoplankton were carrying out much of the nitrate uptake. Dugdale et al. (2007), using data from the 1101W and equatorial transects, reported size-fractioned rates of nitrate and ammonium uptake in daylight hours. Along the 1101W transect, the near-surface f-ratio (rNO3/(rNO3 + rNH4)) for the 45 mm size-fraction ranged from 0.4 to 0.6 (except for 41N, 1101W in the region of overwash, which had very low, gyrelike [NO–3]), while f-ratio for the o5 mm size-fraction ranged from  0.1 to 0.3. Near-surface f-ratios on the equator (from 116.71W to 1401W) ranged from 0.5 to 0.8 for the 45 mm size-fraction, and  0.1 to 0.4 for the o5 mm fraction. These results demonstrate that 45 mm cells (e.g. diatoms, dinoflagellates, etc.) proportionally use more NO–3 for production than do o5 mm cells (e.g. pico-eukaryotes and cyanobacteria). To estimate the contribution of diatoms to primary production and nitrate uptake, we extrapolate the fraction of N and C uptake done by diatoms from the profiles of silica production using direct measures of the elemental composition of diatoms measured on our cruises. The larger data set collected by project members during the EB04 and EB05 cruises allow for direct comparison of R vertically-integrated daily-uptake rates for C, N, and Si (i.e. r24-H). The reported bSiO2 production rates in this study were measured in 32Si incubations run concurrently with 14C (Balch et al., 2011) and 15N (Parker et al., 2011) uptake experiments. Previous studies in the equatorial Pacific have demonstrated that the resident diatom community is dominated by small pennate genera (Blain et al., 1997; Kobayashi and Takahashi, 2002). The elemental composition of the dominant pennate diatoms in the system during our study was measured directly for individual cells by a synchrotron-based x-ray fluorescence (SXRF) microprobe analysis (Twining et al., 2004). Using SXRF, the EB04 cruise average Si:C and Si:N mole ratios for the dominant pennate diatoms present were 0.13 and 0.85, respectively, and 0.14 (Si:C) and 0.95 (Si:N) in EB05 (S. Baines, pers com). These in situ diatom Si:C and Si:N ratios are similar to the mean Si:C and Si:N ratios reported for

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Table 2 Estimated diatom contribution (% of total) to euphotic zone integrated primary production (PP) and nitrogen uptake (rN). The total PP mean from each cruise are from R R R R R Balch et al. (2011). The rN (i.e. rNH4+ + rNO–3), total rNO–3, and 45 mm size-fractioned rNO–3 means from each cruise are from Parker et al. (2011). r24-H was R converted to C and N by Si:C ¼0.15, and Si:N¼ 1.01 in EB04, and Si:C¼ 0.14, and Si:N ¼0.95 in EB05 (S. Baines, pers com). The calculated contribution to rNO–3 assumes R that diatoms use NO–3 as their only N source, and the contribution to rNO–3 in parentheses applies the mean 4 5 mm size-fractioned f-ratio of 0.34 to diatom N uptake (note: the 45 mm f-ratio cruise mean was the same for both EB04 and EB05, see Parker et al., 2011). Error term is the standard deviation. Euphotic-zone integrated rate

EB04

Diatom contribution

EB05

Diatom contribution

R r (mmol Si m  2 d  1) R 24-H PP (mmol C m  2 d  1) R rN (mmol N m  2 d  1) R rNO–3 (mmol NO–3 m  2 d  1) R rN, 45 mm (mmol N m  2 d  1) R rNO–3, 45 mm (mmol N m  2 d  1)

1.58 7 0.56 56.93 7 16.69 10.19 7 5.35 2.89 7 1.44 3.62 7 1.18 1.22 7 0.72

100% 18.5% 15.4% 54.1% (18.4%) 43.2% 128.2% (43.6%)

