Iron, manganese, and lead at Hawaii Ocean Time-series station ALOHA: Temporal variability and an intermediate water hydrothermal plume

Geochimica et Cosmochimica Acta, Vol. 69, No. 4, pp. 933–952, 2005 Copyright © 2005 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/05 $30.00 ⫹ .00

doi:10.1016/j.gca.2004.07.034

Iron, manganese, and lead at Hawaii Ocean Time-series station ALOHA: Temporal variability and an intermediate water hydrothermal plume EDWARD A. BOYLE,1,* BRIDGET A. BERGQUIST,1 RICHARD A. KAYSER,1 and NATALIE MAHOWALD2 1

Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USA 2 National Center for Atmospheric Research, PO Box 3000, Boulder, CO 80307, USA (Received September 13, 2003; accepted in revised form July 23, 2004)

Abstract—Trace metal clean techniques were used to sample Hawaii Ocean Time-series (HOT) station ALOHA on seven occasions between November 1998 and October 2002. On three occasions, full watercolumn profile samples were obtained; on the other four occasions, surface and near-surface euphotic zone profiles were obtained. Together with three other published samplings, this site may have been monitored for “dissolved” (ⱕ0.4 or ⱕ0.2 ␮m) Fe more frequently than any other open ocean site in the world. Low Fe concentrations (⬍0.1 nmol kg⫺1) are seen in the lower euphotic zone, and Fe concentrations increase to a maximum in intermediate waters. In the deepwaters (⬎2500 m), the concentrations we observe (0.4 – 0.5 nmol kg⫺1) are significantly lower than some other deep North Pacific stations but are similar to values that have been reported for a station 350 miles to the northeast. We attribute these low deepwater values to transport of low-Fe Antarctic Bottom Water into the basin and a balance between Fe regeneration and scavenging in the deep water. Near-surface waters have higher Fe levels than observed in the lower euphotic zone. Significant temporal variability is seen in near-surface Fe concentrations (ranging from 0.2– 0.7 nmol kg⫺1); we attribute these surface Fe fluctuations to variable dust deposition, biological uptake, and changes in the mixed layer depth. This variability could occur only if the surface layer Fe residence time is less than a few years, and based on that constraint, it appears that a higher percentage of the total Fe must be released from North Pacific aerosols compared to North Atlantic aerosols. Surprisingly, significant temporal variability and high particulate Fe concentrations are observed for intermediate waters (1000 –1500 m). These features are seen in the depth interval where high ␦3He from the nearby Loihi Seamount hydrothermal fields has been observed; the total Fe/3He ratio implies that the hydrothermal vents are the source of the high and variable Fe. The vertical profile of Mn at ALOHA qualitatively resembles other North Pacific Mn profiles with surface and intermediate water maxima, but there are some significant quantitative differences from other reported profiles. The ⱕ0.4 ␮m Mn concentration is highest near the surface, decreases sharply in the upper 500 m, then shows an intermediate water maximum at 800 m and then decreases in the deepest waters; these concentrations are higher than observed at a station 350 miles to the northeast that shows similar vertical variations. It appears that there is a significant Mn gradient (throughout the water column) from HOT towards the northeast. Compared to the first valid oceanic Pb data for samples collected in 1976, Pb at ALOHA in 1997–1999 shows decreases in surface waters and waters shallower than 200 m. Pb concentrations in central North Pacific surface waters have decreased by a factor of 2 during the past 25 yr (from ⬃65 to ⬃30 pmol kg⫺1); surface water Pb concentrations in the central North Atlantic and central North Pacific are now comparable. We attribute the surface water Pb decrease to the elimination of leaded gasoline in Japan and to some extent by the U.S. and Canada. We attribute most of the remaining Pb in Pacific surface waters to Asian emissions, more likely due to high-temperature industrial activities such as coal burning rather than to leaded gasoline consumption. A 3-year mixed-layer time series from the nearby HALE-ALOHA mooring site (1997–1999) shows that there is an annual cycle in Pb with concentrations ⬃20% higher in winter months; this rise may be created by downward mixing of the winter mixed layer into the steep gradient of higher Pb in the upper thermocline (Pb concentrations double between the surface and 200 m). From 200 m to the bottom, Pb concentrations decrease to levels of 5–9 pmol kg⫺1 near the bottom; for most of the water column, thermocline and deepwater Pb concentrations do not appear to have changed significantly during the 23-yr interval. Copyright © 2005 Elsevier Ltd 1988; Martin and Gordon, 1988) and possibly as a factor limiting nitrogen fixation by diazotrophs (Falkowski, 1997; Kustka et al., 2002) is generally recognized, but data on the spatial and temporal variability of iron in the ocean are extremely sparse. This paucity of data results from the low concentrations of iron (Fe) in seawater combined with the high concentrations of Fe in the human environment; uncontaminated sampling and analysis is difficult and time-consuming. Each new Fe profile leads to significant improvements in our understanding of the oceanic Fe distribution; this distribution

1. INTRODUCTION

Studies of Fe, Mn, and Pb in the ocean contribute to our understanding of larger issues in global biogeochemical cycling. Iron’s role as a limiting nutrient in high-nutrient lowchlorophyll environments (Boyd et al., 1999; Coale et al., 2004; Coale et al., 1996; Martin et al., 1994; Martin and Fitzwater,

* Author to whom ([email protected]).

correspondence

should

be

addressed 933

934

E. A. Boyle et al.

Fig. 1. Map showing location of HOT-ALOHA and Johnson et al. (2003) and MP5 (this paper) surface water transects, plotted on counters of mean annual dust fluxes (1980 –2002) from model of Luo et al., 2003. S&P denotes the 1976 Pb station of Schaule and Patterson (1981).

has proven enigmatic and difficult to model (Archer and Johnson, 2000; Lefevre and Watson, 1999; Parekh et al., 2004; Watson et al., 2000; Watson et al., 1994). Similarly, beginning with the work of Patterson and colleagues (Flegal and Patterson, 1983; Schaule and Patterson, 1983; Schaule and Patterson, 1981), it is widely acknowledged that anthropogenic lead dominates the lead distribution of the ocean, but there is very little data documenting the time-dependent response of the ocean in this ongoing global geophysical experiment. Finally, although there is a more significant body of information on the distribution of manganese in the ocean (Bruland and Franks, 1983; Klinkhammer and Bender, 1980; Landing and Bruland, 1987), the oceanic distribution of Mn is complicated and we do not fully understand the mechanisms controlling its distribution. Because Mn (in common with Fe) has a significant aerosol input, surface water Mn data may help unravel some questions about oceanic Fe as well. We have had the opportunity to sample Hawaii Ocean Time-series (HOT) station ALOHA (Karl et al., 2002; Karl and Lucas, 1996) seven times during the course of a 4-year period (1998 –2002). These data together with three previously published station occupations make this station possibly the most intensely studied site for Fe in the ocean. Some of these samples were also analyzed for Mn and Pb, with the Pb station data supplemented by surface water data from a moored in situ trace element serial sampler (MITESS, Bell et al., 2002).

2. FIELD WORK, SAMPLE HANDLING, AND ANALYSIS

We sampled station HOT-ALOHA (22°45=N, 158°00=W, 4800 m bottom depth; Fig. 1) on seven occasions during five cruises: November 1998 (R/V Moana Wave, surface samples); May 1999 (R/V Moana Wave, deep profile); April 2001 (R/V Wecoma, deep profile and detailed euphotic zone); R/V Kaimikai-O-Kanaloa (KOK) July 2002 (deepwater samples and detailed euphotic zone); October 2002 (R/V Kilo Moana, surface and mixed layer samples). On two of these cruises, we sampled the station at the beginning and end of the cruise (July 2002 and October 2002). We collected most of the samples using various configurations of the MITESS Water sampler (Bell et al., 2002). Some near-surface samples were collected with an underway “towed fish” device (Vink et al., 2000), and a few other “pole” samples were obtained while steaming at two knots and extending a sample bottle from the side of the ship using a long pole. The MITESS sampler opens and closes an acid-cleaned 500-mL polyethylene or polymethylpentene bottle in situ, minimizing opportunities for external contamination during sample collection. The various configurations of the device are: (a) 12 unit sampler suspended beneath a hydrowire for deep profiling; (b) Automated Trace Element (“ATE”), a single-module sampler deployed at 10 m depth suspended from plastic-coated wire; and (c) “ATE/Vane,” a single MITESS module attached to a wire-mounted “weathervane” device that enables sample collection upstream of wire contamination. On these cruises,

Fe, Pb, and Mn variability in central North Pacific (HOT)

MITESS was deployed with a temperature recording device; sample depths are derived by comparison of the MITESS sample temperatures to a conductivity-temperature-depth recorder (CTD) deployed at the same site on the same cruise. Within 12–24 h upon retrieving a profiling MITESS, sealed sample bottles are taken into a class 100 clean flow environment for filtration of a subsample through acid-leached 0.4-␮m Nuclepore filters and acid-rinsed 0.02 ␮m Anotop cartridges containing Anopore alumina filters (see Wu et al., 2001 for further details on this filtration). Some workers prefer to use 0.2-␮m filters to avoid possible bacterial leakage, but our tests of 0.4-␮m Nuclepore and 0.1-␮m Teflon filtrates on samples of Hawaiian surface waters showed little difference between the Fe concentrations of the filtrates, suggesting that there is little Fe in the 0.1– 0.4-␮m size class. Filters are rinsed several times each with dilute acid and the seawater sample before collecting filtrate. Usually, we collect two or three sequential portions into separate bottles with the filtration sequence noted. The last filtrate is considered to be the least likely to suffer from filter holder contamination (because it has been rinsed the most), and therefore this sample is usually the one selected for analysis. In some cases bottles are randomly contaminated, so the previous filtrate is analyzed and retained as the correct value if it is lower and oceanographically consistent. The MITESS sampler also was used in programmed automated sampling mode for five 6-month deployments (May 1997–January 2000) at the HALE-ALOHA mooring (22.5°N, 158.2°W). Two mooring units (designated “A” for the upper unit and “B” for the lower unit), each with 12 sampling modules, were deployed 3 m apart in the 30 – 40 m depth range (almost always within the mixed layer at this site). In the November 1997–May 1998 deployment, a possible artifact in the “A” unit should be noted: a somewhat rusty recording CTD was deployed a few meters above the “A” unit, and the opening between the top of MITESS and the enclosed sample bottles was not sealed as it should have been. Four of the winter “A” samples had higher Pb concentrations than the “B” samples during this deployment, therefore we plot these results in parentheses because we are uncertain of the veracity of these samples. However, the next year, four similar high values are observed in both “A” and “B” units and an independent surface sample during the same season, therefore it is possible that the four high values from winter in the previous winter being seen in the same sampler could just be a statistical coincidence, hence we cannot confidently reject high samples from the previous year. Subsamples (1.3 mL) are analyzed for Fe by isotope dilution inductively coupled plasma mass spectrometry (ICPMS). The 1.3-mL samples analyzed by Wu et al. (2001) and discussed here, as well as the “Honolulu to HOT” surface water transect to be shown later, were determined using 57Fe isotope dilution on a Finnegan Element at Rutgers University using medium mass resolution to eliminate 40Ar16O⫹ and 40Ca16O⫹ interferences from the 56Fe⫹ signal (Wu and Boyle, 1998). Most of the samples reported here were analyzed by 54Fe isotope dilution plasma mass spectrometry on a Micromass IsoProbe multicollector instrument using a hexapole collision cell using Ar and H2 collision cell gases that destroy polyatomic 40Ar16O⫹ and 40 Ar14N⫹ interferences. To intercompare these two methods, we repeatedly analyzed a seawater sample where 11 replicates