1.28 7 0.67 52.08 7 19.37 19.75 7 5.36 3.44 7 1.84 8.27 7 3.41 2.39 7 1.43

100% 17.6% 6.8% 39.2% (12.5%) 16.3% 56.4% (18.0%)

nutrient replete diatoms in culture (Si:CE0.13, Si:NE1.1, Brzezinski, 1985). The 32Si, 14C, 15N and SXRF data enable us to make a novel estimate of the diatoms’ contribution to primary production and nitrate uptake, in that it is constrained by direct contemporaneous measurement of the relevant variables. While we performed daylight, night, and 24-h 32Si tracer incubations to measure bSiO2 production rates, primary production and nitrogen uptake experiments focused on the daylight period only. Balch et al. (2011) calculated vertically integrated daily rates of R primary production ( PP) during dawn to dusk incubations. Parker et al. (2011) derived daily 15N-based rates of NO–3 and NH4+ uptake by R R multiplying daytime hourly rates for rNO–3 and rNH4+ by 12 and 18, respectively (McCarthy et al., 1996). We estimate the percentage contribution of diatoms to primary production (Diatom % PP), total nitrogen uptake (Diatom % rN) and nitrate uptake (Diatom % rNO–3) as follows:  Z  Z Diatom% PP ¼ 100  r24-H C ðSi : CÞ  PP ð7Þ

Diatom% rN ¼ 100 

Z



r24-H C ðSi : NÞ 

Z

rNO3  þ

Z

rNH4 þ



ð8Þ Diatom% rNO3- ¼ 100 

Z



r24-H C ðSi : NÞ 

Z

rNO3 

 ð9Þ

Consistent with previous studies, these calculations suggest that diatoms accounted for a disproportionately high amount of primary production and nitrate uptake (Table 2). Diatoms accounted for  18% of the mean total PP in both years (Table 2). Landry et al. (2011), using independent dilution-experiment methodology, also estimated that diatoms contribution to total primary production was 18% for both cruise years. R R The diatom contribution to total nitrogen (i.e. rNO–3 + rNH4+ ) uptake was similar to their contribution to total PP in EB04 (mean 15%, Table 2) but lower than their contribution to PP in EB05 (mean 7%, Table 2). Eq. (9) derives diatom rNO–3 by assuming that NO–3 is their sole source of N (e.g. Dugdale and Wilkerson, 1998), thus it is a maximum estimate of the diatoms’ contribution to rNO–3. By this maximum estimate, diatoms were responsible for 39-54% of the total nitrate uptake in both years (Table 2). During both EB04 and EB05, size-fractionated (total and 45 mm) NO–3 uptake samples were taken throughout the entire euphotic zone (Parker et al., 2011). Again, assuming the diatoms’ sole source of N was NO–3 indicates they potentially carried out  100% and 56% of the 45 mm sizefractionated rNO–3 in the euphotic zone on the EB04 and EB05 R cruises, respectively (Table 2, note: the 45 mm rNO–3in EB04 was based on three-depth profiles, opposed to eight-depth profiles in EB05, hence the estimate is less constrained, see Parker et al., 2011). Parker et al.’s (2011) assessment of diatom N demand versus 45 mm

rNO–3 indicated that at only a few stations during the EB cruises could diatom N demand have been met completely by NO–3. Parker et al. R (2011) report an f-ratio of 0.34, for the 45 mm size-fractioned rN, + in both years. Applying this f-ratio to diatoms accounts for NH4 -based N uptake and significantly decreases the estimated proportion of nitrate uptake attributable to diatoms. In both years they were responsible for 13-18% of total nitrate use and 18-44% of the nitrate taken up by the 45 mm size-fraction (Table 2). These lowered estimates are consistent the findings of Parker et al. (2011), which indicate that during the EB cruises autotrophic dinoflagellates, in addition to diatoms, were likely strong contributors to nitrate uptake. Our f-ratio-derived estimates for diatom nitrate uptake also differ from previous reports suggesting that diatoms alone carry out most of the total community nitrate uptake in the equatorial Pacific (Dugdale and Wilkerson, 1998; Leynaert et al. 2001). Significant regeneration of NO–3 has been reported within the upper 120 m of the equatorial Pacific upwelling zone, which can fuel between 20 and 100% of nitrate uptake (Raimbault et al., 1999). Demarest et al. (2011) have presented an analysis suggesting that this was also the case during the EB05 cruise. If so, then measured rates of total rNO–3 may significantly overestimate rates of new production, as nitrate is not an entirely ‘new’ form of N in the system. For this reason, the estimates of diatoms’ contributions to nitrate uptake (Table 2) likely overestimate their gross amount of new production. These results suggest that the diatoms’ contribution to primary production and nitrate uptake in the equatorial Pacific is multiple times higher than their contribution to autotrophic biomass ( 6%; Taylor et al., 2011). This situation is similar to that estimated by Nelson and Brzezinski (1997) in the Sargasso Sea. Diatoms in the Sargasso Sea are similarly a very minor component of the autotrophic community in terms of biomass (typically o5% see Goericke, 1998; Steinberg et al., 2001). But the diatom contribution to annual organic matter production, estimated by methods similar to those used here (but without SXRF-derived diatom stoichiometry) is 20% on an annual basis (Nelson and Brzezinski, 1997). Their contribution to organic matter export, based on organic carbon, nitrogen, and bSiO2 collection over periods of 2-4 days by sediment traps at 150, 200 and 300 m, was estimated to be  30% on an annual basis, increasing to 490% during the spring bloom period. Our data, coupled with those of Nelson and Brzezinski (1997), indicate that diatoms have a disproportional impact on biogeochemical processes in these open-ocean systems.

5. Conclusion During cruises in the eastern equatorial Pacific in December 2004 and September 2005 we measured biogenic silica production rates on four transects. We observed significantly higher biogenic silica production rates at stations within 11 of the

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equator at 1101W than at 1401W, likely driven by greater Si(OH)4 availability. On zonal transects along the equator and 0.51N, we found a spatial coherence between bSiO2 production rates and TIWs, resulting in enhanced silica production from increased local upwelling (0.51N transect in EB05) or the north-to-south transport of the surface-chlorophyll maximum (equator transect in EB04). Over the broad sampling area variance in bSiO2 production rates was within a narrow range compared to other systems, despite clear responses to upwelling at the equator and in association with TIW activity. This suggests variability in bSiO2 production throughout much of the equatorial Pacific is dampened, possibly due to grazing regulation or iron limitation. Light and [Si(OH)4] affected silicification rates independently and interactively, whereby the response to increased Si(OH)4 availability was a linear function of relative light intensity. This is the only open-ocean region studied to date where such a strong dependence of silicification on the ambient light field has been observed. Such an effect may be the result of Fe-light colimitation, as previously shown in culture studies. In this system diatoms are estimated to account for 18% of the total primary production and 13-18% of the total nitrate uptake in this system. These estimates all indicate a disproportionately large role for diatoms in C and N cycling in the surface waters of this region.