935

had been measured using the high-resolution method on a single day giving 0.52 ⫾ 0.03 nmol kg⫺1. The hexapole/ multicollector method replicates analyzed on several different analytical sessions differ from that value by less than 0.05 nmol kg⫺1. The multicollector/hexapole collision cell method simultaneously provides Mn data included here (the Mn recovery and ionization efficiencies relative to 54Fe are determined by analysis of a seawater Mn standard and standard additions experiments on some samples); where samples have been analyzed by this method, the Mn concentration is also reported. Typical 1␴ precisions for Fe and Mn analyses are 0.03 and 0.05 nmol/kg. Subsamples (1 mL) are analyzed for Pb by 204Pb isotope dilution ICPMS after Mg(OH)2 separation and preconcentration; the 208Pb/204Pb ratio is determined by quadrupole ICPMS using a VG/Fisons PQ2⫹ (Wu and Boyle, 1997a). The Pb analyses are performed on unfiltered samples to minimize the chance for contamination; however, studies have shown that the particulate Pb is low (e.g., see Bacon et al., 1976) and hence these unfiltered analyses essentially give dissolved Pb. For some samples, we analyzed acidified unfiltered sample aliquots for total Fe using the same Mg(OH)2 co-precipitation method. Immediately before subsampling, these unfiltered samples are shaken thoroughly to distribute the particulate Fe as evenly as possible. Although data from samples that have been acidified and then preconcentrated for trace metal analysis are often referred to as “total dissolvable” metal concentration (to allow for some portion of the particulate metal that is not dissolved), we believe that the data obtained with our method represent total Fe because dissolved Fe is recovered ⬎90% by the co-precipitation and particulate Fe is trapped within the Mg(OH)2 pellet. After dissolution of the soluble iron with the acidification of the Mg(OH)2, the particles are entrained into the nebulized solution and aspirated into the plasma where they are vaporized and ionized. During analysis of our “total Fe” samples, we observe a higher noise level in the Fe signal as individual particles are vaporized and ionized in pulses. All samples are analyzed in triplicate. The Fe and Mn analyses are conducted simultaneously when the multicollector/ hexapole method is used. Pb is always measured in a separate analysis. When at least two of the triplicates agree within expected reproducibility, the average of the two or three concordant samples is taken as the correct concentration. If all three samples deviate beyond expectations, or if a single anomalous sample is too low (rather than too high as expected from erratic procedural contamination), then the samples are run in triplicate again. We report the data as the average of the concordant replicates, showing the standard deviation and the number of analyses included in the mean. In some cases, “oceanographic consistency” (the expectation that coarse resolution vertical profiles of ocean properties will vary relatively smoothly with depth) indicates that a sample bottle is contaminated even if analytical replication is good. In that case, if we have a second sample bottle (the duplicate filtrate) available, we run that sample bottle in triplicate. If the new result is lower and “consistent,” we then report that analysis and discard the contaminated bottle data. If there is no replicate filtrate, or if both bottles indicate high values, then we report the result in the data table but indicate our belief that the sample is contaminated in the data table with a question mark, omitting that sample from consideration in the discussion.

936

E. A. Boyle et al.

sea). Deepwater Fe concentrations obtained from these latest samples agree well with reanalyses of the April 2001 samples after acidification for more than a year (Fig. 2).We also note that our deepwater values in the 3000 – 4000 m range are similar to those of Bruland et al. (1994) for a station northeast of HOT (see later). Hence, we believe that the analyses reported here from the long-term acidified 2001 samples are correct, and we will use these in preference to values reported by Wu et al. (2001), but it should be emphasized that the differences are small (⬍0.1 nmol kg⫺1) and only slightly exceed our analytical error. Also note that the “depths” reported by Wu et al. (2001) were actually “wire out” and the depths we report here are based on the temperature recorder-CTD comparison as discussed above. Despite the absence of a direct intercomparison, we believe that our data are accurate to better than 10% and intercomparable to the datasets of the laboratories of Patterson, Bruland, and Martin, Coale, and Johnson, based on these considerations: (1) the similarity of the Pb data of Schaule and Patterson and our Pb data in the deep water column of the Pacific (which has a residence time of centuries and hence could not have changed in 26 yr); (2) our 20-yr time series of Pb in Bermuda surface waters agrees with the less frequent samplings of Patterson’s laboratory (Wu and Boyle, 1997b); and (3) the similarity of our deep-water-column data for Fe with the data of Bruland et al. at a nearby station. Fig. 2. ALOHA Fe profile. Lines are drawn between adjacent samples for a given station occupation except when a significant depth gap occurs between samples.

Procedural blanks are established by running a 50-␮L aliquot of a low-Fe seawater sample through the same procedure as samples. Measurement precision and accuracy is a function of several factors: (a) reproducibility of isotope ratio measurements on the mass spectrometer (typically 1–2% for Fe, Mn, and Pb); (b) uncertainty in the true analytical blank value (typically ⬃0.2 ⫾ 0.03 nmol kg⫺1 for Fe and Mn; 3 ⫾ 1 pmol kg⫺1 for Pb); and (c) random contamination during sample handling and analysis (indeterminate). We minimize the latter error by running samples in triplicate to assess procedural random contamination and by collecting duplicate samples that are both analyzed if one gives unexpectedly high results. We collected water column vertical profile samples on three occasions. On the HOT 105 R/V Moana Wave 1999 cruise, particle leakage around the insufficiently tightened 0.4-␮m filter holder compromised the filtrates, but “total” and ⬍0.02-␮m filtrates from this cruise were satisfactory. We resampled the station on the R/V Wecoma in April 2001. However, samples from this cruise were left unacidified for nearly a month, with the potential for loss of Fe onto the container wall during that time. Wu et al. (2001) reported data from these samples for analyses performed 8 days after the samples were acidified. Our reanalyses here of the 2001 samples 1.5 yr after acidification indicate that some of the numbers reported by Wu et al. (2001) may have been slightly low because of incomplete redissolution from the bottle walls. This conclusion is supported by our data from another deep-water sampling effort (KOK, July 2002; samples filtered and immediately acidified at

Fig. 3. CTD T, S, O2, and sigma theta profiles from our HOT profile occupations.

Fe, Pb, and Mn variability in central North Pacific (HOT)

937

Table 1. ALOHA profile samples. TOTAL (unfiltered) Cruise HOT 105 (MW) May 11, 1999

Depth 1 35 70 102 192 452 686 730 1118 1118 1405 1724 1724 3000 4000

T

22.719 21.469 16.748 7.735 4.771 4.636 3.653 3.651 3.002 2.503 2.501 1.518 1.453

est. S

35.372 35.322 34.743 34.051 34.230 34.291 34.512 34.512 34.564 34.603 34.603 34.679 34.691

est. O2

Fe

␴

n

0.31 0.06

221.7 216.3 202.9 146.2 29.1 25.3 46.8 46.8 60.1 73.7 73.7 127.3 147.3

*0.92 *0.76 *0.53 *0.46 *0.76 *0.81 0.26 1.93 *2.34 2.54 2.41 *1.94 1.68 *1.67 *1.89

2 2 1 2 2 2 2 2 2 1 2 2 2 1 2

0.05 0.18 0.04 0.04 0.19 0.01 0.05 0.06 0.07 0.18

Mn

␴

n

0.67 0.61

0.01 0.01

2 2

0.62 0.56

0.02

1 2

0.47

0.01

2

* High-resolution ICPMS ⱕ0.4 ␮m (filtered)

TOTAL (unfiltered) Cruise

Depth

MP2 (Wec) April 26, 2001

1 20 75 317 476 807 998 1162 1260 1517 1951 2848 3322 3797 10 20 30 70 110 986 1194 1429 1739 2432 2886 3848

MP5 (KOK) July 1, 2002

T

est. S

est. O2

23.611 23.595 12.125 7.587 4.494 3.869 3.519 3.240 2.735 2.085 1.569 1.472 1.464

35.100 35.153 34.268 34.022 34.311 34.467 34.512 34.536 34.576 34.621

205.6 203.3 222.4 151.8 25.9 45.3 54.7 61.2 74.9 97.9

23.559 22.086 4.037 3.466 2.970 2.515 1.806 1.599 1.461

Fe

3.48 2.35 3.93 2.28 2.65 2.79

␴

0.10 0.25 0.61 2.16 0.90 0.31

n

2 4 3 2 2 2

Mn

0.74 0.68 0.61 0.55 0.59 0.51

s

0.02 0.06 0.10 0.35 0.00 0.53

n

2 4 4 2 2 2

3.25

0.60

2

0.66

0.03

2

2.03 1.64

0.66 0.01

2 2

0.64 0.57

0.14 0.01

2 2

Fe

␴

n

Mn

s

n

0.65 0.74 0.41 0.40 0.44 0.50 0.47 0.76 0.65 0.58 0.58 0.41 0.45 0.41 0.46 0.32 0.44 0.27 0.09 0.84 1.00 0.75 0.77 0.48 0.49 0.44

0.11 0.05 0.07 0.06 0.09 0.02 0.01 0.04 0.09 0.03 0.03 0.03 0.02

2 2 2 2 3 3 2 3 2 3 3 3 2 1 2 3 4 2 2 7 9 4 4 4 4 7

1.39 1.46 1.35 0.49 0.38 0.64 0.54 0.44 0.50 0.50 0.41 0.43 0.38 0.31 1.62 1.67 1.55 1.35 1.25 0.60 0.54 0.42 0.42 0.43 0.37 0.26

0.02 0.02 0.07 0.02 0.02 0.00 0.01 0.01 0.01 0.01 0.01 0.00 0.02 0.01 0.09 0.02 0.11 0.00 0.00 0.04 0.05 0.07 0.02 0.05 0.02 0.03

3 2 5 3 3 3 3 3 3 3 3 3 3 3 4 3 5 2 3 11 9 9 6 6 9 9

0.02 0.01 0.04 0.04 0.04 0.05 0.05 0.03 0.09 0.06 0.09 0.03

MW ⫽ Moana Wave; Wec ⫽ Wecoma; KOK ⫽ Kaimikai-O-Kanaloa. ? Indicates replicable analysis rejected as a contaminated sample by “oceanographic consistency”. Fe, Mn in nmol kg⫺1; T in °C, S in ppt; O2 in ␮mol kg⫺1. Est. salinity and est. O2 are derived from CTD value that matches MITESS temperature.