Acknowledgments We thank S. Baines, W. Balch, R. Dugdale, A. Parker, B. Twining, and F. Wilkerson for permission to use their data. T. Cowles and P. Strutton provided insightful comments. A. Palacz processed and distributed the ADCP data. J. Arrington, C. Beucher, M. Demarest, C. Dumousseaud, and J. Jones provided valuable shipboard and laboratory assistance. Thanks to the science party and crew on each cruise aboard the R.V. Roger Revelle for outstanding technical and logistical support, specifically technician G. Pillard. This work was funded by the National Science Foundation Biocomplexity Grant OCE 03-22074 awarded to DMN, additional support awarded to JWK was provided by the Oregon Space Grant and Leynore Bayley Fellowships at Oregon State University. References Archer, D.E., Takahashi, T., Sutherland, S., Goddard, J., Chipman, D., Rodgers, K., Ogura, H., 1996. Daily, seasonal and interannual variability of sea-surface carbon and nutrient concentration in the equatorial Pacific Ocean. Deep-Sea Research II 43, 779–808. Azam, F., Chisholm, S.W., 1976. Silicic acid uptake and incorporation by natural marine phytoplankton populations. Limnology and Oceanography 21, 427–435. Balch, W.M., Poulton, A.J., Drapeau, D.T., Bowler, B.C., Windecker, L.A., Booth, E.S., 2011. Zonal and meridional patterns of phytoplankton biomass and carbon fixation in the equatorial Pacific Ocean between 1101W and 1401W. Deep-Sea Research II 58 (3–4), 400–416. Bidigare, R.R., Ondrusek, M.E., 1996. Spatial and temporal variability of phytoplankton pigment distributions in the central equatorial Pacific Ocean. Deep-Sea Research II 43, 809–833. Blain, S., Leynaert, A., Treguer, P., Chretiennot-Dinet, M.J., Rodier, M., 1997. Biomass, growth rates and limitation of Equatorial Pacific diatoms. Deep-Sea Research I 44, 1255–1275. Brown, L., Sanders, R., Savidge, G., Lucas, C.H., 2003. The uptake of silica during the spring bloom in the Northeast Atlantic Ocean. Limnology and Oceanography 48, 1831–1845. Brzezinski, M.A., 1985. The Si:C:N ratio of marine diatoms: interspecific variability and the effect of some environmental variables. Journal of Phycology 21, 347–357. Brzezinski, M.A., Nelson, D.M., 1989. Seasonal changes in the silicon cycle within a Gulf Stream warm-core ring. Deep-Sea Research 36, 1009–1030. Brzezinski, M.A., Nelson, D.M., 1995. The annual silica cycle in the Sargasso Sea near Bermuda. Deep-Sea Research I 42, 1215–1237. Brzezinski, M.A., Villareal, T.A., Lipschultz, F., 1998. Silica production and the contribution of diatoms to new and primary production in the central North Pacific. Marine Ecology-Progress Series 167, 89–104.

447

Brzezinski, M.A., Nelson, D.M., Franck, V.M., Sigmon, D.E., 2001. Silicon dynamics within an intense open-ocean diatom bloom in the Pacific sector of the Southern Ocean. Deep-Sea Research II 48, 3997–4018. Brzezinski, M.A., Jones, J.L., Bidle, K.D., Azam, F., 2003. The balance between silica production and silica dissolution in the sea: Insights from Monterey Bay, California, applied to the global data set. Limnology and Oceanography 48, 1846–1854. Brzezinski, M.A., Dumousseaud, C., Krause, J.W., Measures, C.I., Nelson, D.M., 2008. Iron and silicic acid concentrations together regulate Si uptake in the equatorial Pacific Ocean. Limnology and Oceanography 53, 875–889. Brzezinski, M.A., Baines, S., Balch, W.M., Beucher, C., Chai, F., Dugdale, R.C., Krause, J.W., Landry, M.R., Marchi, A., Measures, C.I., Nelson, D.M., Parker, A., Poulton, A., Selph, K.E., Strutton, P., Taylor, A.G., Twining, B.S., 2011. Co-limitation of diatoms by iron and silicic acid in the equatorial Pacific. Deep-Sea Research II 58 (3–4), 493–511. Chavez, F.P, Barber, R.T, 1987. An estimate of new production in the equatorial Pacific. Deep-Sea Research 34, 1229–1243. Chisholm, S.W., Azam, F., Eppley, R.W., 1978. Silicic acid incorporation in marine diatoms on light: dark cycles: use as an assay for phased cell division. Limnology and Oceanography 23, 518–529. Chisholm, S.W., Morel, F.M.M., W.S. Slocum, W.S., 1980. The phasing and distribution of cell division cycles in marine diatoms. In: Falkowski, P. (Ed.), Primary Productivity in the Sea.. Plenum Press, New York, pp. 281–300. Coale, K.H., Fitzwater, S.E., Gordon, R.M., Johnson, K.S., Barber, R.T., 1996a. Control of community growth and export production by upwelled iron in the equatorial Pacific Ocean. Nature 379, 621–624. Coale, K.H., Johnson, K.S., Fitzwater, S.E., Gordon, R.M., Tanner, S., Chavez, F.P., Ferioli, L., Sakamoto, C., Rogers, P., Millero, F., Steinberg, P., Nightingale, P., Cooper, D., Cochlan, W.P., Landry, M.R., Constantinou, J., Rollwagen, G., Trasvina, A., Kudela, R., 1996b. A massive phytoplankton bloom induced by an ecosystems-scale fertilization experiment in the equatorial Pacific Ocean. Nature 383, 495–501. Demarest, M.S., Brzezinski, M.A., Nelson, D.M., Krause, J.W., Jones, J.L., Beucher C., 2011. Net biogenic silica production and nitrate regeneration determine the strength of the silica pump in the eastern equatorial Pacific. Deep-Sea Research II 58 (3–4), 462–476. Dugdale, R.C., 1967. Nutrient limitation in the sea: dynamics, identification and significance. Limnology and Oceanography 12, 685–695. Dugdale, R.C., Wilkerson, F.P., 1998. Silicate regulation of new production in the equatorial Pacific upwelling. Nature 391, 270–273. Dugdale, R.C., Wilkerson, F.P., Chai, F., Feely, R., 2007. Size-fractionated nitrogen uptake measurements in the equatorial Pacific and confirmation of the low Sihigh-nitrate low-chlorophyll condition. Global Biogeochemical Cycles, 21. doi:10.1029/2006GB002722. Dunne, J.P, Murray, J.W., Aufdenkampe, A.K., Blain, S., Rodier, M., 1999. Silicon– nitrogen coupling in the equatorial Pacific upwelling zone. Global Biogeochemical Cycles 13, 715–726. Fiedler, P.C., Talley, L.D., 2006. Hydrography of the eastern tropical Pacific: a review. Progress in Oceanography 69, 143–180. Franck, V.M., Brzezinski, M.A., Coale, K.H., Nelson, D.M., 2000. Iron and silicic acid concentrations regulate Si uptake north and south of the Polar Frontal Zone in the Pacific Sector of the Southern Ocean. Deep-Sea Research II 47, 3315–3338. Frost, B.W., Franzen, N.C., 1992. Grazing and iron limitation in the control of phytoplankton stock and nutrient concentration: a chemostat analogue of the Pacific equatorial upwelling zone. Marine Ecology Progress Series 83, 291–303. Goericke, R., 1998. Response of phytoplankton community structure and taxon specific growth rates to seasonally varying physical forcing in the Sargasso Sea off Bermuda. Limnology and Oceanography 43, 921–935. Kaczmarska, I., Fryxell, G.A., 1995. Micro-phytoplankton of the equatorial Pacific: ˜ o. Deep-Sea Research II 42, 1401W meridianal transect during the 1992 El Nin 535–558. Kessler, W.S., 2006. The circulation of the eastern tropical Pacific: a review. Progress in Oceanography 69, 181–217. Kobayashi, F., Takahashi, K., 2002. Distribution of diatoms along the equatorial transect in the western and central Pacific during the 1999 La Nina conditions. Deep-Sea Research II 49, 2801–2821. Krause, J.W., Nelson, D.M., Lomas, M.W., 2009a. Biogeochemical responses to latewinter storms in the Sargasso Sea. II. Increased rates of biogenic silica production and export. Deep-Sea Research I 56, 861–874. Krause, J.W., Lomas, M.W., Nelson, D.M., 2009b. Biogenic silica at the Bermuda Atlantic Time-series Study site in the Sargasso Sea: temporal changes and their inferred controls based on a 15-year record. Global Biogeochemical Cycles 23, GB3004. doi:10.1029/2008GB003236. Ku, T.-L., Luo, S., Kusakabe, M., Bishop, J.K.B., 1995. 228Ra-derived nutrient budgets in the upper Equatorial Pacific and the role of ’’new’’ silicate in limiting productivity. Deep-Sea Research II 42, 479–497. Landry, M.R., Barber, R.T., Bidigare, R.R., Chai, F., Coale, K.H., Dam, H.G., Lewis, M.R., Lindley, S.T., McCarthy, J.J., Roman, M.R., Stoecker, D.K., Verity, P.G., White, J.R., 1997. Iron and grazing constraints on primary production in the central equatorial Pacific: an EqPac synthesis. Limnology and Oceanography 42, 405–418. Landry, M.R., Selph, K.E., Taylor, A.G., De´cima, M., Balch, W.M., Bidigare, R.R., 2011. Phytoplankton growth, grazing and production balances in the HNLC equatorial Pacific. Deep-Sea Research II 58 (3–4), 524–535.