3. RESULTS AND DISCUSSION

3.1. Standard Oceanographic Properties The hydrography of station ALOHA has been discussed in detail (Karl and Lucas, 1996). For guidance in interpreting the metal data, CTD temperature, salinity, dissolved oxygen, and sigma theta profiles obtained at station ALOHA from two station occupations (May 1999 and April 2002) are shown in Figure 3.

3.2. Fe Water Column Profiles Water column profiles for ⱕ0.4 ␮m Fe from the April 2001 and July 2002 samples (Table 1) are shown in Figure 2. The Fe concentrations decrease from the surface to a minimum in the lower euphotic zone, similar to the observations by Bruland et al. (1994) for a station 350 miles to the northeast (28°N, 155°W) and previously at HOT-ALOHA by Rue and Bruland (1995). This Fe decrease from the surface has been attributed to

938

E. A. Boyle et al.

Fe input to the surface waters by atmospheric dust followed by downward mixing and biologic uptake and/or scavenging within the euphotic zone (Bruland et al., 1994; Wu et al., 2001). The elevated surface water concentrations differ significantly between these station occupations. We attribute these differences to sporadic and seasonal changes in dust input (discussed later in this paper). In the deep water column (⬎2500 m), Fe concentrations are 0.4 – 0.5 nmol kg⫺1, with both of our occupations yielding similar values. These deepwater values are comparable to those Bruland et al. (1994) reported farther to the north (28°N, 155°W). They report 0.43 nmol kg⫺1 at 3491 m and 0.40 nmol kg⫺1 at 3974 m; we find 0.45 nmol kg⫺1 at 3342 m, 0.42 nmol kg⫺1 at 3797 m, and 0.45 nmol kg⫺1 at 3848 m. We attribute these low deepwater concentrations to advection of low-Fe (⬃0.4 nmol kg⫺1) waters derived from the Southern Ocean (De Baar et al., 1999; Measures and Vink, 2001; Sedwick et al., 2000) into the Pacific, accompanied by a negligible net Fe addition or removal because of balanced sources and sinks. 3.2.1. Intermediate Water Hydrothermal Fe Plume Although it is expected that Fe concentrations will be higher in intermediate waters because of regeneration from sinking biogenic particles, the behavior of Fe at HOT-ALOHA is anomalous because the middepth concentrations change on time scales of a year or less. At ⬃1250 m, Fe in the July 2002 occupation was ⬃25% higher than the Fe found during the April 2002 station occupation. This degree of temporal variability is unusual for nutrient-type elements in this depth range. Properties such as P, NO3⫺, or Si show little temporal variability at depth (e.g., see data from Bermuda Atlantic Time Series (BATS) and HOT time-series stations). We attribute part of the intermediate water Fe enrichment and the temporal variability to the proximity to HOT relative to the nearby Loihi Seamount submarine hydrothermal vents (HOT is 463 km northwest of Loihi) and near-field hydrodynamic variability. Loihi seamount (18°55=N; 156°15=W) is the southernmost extension of the Hawaiian Island volcanic chain. This volcanic complex erupts at depths of 1000 –1500 m (Wheat et al., 2000). The interaction of seawater, hot rock, and mantle plume volatiles creates hydrothermal vent fields that emit elevated concentrations of 3He and transition metals. At the highest observed temperatures (a bit over 30°C) from the first observations at this site, dissolved Fe occurs at levels of ⬃750 ␮mol kg⫺1 and Mn at levels of ⬃20 ␮mol kg⫺1 (Sedwick et al., 1992). The vent fluids have a 3He/4He ratio ⬃27 times the atmospheric ratio (Hilton et al., 1998), a CO2/He ratio ⬃3 ⫻ 109 (with a range of 6 ⫻ 108 to 8 ⫻ 109: Hilton et al., 1998), and a CO2 content of ⬃0.3 mol kg⫺1 (Sedwick et al., 1992). This chemical composition differs from typical ridge crest systems in having extremely high CO2 levels and an order-ofmagnitude higher Fe/Mn ratio. Yuan et al. (1994) surveyed the total Fe concentration of samples immediately above Loihi hydrothermal fields and observed elevated Fe concentrations in the depth range from 1100 to 1200 m. Lupton (1998) demonstrated that volcanic 3He emissions from these seamounts influence oceanic ␦3He values over a broad region of the intermediate Pacific Ocean near Hawaii (Fig. 4a). A ␦3He profile from the vicinity of the HOT-ALOHA

Fig. 4. (a) ␦3He Pacific Ocean horizontal distribution at 1100 m and (b) ␦3He profile (22.5°N, 155°W) near HOT station (Lupton, 1998).

station (463 km northwest of Loihi) shows elevated ␦3He from 900 to 1500 m with a sharp maximum at 1100 m (⬃25% compared to ⬃20% at 2500 m; Fig. 4b). The ␦3He data are incontrovertible evidence for the presence of a hydrothermal component from Loihi seamount at this depth range. A 500,000-fold dilution of observed 30°C vent fluids with ambient seawater produces ␦3He ⫽ 25% (as observed) and indicates that seawater near HOT-ALOHA at ⬃1100 m should be significantly enriched in Fe (1.6 nmol kg⫺1) and only slightly enriched for Mn (0.04 nmol kg⫺1) if these elements are not lost by precipitation and settling in transit. These estimates are consistent with the observations of total Fe and total Mn in the

Fe, Pb, and Mn variability in central North Pacific (HOT)

939

concentrations, but we do not know how to estimate their magnitude and influence at this site. Given that the hydrothermal influence on ␦3He is incontrovertible and the dilution factors are consistent for 3He and Fe, we believe that hydrothermal discharge is the most likely source of the intermediate water Fe anomaly. 3.3. Near-Surface Fe Variability: Evidence for High Dust Solubility Measurements of dust (mineral aerosols) at the Moana Loa Observatory (MLO) demonstrate a strong seasonality with maximum concentrations in early spring. Parrington et al. (1983) measured aerosol Al at MLO from 1979 to 1982, Perry et al. (1999) measured fine-soil (⬍2.5 ␮m) aerosols at MLO from 1989 to 1996, and Perry (in Johnson et al., 2003) measured MLO aerosol Fe at high temporal resolution from midMarch 2001 until mid-April 2001 (Fig. 7). These data demonstrate a pronounced seasonal cycle in dust aerosols of more than an order of magnitude, with lowest aerosol concentrations occurring between July and January and highest concentrations observed from March through June. The dust can be sporadic with very strong events lasting only a few days (Gong et al., 2003). Mineral aerosols at MLO also show significant interannual differences in onset and duration of the high dust season. This aerosol variability is caused by outbreaks of dust from central Asia. The Total Ozone Mapping Spectrometer (TOMS) aerosol index (AI) (derived from spectral attenuation of Rayleigh scattering caused by aerosol absorption using 340- and 380-nm bands) allows for the determination of the temporal Fig. 5. Total Fe compared to ⬍0.4 ␮m Fe at HOT. Note ⬃2–3 nmol kg⫺1 maximum in total Fe at 700 –1300 m depth.

Loihi plume near station ALOHA (Table 1, Figs. 5 and 6). Elevated total Fe is seen at depths ranging from 700 to 1300 m. The Fe concentration increase appears most strongly in the particulate form, exceeding ⱕ0.4 ␮m Fe by a factor of 4 – 8. For Mn, dissolved⫹particulate levels are only slightly higher than the dissolved fraction, and only small enhancements above ambient Mn in the region (⬃0.5 nmol kg⫺1) are seen in the Loihi plume depth range. Dissolved Fe and total Fe show significant variability between station occupations. We attribute this variability to the near-field variability of deep-sea transport away from Loihi to station HOT-ALOHA. Just as a smokestack plume shows “puffs” in the vicinity of the source, it is expected that the dilution factors for even a relatively steady hydrothermal source will be reflected by pulsations caused by the unsteady deep sea circulation field. The hydrothermal source itself varies during an interval of several years (Hilton et al., 1998) although the chemical composition of the estimated end-member remains constant. Moore et al. (2001) have suggested that deep groundwater discharges could serve as a source of middepth radioisotope anomalies near Hawaii. Groundwater discharges could lead to trace metal anomalies as well. We cannot dispute the possibility of groundwater discharges having an influence on Fe and Mn

Fig. 6. Total and ⬍0.4 ␮m Mn profile at HOT.