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Latasa, M., Landry, M.R., Schluter, L., Bidigare, R.R., 1997. Pigment-specific growth and grazing rates of phytoplankton in the central equatorial Pacific. Limnology and Oceanography 42, 289–298. Leynaert, A., Treguer, P., Lancelot, C., Rodier, M., 2001. Silicon limitation of biogenic silica production in the Equatorial Pacific. Deep-Sea Research I 48, 639–660. Lukas, R., 1986. The termination of the equatorial undercurrent in the Eastern Pacific. Progress in Oceanography 16, 63–90. McCarthy, J.J., Garside, C., Nevins, J.L., Barber, R.T., 1996. New production along ˜ o event. 1401W in the equatorial Pacific during and following the 1992 El Nin Deep-Sea Research II 43, 1065–1093. Minas, H.J., Minas, M., Packard, T.T., 1986. Productivity in upwelling areas deduced from hydrographic and chemical fields. Limnology and Oceanography 31, 1182–1206. Mosseri, J., Queguiner, B., Armand, L., Cornet-Barthaux, V., 2008. Impact of iron on silicon utilization by diatoms in the Southern Ocean: a case study of Si/N cycle decoupling in a naturally iron-enriched area. Deep-Sea Research II 55, 801–819. Murray, J.W., Johnson, E., Garside, C., 1995. A U.S. JGOFS Process Study in the equatorial Pacific (EqPac): introduction. Deep-Sea Research II 42, 275–293. Nelson, D.M., Brand, L.E., 1979. Cell division periodicity in 13 species of marine phytoplankton on a light:dark cycle. Journal of Phycology 15, 67–75. Nelson, D.M., Brzezinski, M.A., 1997. Diatom growth and productivity in an oligotrophic midocean gyre: a 3-yr record from the Sargasso Sea near Bermuda. Limnology and Oceanography 42, 473–486. Nelson, D.M., Tre´guer, P., Brzezinski, M.A., Leynaert, A., Queguiner, B., 1995. Production and dissolution of biogenic silica in the ocean: revised global estimates, comparison with regional data and relationship to biogenic sedimentation. Global Biogeochemical Cycles 9, 359–372. Nelson, D.M., Brzezinski, M.A., Sigmon, D.E., Franck, V.M., 2001. A seasonal progression of Si limitation in the Pacific sector of the Southern Ocean. DeepSea Research II 48, 3973–3995. Parker, A.E., Wilkerson, F.P., Dugdale, R.C., Marchi, A., Hogue, V.E., Landry, M.R., Taylor, A.G., 2011. Spatial patterns of nitrogen uptake and phytoplankton in the equatorial upwelling zone (1101W–1401W) during 2004 and 2005. DeepSea Research II 58 (3–4), 417–433. Pennington, J.T., Mahoney, K.L., Kuwahara, V.S., Kolber, D.D., Calienes, R., Chavez, F.P., 2006. Primary production in the eastern tropical Pacific: a review. Progress in Oceanography 69, 285–317. Philander, S.G.H., 1978. Instabilities of zonal equatorial currents. Journal of Geophysical Research 83, 3679–3682. Raimbault, P., Slawyk, G., Boudjellal, B., Coatanoan, C., Conan, P., Coste, B., Garcia, N., Moutin, T., Pujo-Pay, M., 1999. Carbon and nitrogen uptake and export in