940

E. A. Boyle et al.

Fig. 7. (A) (Parrington et al., 1983) Mauna Loa aerosol Al data, 1978 –1982. (B) (Perry et al., 1999) Moana Loa fine soil aerosol from 1989 to 1996. (C) (Johnson et al., 2003) aerosol Fe data from March–May 2001.

and spatial evolution of these dust outbreaks (⬍http://toms. gsfc.nasa.gov/aerosols/aerosols.html⬎). A very strong dust event occurred originating in China on April 6, 2001, spreading across most of the North Pacific from April 11–14. This dust storm was intensively studied by the Aerosol Characterization Experiment-Asia (ACE-Asia) project (Gong et al., 2003). Although vertical column integrated dust dominantly occurred at higher latitudes than HOT-ALOHA, high dust concentrations were seen at MLO during this event (although at MLO this event is not unusually strong, with aerosol Fe concentrations being lower than seasonal averages for 1989 –1996: Johnson et al., 2003; Perry et al., 1999). Our surface water samples at station HOT-ALOHA allow us to make a preliminary evaluation of the effect of this dust variability on Fe variability in surface waters near Hawaii. With the addition of our seven stations, the near-surface waters of station HOT-ALOHA have been sampled 10 times between January 1994 and October 2002, making HOT-ALOHA possibly the most intensively studied open-ocean site in the world for ⱕ0.4 ␮m Fe (Table 2). In Figure 8, we collapse all of the

samplings onto a single “annual cycle.” These data demonstrate several important features about Fe in the central North Pacific: (1) Although it is commonly believed that ⱕ0.4 ␮m Fe concentrations in the surface waters of the central Pacific are extremely low, this impression is largely based on samples from 50 to 100 m depth (Johnson et al., 1997). As originally shown by Bruland et al. (1994) and confirmed here by two euphotic zone profiles (Fig. 2), ⱕ0.4 ␮m Fe in the central North Pacific is high in the mixed layer and decreases to a minimum deeper in the euphotic zone. As our repeated samplings at station HOT-ALOHA demonstrate, near-surface waters at HOT-ALOHA have moderate levels of Fe, with a range of 0.21– 0.71 nmol kg⫺1 and a median value of 0.44 nmol kg⫺1). Clearly, at this mid-Pacific site Fe concentrations in nearsurface waters are not as low as previously expected. Although the biological availability of this Fe is difficult to assess, the ⱕ0.4 ␮m Fe concentrations present at this site for most of the year generally are not considered sufficiently low as to limit phytoplankton activity in surface waters (e.g., Blain et al., 2002 report half saturation constants of ⬍0.5 nmol/kg for phyto-

Fe, Pb, and Mn variability in central North Pacific (HOT)

941

Table 2. Other near-surface samples at HOT. Date

Cruise/Ship or reference

Depth (m)

Fe

1/20/94 11/7/98 5/11/99 5/11/99 5/11/99 3/26/01 4/27/01 4/27/01 5/22/01 7/3/02 7/3/02 7/17/03 10/1/02 10/1/02 10/1/02 10/13/02 10/13/02

Rue and Bruland (1995) HA-5B (Moana Wave) HOT 105 (Moana Wave) HOT 105 (Moana Wave) HOT 105 (Moana Wave) Johnson et al. (2003) MP2 (Wecoma) MP2 (Wecoma) Johnson et al. (2003) MP5 (KOK) MP5 (KOK) MP5 (KOK) MP6 (Kilo Moana) MP6 (Kilo Moana) MP6 (Kilo Moana) MP6 (Kilo Moana) MP6 (Kilo Moana)

20 1 1 1 10 1 1 20 1 10 20 10 0 10 40 10 40

0.24 0.50 0.51 0.43 0.30 0.21 0.63 0.69 0.33 0.47 0.35 0.36 0.57 0.38 0.54 0.42 0.50

␴

n

Mn

␴

n

0.03

1.53

0.02

2

0.01

2 1 2

1.48

0.01

2

0.10 0.04

2 2

1.39 1.46

0.02 0.02

3 2

0.04 0.02 0.05 0.04 0.04 0.04 0.05

1 3 2 2 2 2 3 3

1.62 1.67 1.42 1.22 1.23 1.24

0.09 0.02 0.01 0.02 0.00 0.02

4 3 3 3 2 3

KOK ⫽ Kaimikai-O-Kanaloa. Fe, Mn: nmol kg⫺1

plankton in the Indian Sector of the Southern Ocean, and Timmermans et al. (in press) report half-saturation constants of ⱕ1 nmol/kg for large Antarctic diatoms; given that smaller phytoplankton are less Fe limited than large diatoms and that Hawaiian surface waters are strongly N-limiting, ⬃0.4 nmol/kg levels are not likely to be limiting relative to N). Deeper in the euphotic zone (near 100 m) where Fe concentrations are ⬍0.1 nmol kg⫺1, Fe levels are likely to be limiting. (2) Fe shows significant temporal variability, both interannually and on a seasonal and subseasonal basis. The lowest concentration (0.24 nmol kg⫺1) was observed for a sample collected in mid-January. This low concentration occurred after 7 months of low aerosol dust and with a deep winter mixed layer (Fig. 9). Average Fe in the upper 110 m in July 2002 was 0.30 nmol kg⫺1, therefore part of the decrease in surface Fe concentrations in January relative to the summer and fall is simply a consequence of mixed layer homogenization of the upper ⬃100 m. The highest Fe concentration (0.71 nmol kg⫺1) is seen in late April 2001, in the middle of the high dust season and just after a strong Asian dust outbreak. This event was recorded at MLO (Fig. 7c) but dust transport was most intense farther north (30 –55°N) at an elevation of several km (Bishop et al., 2002; Gong et al., 2003; Johnson et al., 2003). From June through November, intermediate Fe concentrations are observed (averaging 0.47 nmol kg⫺1 ⫾ 0.08, 1␴). Significant interannual and subseasonal variability is evident. Johnson et al. (2003) report moderately low surface ocean Fe just before (0.21 nmol kg⫺1 on March 21, 2001) and just after (0.33 nmol kg⫺1 on May 23, 2001) our April 27, 2001 peak value (0.71 nmol kg⫺1). We suggest that this pattern makes the most sense if the residence time of “dissolved” Fe in the surface ocean is on the order of a half year in this oceanic regime. This time is sufficiently short to allow heterogeneities to develop in surface waters despite horizontal mixing, accounting for relatively rapid variations (as seen in spring 2001) due to dust input pulses and ocean surface water mesoscale transport variability. The high surface Fe induced by the annual dust peak will accumulate somewhat during spring and then begin to decrease

after dust deposition drops in May. If the residence time was much shorter than a half year, the Fe would decrease to very low levels by November, in contrast to our observations. The minimum Fe concentration occurs in early winter as the mixed layer deepens and surface Fe mixes with the low Fe waters deeper in the euphotic zone. For example, in late May 2001, Johnson et al. (2003) observed 0.33 nmol kg⫺1 “reactive” Fe in near-surface waters at HOT-ALOHA, but Fe was 0.67 nmol kg⫺1 at a station 200 miles to the northeast (similar to our observation at HOT-ALOHA in the previous month). From this evidence, we suggest that near-surface Fe in the central North Pacific will show (sporadically) high levels in the early spring, moderate concentrations in the summer and fall, and lowest concentrations in early winter. We also expect that significant interannual differences in near-surface Fe will occur as a result of interannual variability in Asian dust transport. Further support for the effect of dust on surface water Fe comes from a deployment of MITESS on the HALE-ALOHA mooring May-October 1999 at 34 –37 m (Table 3). This deployment depth is usually within the mixed layer at station HOT-ALOHA (e.g., from 1989 to 1993, the mixed layer depth at HOT-ALOHA varied between 30 and 120 m). On this deployment, we collected duplicate samples (using different MITESS modules) for each sample date. This duplicate sampling strategy was chosen because samples have a ⬃10% chance of being contaminated and duplicate sampling allows us to eliminate these by comparison to the lower replicate sample. For example, for this deployment 2 of 19 samples were high compared to the duplicate sample. As seen in Figure 10, total Fe (particulate⫹colloidal⫹soluble) decreases from 1.84 nmol kg⫺1 on June 1, 1999 to 0.95 nmol kg⫺1 on July 1, 1999. Except for a brief rise to 1.07 nmol kg⫺1 in mid-August, total Fe remains low through October (six samples averaged 0.78 nmol kg⫺1, excluding the mid-August sample). Although some of this total Fe may be bound in refractory sites that are not released into the ocean biogeochemical cycle, these data are consistent with the influence of spring dust deposition at station ALOHA deduced from the “dissolved” data.

942

E. A. Boyle et al.

Fig. 8. Near-surface Fe at HOT-ALOHA (by time of year).

Johnson et al. (2003) noted that the rise in Fe concentrations at HOT between late March and late April 2002, together with estimates of dust flux from models and aerosol deposition velocities, implies solubilization fractions much greater than 10% for this event. Now that we have shown that moderate Fe levels occur throughout the year, it appears that Fe solubilization fractions must be high in general. For example, taking the typical annual concentration of Fe at HOT-ALOHA as 0.4 nmol kg⫺1 for the upper 75 m and taking the annual average Fe flux as 10 mg m⫺2 yr⫺1 (Gao et al., 2001; Mahowald et al., 2003; Parrington et al., 1983), the residence time of Fe at HOT-ALOHA would be 17 yr if only 1% of the Fe were solubilized. That residence time estimate is clearly too high given the observed temporal variability. It seems more likely that a higher fraction of the aerosol Fe dissolves; if ⱖ10% of the aerosol Fe dissolves, the residence time would be shorter than a couple of years and the observed temporal and spatial variability would be reasonable. These data imply that the solubilization fraction (the Fe released to seawater relative to the total aerosol Fe deposition) of Fe for Pacific aerosols must be ⱖ10%, and if

our estimate for a residence time of a half year is assumed, the solubilization fraction would be ⬃30 – 40%. This observation points to an important difference between Fe cycling in the central North Pacific compared to the North Atlantic. At the same latitude in the western North Atlantic, we find that ⬍0.4 ␮m Fe is ⬃0.9 nmol kg⫺1 (only a factor of 2–3 higher than typical levels at HOT-ALOHA; Bergquist, 2004), but dust flux estimates for the North Atlantic are much higher than at HOT-ALOHA. For example, estimated dust fluxes are ⬃40-fold higher near Barbados than near Hawaii (Gao et al., 2001; Mahowald et al., 2003). Given surface water Fe concentrations that differ by less than a factor of 3, either the dust solubilization fractions must be an order of magnitude higher in the central North Pacific, or the residence time of Fe must be an order of magnitude longer in the central North Pacific (which seems unlikely given the observed temporal variability of Fe). It seems that significant differences must exist in the net solubilization of aerosol dust in the central North Atlantic and central North Pacific. Several factors could account for these differences, such as ecological differences that enhance the extraction of Fe from North Pacific aerosols or from differences

Fe, Pb, and Mn variability in central North Pacific (HOT)

943

Fig. 9. Moored sampler (MITESS) Pb data from the HALE-ALOHA mooring, 1997–1999, with mixed layer depth (from HOT annual reports [Tupas et al., 1997–2000]) and primary productivity (Karl et al., 2002).

in the mineral aerosol source chemistry and/or its subsequent atmospheric chemistry. A shorter residence time for Fe in the North Atlantic could result (even with the same high solubilization fraction) if the solubilized Fe exceeded a “solubility” limit (either as truly soluble Fe-organic complexes or as Fe colloids). Hand et al. (in press) suggest that the long transit times (8 –14 days to Midway compared to 7 days from North Africa to Barbados, Jickells and Spokes, 2001) and atmospheric chemical reactions may be responsible for a greater degree of solubilization of Asian dust. Meskhidze et al. (2003) observed high concentrations of free gaseous HNO3 in an Asian dust storm that had traversed an urban area. They pointed out that this free HNO3 could only exist if dust particles were moistened with very low pH solutions, and they hypothesized that Asian acid urban pollution (by sulfuric and nitric acid) could increase the solubilization of iron from Asian desert dusts that pass over urban areas. It is possible that the high apparent dust solubilizations required to account for surface Fe at HOTALOHA may be caused in part by this urban acidification. If this effect is significant, it may also be that economic growth in China during the past three decades has resulted in higher iron solubilization. Coal production in China doubled between 1981

and 1994 (http://www.chinaenvironment.com/english/channel/ energy/coal.html). These data are consistent with the estimate by Prospero et al. (2003) that anthropogenic nitrate and sulfate at Midway Island doubled in the same period. It is also consistent with estimates of Chinese NOx emissions associated with fossil fuel burning (Zheng et al., 2002) indicating that from 1960 to 1977, NOx emissions increased from 0.5 Tg N yr⫺1 to 3.6 Tg N yr⫺1 and then increased again to 6.8 Tg N yr⫺1 by 1997. If a pollution-induced shift in iron availability has significantly altered the flux of Fe to the North Pacific ocean, it may also play a factor in the decadal ecological and chemical shifts in North Pacific surface waters (Karl et al., 1997). Substantially different treatments have been used to estimate the percentage of dust solubilization in North Atlantic and North Pacific aerosols, hence it is difficult to compare experiments between different datasets. Zhuang et al. (1990) placed North Pacific Island aerosols in surface seawater and found ⬃50% release of Fe after 1 h for their lowest aerosol Fe/seawater ratio (2 nmol kg⫺1; higher aerosol/seawater ratios showed much lower release percentages). Zhuang et al. (1992a) performed experiments with aerosols collected at

944

E. A. Boyle et al. Table 3. HALE-ALOHA MITESS Fe time-series samples.

Sampler

Date

Fe (nmol kg⫺1)

s

n

A1 B1 A2 B2 A3 B3 A4 B4 B5 A6 B6 B7 A8 B8 A10 B10

06/01/99 06/01/99 06/16/99 06/16/99 07/01/99 07/01/99 07/16/99 07/16/99 07/31/99 08/15/99 08/15/99 08/30/99 09/14/99 09/14/99 10/14/99 10/14/99

1.84 ?2.66 1.62 1.34 0.95 0.81 0.77 0.81 0.77 1.11 1.03 0.78 1.86 0.79 ?1.28 0.69

0.02 0.02 0.03 0.01 0.06 0.08 0.02 0.12 0.06 0.11 0.07 0.02 0.02 0.03 0.07 0.20

2 2 3 2 3 3 2 2 2 3 3 3 3 2 3 3

Samples are collected just after 00:00 AM.

four remote Pacific island stations. They stated that 2N acidic solutions solubilized more than 90% of the aerosol Fe. Hand et al. (in press) found that 1.7 ⫾ 0.8% Fe(II) was released into pH 4.5 formic acid solution for ⬍2.5 ␮m North Pacific aerosol samples collected in April 2001. Bergquist (2004) treated two fine-fraction North Pacific aerosol samples (April 2001) with surface seawater (refreshed daily) for 3 days and found 7% and 37% release of ⬍0.4 ␮m Fe. Zhuang et al. (1992b) reported that ⬃15% of total Fe was present as Fe(II) in Barbados aerosols (after correction for a typographic error noted by Zhu et al., 1997). However, Hand et al. (in press) found that 7 ⫾ 13% Fe(II) was released into pH 4.5 formic acid solution for ⬍2.5 ␮m western tropical Atlantic aerosol samples collected in January 2001. Zhu et al., (1997) reported that fresh Barbados aerosols only released 6% of the total Fe from the dust when leached with pH ⫽ 1 sodium chloride solution for 5 min. Spokes and

Fig. 11. Surface water Fe transects near Hawaii. Note that Johnson et al. (2003) data are “reactive” Fe (Fe made reactive from an unfiltered sample after brief mild acidification). MP5 data are total Fe from filtered samples (⬍0.4 and ⬍0.02 ␮m). Also note that Johnson et al. (2003) data plotted here are the average of numerous samples within a latitude band, whereas our MP5 data are for individual samples.

Jickells (1996) also found a relatively low dissolution percentage from stored Sahara aerosols (⬃2.5%) with only 40% of that occurring as Fe(II) after exposure to sunlight when ˜ 2, pH ˜ 5, treated with a pH cycle (24 h distilled water, 24 h pH ˜ 8). Bonnet and Guieu (2004) placed Sahara soils in then pH contact with uncontaminated surface seawater and found that ⬃2% of the Fe was released after 7 days for the experiments with the lowest soil/seawaters ratios. At present, there is little direct experimental evidence to indicate high percentages of Fe release from Pacific aerosols (and the one study that included aerosols from both oceans specifically did not find higher Pacific aerosol Fe release: Hand et al., in press). Nonetheless, the surface water data presented by Johnson et al. (2003) and that presented here seem to require high release percentages than observed in the dust release experiments. It is crucial to obtain a better understanding of the release of aerosol Fe in the North Pacific to account for the oceanic distribution of Fe and the biogeochemistry of the ocean. 3.4. Fe: Spatial Variations in Surface Waters Along Transects 174°E-123°E and From Honolulu to HOT

Fig. 10. HALE-ALOHA total Fe, April–June 1999.

Johnson et al. (2003) obtained surface water Fe transects from Hawaii to California showing the highest “reactive” Fe concentrations near Hawaii in March and May 2001 (“reactive” Fe refers to the Fe that is detected by their analytical method after unfiltered samples are very briefly acidified). During our July 2002 cruise, we obtained several surface samples going northwest from Hawaii to 173.5°W (see Fig. 1) that extends

Fe, Pb, and Mn variability in central North Pacific (HOT)

945

Fig. 12. Surface water Fe transect from Honolulu to HOT-ALOHA, November 7– 8, 1998, with total, ⬍0.4 ␮m, and ⬍0.02 ␮m Fe size fractions.

this spatial perspective (Fig. 11). Westward from HOTALOHA in July 2002, ⱕ0.4 ␮m Fe decreases from 0.47 nmol kg⫺1 to somewhat lower values (0.3– 0.4 nmol kg⫺1) west of 164°W, seemingly as if higher Fe concentrations were either connected with proximity to Hawaii or to more southerly latitudes. We are skeptical that proximity to the Hawaiian Islands could be related to the higher Fe near HOT-ALOHA, because the station was chosen to be normally upwind of the islands (Karl and Lucas, 1996). Furthermore, a near-shore surface water Fe transect out of Hawaii does not show any island influence beyond 30 –50 km: on November 7– 8, 1998 (pre-HOT 99) we used an underway sampler to collect nearsurface samples as the R/V Moana Wave steamed from Honolulu to HOT-ALOHA (Fig. 12). Up to 70 km away from Honolulu, total Fe is higher than at HOT-ALOHA. Our first ⱕ0.4 ␮m Fe concentration (31 km north of Honolulu) is slightly elevated, but from 38 to 100 km north of Honolulu, ⬍0.4 ␮m Fe concentrations are consistently 0.5 ⫾ 0.1 nmol kg⫺1. Seven-day air mass back-trajectories from Station HOTALOHA for April 10 –22, 2001 indicate that the air at HOTALOHA typically originates further north at the latitudes of higher dust transport, sometimes from higher levels in the atmosphere where dust concentrations are higher (Fig. 13). Hence it is plausible that dust could be transported to this site either by descending air trajectories from the northern dust plume, or by lower level transport of dust that settles out of the higher level dust plume, or wet deposition processes which

dominate deposition in this region in most model simulations (e.g., Luo et al., 2003). Johnson et al. (2003) compared their transect Fe data to the surface soil aerosol concentration calculated with the Navy Aerosol Analysis and Prediction System (NAAPS) model from March 1 to May 31, 2001. The NAAPS model estimated higher aerosol concentrations near the northeast end of the transect; Johnson et al. (2003) noted that it was puzzling that higher Fe levels are found closer to HOT-ALOHA despite higher estimated aerosol concentrations to the northeast. Calculated dust fluxes from the model of Luo et al. (2003) for the same period March-April-May, (MAM) are shown in Figure 14. These differences are likely to be due to differences in the models, since aerosol concentrations in Luo et al. (2003) for MAM, 2001 look more similar to the dust fluxes than to the aerosol concentrations from the NAAPS model. There is less contrast between the Luo et al. (2003) model aerosol flux and the Johnson et al. (2003) surface water Fe variability, with minima in the middle of the transect (and the lower Fe values near the California Coast due to higher productivity in the California Current). None of these considerations account for the slight decrease in ⬍0.4 ␮m Fe concentrations towards the northwest of Hawaii during July 2002. Also, in contrast to ⬍0.4 ␮m Fe, ⬍0.02 ␮m Fe shows no apparent variation from Hawaii to 173.5°W. These data indicate that the variations in ⬍0.4 ␮m Fe are due to variations in the concentration of colloidal Fe.

946

E. A. Boyle et al.

Fig. 13. Seven-day back trajectories (above) from station ALOHA calculated using NOAA HYSPLIT program using NCAR Global Reanalysis (1948 –2002; url:http://www.arl.noaa.gov/ready/hysplit4.html). Numbers are placed near the beginning of the trajectory that will arrive at the surface of ALOHA on the date in April 2001 (midnight, UCT) indicated by that number.

Fe, Pb, and Mn variability in central North Pacific (HOT)

947

Fig. 14. Seasonal dust fluxes for March 1, 2001 to May 31, 2001 as estimated by the model of Luo et al. (2003).