the equatorial Pacific at 1501W: evidence of an efficient regenerated production cycle. Journal of Geophysical Research 104, 3341–3356. Schlitzer, R., 2006. Ocean Data View, http://www.odv.awi.de. Shipe, R.F., Brzezinski, M.A., 2001. A time series study of silica production and flux in an eastern boundary region: Santa Barbara Basin, California. Global Biogeochemical Cycles 15, 517–531. Smith, C.R., Hoover, D.J., Doan, S.E., Pope, R.H., DeMaster, D.J., Doobs, F.C., Altabet, M.A., 1996. Phytodetritus at the abyssal seafloor across 10 of latitude in the central equatorial Pacific. Deep-Sea Research II 43, 1309–1338. Steinberg, D.K., Carlson, C.A., Bates, N.R., Johnson, R.J., Michaels, A.F., Knap, A.H., 2001. Overview of the US JGOFS Bermuda Atlantic Time-series Study (BATS): a decade-scale look at ocean biology and biogeochemistry. Deep-Sea Research II 48, 1405–1447. Strutton, P.G., Palacz, A.P., Dugdale, R.C., Chai, F., Marchi, A., Parker, A.E., Hogue, V., Wilkerson, F.P., 2011. The impact of equatorial Pacific tropical instability waves on hydrography and nutrients: 2004–2005. Deep-Sea Research II 58 (3–4), 284–295. Sunda, W.G., Huntsman, S.A., 1997. Interrelated influence of iron, light and cell size of marine phytoplankton growth. Nature 390, 389–392. Takahashi, T., Sutherland, S.C., Wanninkhof, R., Sweeney, C., Feely, R.A., Chipman, D.W., Hales, B., Friederich, G., Chavez, F., Sabine, C., Watson, A., Bakker, D.C.E., Schuster, U., Metzl, N., Yoshikawa-Inoue, H., Ishii, M., Midorikawa, T., Nojiri, Y., ¨ Kortzinger, A., Steinhoff, T., Hoppema, M., Olafsson, J., Arnarson, T.S., Tilbrook, B., Johannessen, T., Olsen, A., Bellerby, R., Wong, C.S., Delille, B., Bates, N.R., de Baar, H.J.W., 2009. Climatological mean and decadal change in surface ocean pCO2, and net sea–air CO2 flux over the global oceans. Deep-Sea Research II 56, 554–577. Takeda, S., 1998. Influence of iron availability on nutrient consumption ratio of diatoms in oceanic waters. Nature 393, 774–777. Taylor, A.G., Landry, M.R., Selph, K.E., Yang, E.J., 2011. Biomass, size structure and depth distributions of the microbial community in the eastern equatorial Pacific. Deep-Sea Research II 58 (3–4), 342–357. Twining, B.S., Baines, S.B., Fisher, N.S., 2004. Element stoichiometries of individual plankton cells collected during the Southern Ocean Iron Experiment (SOFeX). Limnology and Oceanography 49, 2115–2128. Wilkerson, F.P., Dugdale, R.C., 1996. Silicate versus nitrate limitation in the equatorial Pacific estimated from satellite-derived sea-surface temperatures. Advances in Space Research 18, 81–89. Yoder, J.A., Ackleson, S.G., Barber, R.T., Flament, P., Balch, W.M., 1994. A line in the sea. Nature 371, 689–692.