3.5. Pb: Decadal Evolution of Pacific Profiles, 1976 –1999 Lead (Pb) in the ocean is dominantly anthropogenic, and its oceanic distribution is evolving on decadal time scales in response to changing Pb emissions and as ocean surface waters penetrate into the ocean interior. The first reliable Pb profile in the ocean was collected by Schaule and Patterson (1981) from a station at 32°41=N, 145°00=W occupied in 1976. This site is located 911 km northeast of HOT-ALOHA. Apart from profiles collected at 39.6°N, 140.8°W and farther north in 1987 (Martin et al., 1989), no other central North Pacific Pb profiles have been reported. To determine how Pb in the central North Pacific has evolved during the past 23 yr, we measured Pb on unfiltered samples from HOT-ALOHA occupation 105, May 11–12, 1999 (Table 4). We chose to use unfiltered samples because Pb is dominantly in the dissolved form (⬎90%) throughout most of the ocean (Bacon et al., 1976) and by eliminating filtration we remove a possible source of contamination. These analyses are compared to the 1976 data of Schaule and Patterson (1981) in Figure 15. Twenty-three years after the Schaule and Patterson profile, Pb concentrations have decreased in near-surface waters; at depths ⱖ200 m, thermocline and deep water concentrations remain largely unchanged. Surface water Pb concentrations in the central North Pacific have decreased by nearly a factor of 2, from 65 pmol kg⫺1 down to 35 pmol kg⫺1, as previously reported by Bell et al. (2002). Lower Pb concentrations in 1997

compared to 1976 are observed in samples from 0 to 102 m; the sample at 192 m has essentially the same Pb concentration seen two decades ago. Because the mixed layer at this site deepens Table 4. HOT 105 Pb profile, May 7– 8, 1999. Depth 0 0 10 35 70 150 192 428 452 570 686 730 1118 1405 1724 2508 3000 3500 4000

T (°C)

est. S (ppt)

est. O2 (nmol kg⫺1)

24.11 22.773 22.799 21.492 16.753 8.447 7.727 5.810 4.771 4.636 3.651 3.002 2.501 1.719 1.518 1.453 1.442

35.372 35.322 34.743 34.070 34.051 34.072 34.230 34.291 34.512 34.564 34.603 34.656 34.679

221.7 216.3 202.9 161.4 146.2 80.9 29.1 25.3 46.8 60.1 73.7 104.6 127.3

34.691

147.3

Pb (pmol kg⫺1)

␴

n

?43.3 35.2 36.6 37.7 40.1 47.8 63.3 60.9 52.4 46.3 10.7 25.5 15.7 15.0 10.6 8.4 8.4 8.3 5.6

0.9 1.8 1.3 1.5 3.1 1.1 0.2 2.4 1.1 0.8 2.1 0.3 0.2 0.3 2.5 0.1 6.3 1.1 1.0

3 2 3 2 3 3 2 3 2 3 4 2 2 2 2 2 2 2 3

? Sample rejected as contaminated based on duplicate sample (below).

948

E. A. Boyle et al.

thermocline in the central North Pacific is in contrast to observations in the western North Atlantic (Wu and Boyle, 1997b) and remains a puzzle yet to be explained. 3.6. Mixed Layer Pb Variability at HALE-ALOHA, 1997–1999

Fig. 15. Pb profiles near Hawaii from 1976 (Schaule and Patterson, 1981) and May 1999 (this paper).

to ⬎100 m in midwinter, Pb concentrations in the upper 100 m reflect the balance between the eolian supply of anthropogenic Pb and biogenic uptake and scavenging onto sinking particles. Given a Pb residence time of ⬃2 yr in oligotrophic surface waters of the central North Pacific (Nozaki et al., 1976), it is expected that Pb concentrations in surface waters will closely track changes in anthropogenic emissions with a lag of only a few years. The apparent absence of changes in Pb concentrations below the upper layer is surprising. This depth range is dominated by the physical circulation of the “ventilated thermocline” where surface waters are subducted into the thermocline and move towards the southwest (Fine et al., 2001; Luyten et al., 1983). Fine et al. (2001) estimated the chlorofluorocarbon (CFC) 11/12 “ages” of waters near Hawaii as ⬃2.5 yr at ⬃200 m, ⬃12 yr at ⬃350 m, ⬃20 yr at ⬃550 m, and ⬃31 yr at ⬃800 m. Assuming that the cooler surface waters to the north show the same decrease in Pb as the waters near Hawaii, we would expect water at 200 m to show less Pb in 1999 compared to 1976, and at least a slight decrease in Pb below that depth. This discrepancy with expectation may in part be due to the distance between the stations (911 km), so that we are not comparing the same site. A Pb profile was collected 250 km to the northwest of the S&P station (39.6°N, 140°W) in 1987 (Martin et al., 1989); upper thermocline Pb concentrations are slightly lower than seen at Schaule and Patterson’s site in 1976 and at HOT in 1997. This difference between Pb and CFCs may be due in part to the known inherent bias of CFC ages towards low values (because of the exponential increase of CFCs for most of the latter half of the twentieth century, as high-CFC waters mix with low-CFC waters, they dominate the “age”). Nonetheless, the lack of an apparent significant reduction of Pb in the upper

All of the surface water Pb concentrations from 1997 to 1999 are lower than the 1976 surface water concentrations observed nearby and farther north by Schaule and Patterson (1981). Their sample collected nearest Hawaii (24°19=N 154°29=W) had 64 pmol kg⫺1; the average of this sample and four others between that site and 33°03=N 140°29=W was 65 pmol kg⫺1 (with a range from 60 to 72 pmol kg⫺1). The average of our 101 samples is 35 pmol kg⫺1 with a standard deviation of 6 pmol kg⫺1 and a range of 25–57 pmol kg⫺1 (Table 5, Fig. 9). We attribute the decrease in surface water Pb to the elimination of leaded gasoline in Japan and to some extent the United States and Canada. Mexico (1998) and China (1999) have begun phasing out leaded gasoline within the past decade, but cars in those countries dominantly used leaded gasoline during our sampling period. Gasoline consumption by Mexico (6% of U.S. gasoline consumption in 1992–94) and China (9%) are small by comparison to Pb gasoline utilization by the U.S., Canada (7%), and Japan (10%) (gasoline consumption statistics are from Thomas, 1995). Although central North Pacific surface Pb concentrations have decreased by a factor of 2 during the past two decades, this decrease is significantly less than the factor of 4 –5 decrease observed near Bermuda (Wu and Boyle, 1997b). In the late 1990s, the Pb concentrations of surface waters near Bermuda and Hawaii are essentially identical. We attribute the difference in Pb surface water concentration reduction to the prevailing westerlies, where winds from the U.S. dominate the North Atlantic and the same westerlies tend to minimize the transport of emissions from the U.S. into the Pacific (although some air mass trajectories take west coast U.S. air over the eastern boundary of the Pacific Ocean). The North Pacific has twice the surface area as the North Atlantic; given similar surface Pb residence times in the two basins (Bacon et al., 1976; Nozaki et al., 1976), then total emissions to the North Pacific in the late 1990s must be about a factor of 2 higher than for the North Atlantic. Leaded gasoline has been completely phased out in the U.S., Canada, and Japan, and gasoline consumption is relatively low in the Far East and Oceania, so we believe that other sources of Pb must be more important than gasoline Pb. Coal production in China doubled from 1981 to 1994 (see above). We suspect that coal combustion and other high-temperature industrial activities in China, Japan, and Korea are likely to be the dominant sources of Pb to the North Pacific in the late 1990s. This conclusion appears to be consistent with the 1989 Pb emission patterns mapped by Pacyna et al. (1995). The time-series data indicate that there is a seasonal cycle in the mixed layer Pb concentration of the central North Pacific (Fig. 16). Pb concentrations are relatively low from May through November (⬃33 pmol kg⫺1) and then rise to ⬃40 pmol kg⫺1 in December through February (and perhaps March). In these years, there were annual cycles in primary productivity and carbon flux at the base of the mixed layer with

Fe, Pb, and Mn variability in central North Pacific (HOT)

949

Table 5. HALE-ALOHA MITESS Pb time series. Date

Sampler

Pb (pmol kg⫺1)

Date

Sampler

Pb (pmol kg⫺1)

05/20/97 05/29/97 06/12/97 07/03/97 07/11/97 07/17/97 07/31/97 08/07/97 08/14/97 08/28/97 09/04/97 09/11/97 09/18/97 09/25/97 10/09/97 10/16/97 10/23/97 10/30/97 11/06/97 12/02/97 12/24/97 12/29/97 01/03/98 01/08/98 01/13/98 01/18/98 01/23/98 01/28/98 02/02/98 02/12/98 02/17/98 02/22/98 03/09/98 03/19/98 04/08/98 04/13/98 04/18/98 04/18/98 04/18/98 04/18/98 05/02/98 05/02/98 05/18/98 06/03/98 06/19/98 07/05/98 07/21/98 08/06/98 08/06/98 08/09/98 09/23/98

ATE 5/20/97 A1 A2 B3 ATE HOT85 B4 B5 A6 B6 B7 A8 B8 A9 B9 B10 A11 B11 A12 B12 MWB (pole) A1 B1 A2 B2 A3 B3 A4 B4 A5 A6 B6 A7 B8 B9 B11 A12 B12 MW10 (pole) MW4 (pole) MW2 (pole) A1 B1 A2 B3 B4 A5 B6 A7 B7 ATE HOT96 B10

31.3 32.6 29.4 30.9 32.0 38.3 31.7 32.9 30.0 30.2 32.1 31.2 26.8 35.1 37.1 33.9 33.4 33.4 33.6 39.4 47.1 34.9 42.2 33.3 44.7 39.7 50.4 34.9 35.4 54.8 36.9 39.4 38.7 33.2 32.2 39.2 35.5 27.5 30.7 27.4

10/09/98 10/25/98 10/25/98 11/07/98 11/07/98 11/20/98 11/20/98 11/28/98 12/05/98 12/13/98 12/21/98 01/06/99 01/14/99 01/22/99 02/02/99 02/07/99 02/15/99 02/23/99 03/11/99 03/18/99 03/19/99 04/06/99 04/15/99 04/24/99 05/02/99 05/09/99 06/01/99 06/01/99 06/16/99 06/16/99 07/01/99 07/01/99 07/16/99 07/16/99 07/31/99 08/15/99 08/15/99 08/30/99 09/14/99 09/14/99 09/29/99 10/14/99 10/14/99 11/15/99 11/15/99 11/30/99 11/30/99 12/15/99 12/30/99 01/14/00 01/14/00

B11 A12 B12 MW (pole) PMP bottle MW (pole) HDPE bottle A1 B1 A2 B2 A3 A4 B5 A6 B6 ATE HOT102 B7 A8 B8 B9 ATE HOT103 A10 A11 ATE HOT104 A12 B12 ATE/vane A1 B1 A2 B2 A3 B3 A4 B4 B5 A6 B6 B7 A8 B8 A9 A10 B10 A1 B1 A2 A3 B3 A4 A5 B7

33.4 32.2 30.3 29.8 28.6 28.3 33.3 33.6 30.0 37.1 35.5 31.0 30.1 28.2 28.2 35.5 34.4 32.1 31.8 29.7 32.2 31.0 29.6

peaks in the summer (Karl et al., 2002). Although midyear productivity maxima may account for the Pb drawdown in spring and summer, this process cannot account for the Pb increase in winter months. This Pb increase does not coincide with the Asian terrigenous dust cycle to MLO (high dust from March through May), as discussed earlier. However, the wintertime increase in Pb concentrations does coincide with the deepest mixed layer depths caused by winter extra-tropical cyclones. North Pacific midlatitude extra-tropical cyclones move from west to east approximately every 5–7 days during winter. Cold fronts associated with these storms sometimes reach Hawaii, resulting in several days of strong northerly and northeasterly winds. The strong winds force mixing in the

upper ocean and deepen the mixed layer (Karl and Lucas, 1996). Perhaps these events also deposit more Asian Pb near Hawaii. However, we think that it is more likely that the additional Pb comes from “mining” higher Pb concentrations in the upper thermocline. Pb increases strongly below the mixed layer: in our May 1999 profile, Pb increases from 41 pmol kg⫺1 at 70 m to 52 pmol kg⫺1 at 102 m to 65 pmol kg⫺1 at 192 m. The seasonal mixed layer winter increase of ⬃7 pmol kg⫺1 could be accounted for by a 22% admixture of water from 192 m. If this interpretation is correct, this seasonal cycle is a feature of the past decade or two, and may not occur in the future when the thermocline Pb maximum eventually is eroded away by ventilation and scavenging.

950

E. A. Boyle et al.

and higher than typically considered as limiting to marine phytoplankton growth. Fe also shows high values and temporal variability at depths between 1000 and 1500 m; high ␦3He data in this depth zone indicate that this high Fe (mainly particulate) and its variability derives from hydrothermal emanations from the nearby Loihi seamount hydrothermal fields and their interaction with intermediate water mesoscale circulation variability. In near-surface waters, Pb shows 20% higher concentrations during the winter months, probably because of entrainment of higher-Pb upper thermocline Pb as the mixed layer deepens in winter. During the 23-yr period from 1976 to 1999, Pb decreased by a factor of 2 in the mixed layer but showed no significant decreases below 200 m. We attribute the decrease in surface Pb to the phasing out of leaded gasoline by the U.S., Japan, and Canada, and we conclude that the current surface water Pb must be supported by Asian emission sources.

Fig. 16. Monthly averaged Pb data from moored sampler (MITESS) on the HALE-ALOHA mooring, 1997–1999.

Interannual and shorter term variability are also evident in these data. Summer 1998 has lower Pb levels than the summers of 1997 or 1999. Note that December 1999 and January 2000 do not have the elevated winter Pb levels evident in the previous years. This result is consistent with our hypothesis of “thermocline mining” because there the mixed layer had not deepened yet that year. The short-term scatter (weekly and bimonthly compared to seasonal trends) is also in excess of the measurement error. We suspect that this residual variability is caused by patchiness in Pb deposition and biological uptake, with perhaps an element of mesoscale circulation variability.

Acknowledgments—This research was funded by NSF grants OCE0002273 and OCE-9981442. We thank Dave Karl, Terry Houlihan, and Louie Tupas for accommodating us on HOT cruises and the HALE/ ALOHA moorings. We thank Tony Michaels, Doug Capone, Ron Siefert, Michael Neumann, and our other colleagues in the nitrogen fixation Biocomplexity project MANTRA for help at sea and helpful discussions. We thank Chris Measures for the use of his lab during MITESS turnarounds. Thanks to Rob Sherrell and Paul Field for allowing us to analyze some of these samples on the Rutgers ICPMS. The R/V Moana Wave cruises were carried out as part of Dave Karl’s ongoing NSF-sponsored time series work at HOT-ALOHA; the other cruises were done as part of an NSF-sponsored Biocomplexity program (PI: A. Michaels). We thank the officers and crews of the R/V Moana Wave, R/V Wecoma, R/V Kaimikai-O-Kanaloa, and R/V Kilo Moana for their efforts on our behalf. The authors gratefully acknowledge the NOAA Air Resources Laboratory (ARL) for provision of the HYSPLIT transport and dispersion models used in this publication. We thank John Edmond for discussions, and we thank the three reviewers for suggestions that improved the clarity of the manuscript. Associate editor: T. Lowenstein

4. CONCLUSIONS

These data provide the first look at Pb near Hawaii since the original observations of Schaule and Patterson (1981) at a station occupied in 1976. These new data demonstrate the central North Pacific Ocean response to the phasing out of leaded gasoline by major industrial nations. These observations also reveal significant temporal variability for Fe and Pb in near-surface waters and for dissolved and particulate Fe in intermediate waters. Fe and Pb show significant temporal variability in nearsurface waters at station HOT-ALOHA in the central North Pacific. The highest observed Fe concentration was seen during the period of peak Asian dust transport, and the lowest observed Fe concentration was seen after more than 6 months of low dust transport and deepened mixed layer (January). Shorter-term variability is also significant, and typical HOT-ALOHA surface Fe concentrations are only a factor of 2–3 lower than typical North Atlantic concentrations at this latitude. In the face of an estimated 40-fold lower dust deposition near Hawaii compared to Barbados, the relatively high surface water Fe concentrations imply a much greater degree of solubilization of Fe from dust in the North Pacific compared to the North Atlantic. Fe concentrations in the HOT-ALOHA mixed layer are higher than typically observed deeper in the euphotic zone

REFERENCES Archer D. E. and Johnson K. (2000) A model of the iron cycle in the ocean. Glob. Biogeochem. Cyc. 14, 269 –279. Bacon M. P., Spencer D. W., and Brewer P. G. (1976) 210Pb/226Ra and 210 Po/210Pb disequilibria in seawater and suspended particulate matter. Earth Planet. Sci. Lett. 32, 277–296. Bell J., Betts J., and Boyle E. (2002) MITESS: A moored in-situ trace element serial sampler for deep-sea moorings. Deep-Sea Res. I 49, 2103–2118. Bergquist B. A. (2004) The marine geochemistry of iron and iron isotopes. Ph.D. thesis, Massachusetts Institute of Technology/ Woods Hole Oceanographic Institution. Bishop J. K. B., Davis R. E., and Sherman J. T. (2002) Robotic observations of dust storm enhancement of carbon biomass in the North Pacific. Science 298, 817– 821. Blain S., Sedwick P. N., Griffiths F. B., Queguiner B., Bucciarelli E., Fiala M., Pondaven P., and Treguer P. (2002) Quantification of algal iron requirements in the Subantarctic Southern Ocean (Indian sector). Deep-Sea Res. II 49, 3255–3273. Bonnet S. and Guieu C. (2004) Dissolution of atmospheric iron in seawater. Geophys. Res. Lett. 31, L03303, doi:10.1029/ 2003GL018423. Boyd P. W., Watson A. J., Law C., Abraham E., Trull T., and Murdoch R. (1999) SOIREE: A Southern Ocean iron release experiment elevates phytoplankton stocks in polar waters. Trans. Am. Geophys. Union 80, OS30.

Fe, Pb, and Mn variability in central North Pacific (HOT) Bruland K. W. and Franks R. P. (1983) Mn, Ni, Cu, Zn, and Cd in the western North Atlantic. In Trace Metals in Seawater (eds. E. Boyle, C. S. Wong, K. W. Bruland, J. D. Burton, and E. D. Goldberg), pp. 395– 414. Plenum, New York. Bruland K., Orians K. J., and Cowen J. P. (1994) Reactive trace metals in the stratified central North Pacific. Geochim. Cosmochim. Acta 58, 3171–3182. 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., and Kudela R. (1996) A massive phytoplankton bloom induced by an ecosystemscale iron fertilization experiment in the equatorial Pacific ocean. Nature 383, 495–501. de Baar H. J. W., Jong J. T. M. d., Nolting R. F., Timmermans K. R., Leeuwe M. A. v., Bathmann U., Loeff M. R. v. d., and Sildam J. (1999) Low dissolved Fe and the absence of diatom blooms in the remote Pacific waters of the Southern Ocean. Mar. Chem. 66, 1–34. Falkowski P. G. (1997) Evolution of the nitrogen cycle and its influence on the biological sequestration of CO2 in the ocean. Nature 387, 272. Fine R. A., Maillet K. A., Sullivan K. F., and Willey D. (2001) Circulation and ventilation flux of the Pacific Ocean. J. Geophys. Res. 106, 22159 –22178. Flegal A. R. and Patterson C. C. (1983) Vertical concentration profiles of lead in the Central Pacific at 15N and 20S. Earth Planet. Sci. Lett. 64, 19 –32. Gao Y., Kaufman Y. J., Tanré D., Kolber D., and Falkowski P. G. (2001) Seasonal distributions of aeolian iron fluxes to the global ocean. Geophys. Res. Lett. 28, 29 –32. Gong S. L., Zhang X. Y., Zhao T. L., McKendry I. G., Jaffe D. A., and Lu N. M. (2003) Characterization of soil dust aerosol in China and its transport and distribution during 2001 ACE-Asia: 2. Model simulation and validation. J. Geophys. Res. 108, 4262,10.1029/ 2002JD002633. Hand J. L., Mahowald N. M., Chen Y., Siefer R. L., Luo C., Subramaniam A., and Fung I. (2004) Estimates of soluble iron from observations and a global mineral aerosol model: Biogeochemical implications. J. Geophys. Res. 109, D17205, doi 10.1029/ 2004JD004574. Hilton D. R., McMurtry G. M., and Goff F. (1998) Large variations in vent fluid CO2/3He ratios signal rapid changes in magma chemistry at Loihi seamount, Hawaii. Nature 396, 359 –362. Jickells T. D. and Spokes L. J. (2001) Atmospheric iron inputs to the ocean. In The Biogeochemistry of Iron in Seawater (eds. D. R. Turner and K. A. Hunter), pp. 85–122. John Wiley and Sons USA. Johnson K. J., Gordon R. M., and Coale K. H. (1997) What controls dissolved iron concentrations in the world ocean? Mar. Chem. 57, 137–161. Johnson K. S., Elrod V. A., Fitzwater S. E., Plant J. N., Chavez F. P., Tanner S. J., Gordon R. M., Westphal D. L., Perry K. D., Wu J., and Karl D. M. (2003) Surface ocean–lower atmosphere interactions in the Northeast Pacific Ocean gyre: Aerosols, iron and the ecosystem response. Glob. Biogeochem. Cycles 17, 1063, doi:10.1029/ 2002GB002004. Karl D. M. and Lucas R. (1996) The Hawaii Ocean Time-series (HOT) program: Background, rationale and field implementation. DeepSea Res. 43, 129 –146. Karl D., Leteller R., Tuypas L., Dore J., Christian J., and Hebel D. (1997) The role of nitrogen fixation in biogeochemical cycling in the subtropical North Pacific Ocean. Nature 388, 533–538. Karl D. M., Bates N. R., Emerson S., Harrison P. J., Jeandel C., Llinás O., Liu K.-K., Marty J.-C., Michaels A. F., Miquel J. C., Neuer S., Nojiri Y. and Wong C. S. (2002) Temporal studies of biogeochemical processes determined from ocean time-series observations during the JGOFS era. In Ocean Biogeochemistry: The Role of the Ocean Carbon Cycle in Global Change (ed. M. J. R. Fasham), pp. 239 –267. Springer-Verlag USA. Klinkhammer G. P. and Bender M. L. (1980) The distribution of manganese in the Pacific Ocean. Earth Planet. Sci. Lett. 46, 361– 384.

951

Kustka A., Carpenter E. J., and Sañudo-Wilhelmy S. A. (2002) Iron and marine nitrogen fixation: Progress and future directions. Res. Microbiol. 153, 255–262. Landing W. M. and Bruland K. W. (1987) The contrasting biogeochemistry of iron and manganese in the Pacific Ocean. Geochim. Cosmochim. Acta 51, 29 – 43. Lefevre N. and Watson A. J. (1999) Modeling the geochemical cycle of iron in the oceans and its impact on atmospheric CO2 concentrations. Glob. Biogeochem. Cycles 13, 727–736. Luo C., Mahowald N. M., and del Corral J. (2003) Sensitivity study of meteorological parameters on mineral aerosol mobilization, transport and distribution. J. Geophys. Res. 108, 4447, doi:10.1029/ 2003JD003483. Lupton J. (1998) Hydrothermal helium plumes in the Pacific Ocean. J. Geophys. Res. 103, 15853–15868. Luyten J. R., Pedlosky J., and Stommel H. (1983) The ventilated thermocline. J. Phys. Oceanogr. 13, 292–309. Mahowald N., Luo C., del Corral J., and Zender C. (2003) Interannual variability in atmospheric mineral aerosols from a 22-year model simulation and observational data. J. Geophys. Res. 108, 4352, 10.1029/2002JD002821. Martin J. H. and Fitzwater S. E. (1988) Iron deficiency limits phytoplankton growth in the north-east Pacific subarctic. Nature 331, 341–343. Martin J. H. and Gordon R. M. (1988) Northeast Pacific iron distributions in relation to phytoplankton productivity. Deep-Sea Res. 35, 177–196. Martin J. H., Coale K. H., Johnson K. S., Fitzwater S. E., Gordon R. M., Tanner S. J., Hunter C. N., Elrod V. A., Nowicki J. L., Coley T. L., Barber R. T., Lindley S., Watson A. J., Scoy K. V., Law C. S., Liddicoat M. I., Ling R., Stanton T., Stockel J., Collins C., Anderson A., Bidigare R., Ondrusek M., Latasa M., Millero F. J., Lee K., Yao W., Zhang J. Z., Friederich G., Sakamoto C., Chavez F., Buck K., Kolber Z., Greene R., Falkowski P., Chisholm S. W., Hoge F., Swift R., Yungel J., Turner S., Nightingale P., Hatton A., Liss P., and Tindale N. W. (1994) Testing the iron hypothesis in ecosystems of the equatorial Pacific ocean. Nature 371, 123–129. Martin J. H., Gordon R. M., Fitzwater S. E., and Broenkow W. W. (1989) VERTEX: Phytoplankton/iron studies in the gulf of Alaska. Deep-Sea Res. 36, 649 – 680. Measures C. I. and Vink S. (2001) Dissolved Fe in the upper waters of the Pacific sector of the Southern Ocean. Deep-Sea Res. II 48, 3913–3941. Meskhidze N., Chameides W. L., Nenes A., and Chen G. (2003) Iron mobilization in mineral dust: Can anthropogenic SO2 emissions affect ocean productivity? Geophys. Res. Lett. 30, 2085; doi: 10.10292003GL018035. Moore W. S, Paull C. K., and, Ussle W. (2001) Short-lived radium isotopes in the Hawaiian Margin: Evidence for large fluid fluxes through the Puna Ridge. Eos Trans. AGU, Fall Meet. Suppl. 82(47), abstract OS42E-10. Nozaki Y., Thomson J., and Turekian K. K. (1976) The distribution of Pb-210 and Po-210 in the surface waters of the Pacific Ocean. Earth Planet. Sci. Lett. 32, 304 –312. Pacyna J. M., Scholtz M. T., and Li Y.-F. (1995) Global budget of trace metal sources. Environ. Rev. 3, 145–159. Parekh P., Follows M., and Boyle E. (2004) Modeling the global iron cycle. Glob. Biogeochem. Cyc. 18, GB1002, doi:10.1029/ 2003GB002061. Parrington J. R., Zoller W. H., and Aras N. K. (1983) Asian dust: Seasonal transport to the Hawaiian Islands. Science 220, 195–198. Perry K. D., Cahill T. A., Schnell R. C., and Harris J. M. (1999) Long-range transport of anthropogenic aerosols to the National Oceanic and Atmospheric Administration baseline station at Mauna Loa Observatory, Hawaii. J. Geophys. Res. 104, 18521–18533. Prospero J. M., Savoie D. L., and Arimoto R. (2003) Long-term record of nss-sulfate and nitrate in aerosols on Midway Island, 1981–2000: Evidence of increased (now decreasing?) anthropogenic emissions from Asia. J. Geophys. Res. 108, 4019, doi:10.1029/ 2001JD001524. Rue E. L. and Bruland K. W. (1995) Complexation of iron(III) by natural ligands in the central North Pacific as determined by a new

952

E. A. Boyle et al.

competitive ligand equilibrium/adsorptive cathodic stripping voltammetry method. J. Mar. Chem. 50, 117–138. Schaule B. K. and Patterson C. C. (1981) Lead concentrations in the northeast Pacific: Evidence for global anthropogenic perturbations. Earth Planet. Sci. Lett. 54, 97–116. Schaule B. and Patterson C. C. (1983) Perturbations of the natural lead depth profile in the Sargasso Sea by industrial lead. In Trace Metals in Seawater (eds. C. S. Wong, E. Boyle, K. W. Bruland, J. D. Burton and E. D. Goldberg), pp. 487–504. Plenum, New York. Sedwick P. N., McMurtry G. M., and MacDougall J. D. (1992) Chemistry of hydrothermal solutions from Pele’s Vents, Loihi Seamount, Hawaii. Geochim. Cosmochim. Acta 56, 3643–3667. Sedwick P. N., DiTullio G. R., and Mackey D. J. (2000) Iron and manganese in the Ross Sea, Antarctica: Seasonal iron limitation in Antarctic shelf waters. J. Geophys. Res. 105, 11331–11336. Spokes L. J. and Jickells T. D. (1996) Factors controlling the solubility of aerosol trace metals in the atmosphere and on mixing into seawater. Aquat. Geochem. 1, 355–374. Thomas V. M. (1995) The elimination of lead in gasoline. Annu. Rev. Energy Environ. 20, 301–324. Timmermans K. R., VanderWagt B., and de Baar H. J. W. Growth rates, half saturation constants, and silicate, nitrate and phosphate depletion in relation to iron availability of four large open ocean diatoms from the Southern Ocean. Limnol. Oceanogr.(in press). Tupas L., Santiago-Mandujano F., Hebel D, Nosse C., Fujieki L., Lukas R., Karl D., Winn C., Bidigare R., Landry M., and, Firing E. (1997–2000) Hawaii Ocean Time-series Data Reports 9 –12: 1997– 2000. University of Hawaii. Vink S., Boyle E. A., Measures C. I., and Yuan J. (2000) Automated high resolution determination of the trace elements iron and aluminium in the surface ocean using a towed fish coupled to flow injection analysis. Deep-Sea Res. I 47, 1141–1156. Watson A. J., Law C. S., Scoy K. A. V., Millero F. J., Yao W., Friederich G. E., Liddicoat M. I., Wanninkhof R. H., Barber R. T., and Coale K. H. (1994) Minimal effect of iron fertilization on sea-surface carbon dioxide concentrations. Nature 371, 143–145.

Watson A. J, Bakker D. C. E, Ridgewell A. J, Boyd P. W., and Law C. S (2000) Effect of iron supply on Southern Ocean CO2 uptake and implications for glacial atmospheric CO2. Nature 407, 730 –733. Wheat C. G., Jannasch H. W., Plant J. N., Moyer C. L., Sansone F. J., and McMurtry G. M. (2000) Continuous sampling of hydrothermal fluids from Loihi Seamount after the 1996 event. J. Geophys. Res. 105, 19353–19367. Wu J. and Boyle E. A. (1997a) Low blank preconcentration technique for the determination of lead, copper and cadmium in small-volume samples by isotope dilution ICPMS. J. Anal. Chem. 69, 2464 –2470. Wu J. F. and Boyle E. A. (1997b) Lead in the western North Atlantic Ocean: Completed response to leaded gasoline phaseout. Geochim. Cosmochim. Acta 61, 3279 –3283. Wu J. and Boyle E. A. (1998) Determination of iron in seawater by high-resolution isotope dilution by high-resolution inductively coupled plasma mass spectrometry after Mg(OH)2 coprecipitation. J. Anal. Chim. Acta 367, 183–191. Wu J., Boyle E., Wen L.-S., and Sunda W. (2001) Occurrence of soluble and colloidal iron in oligotrophic oceans. Science 293, 847– 849. Yuan J., Measures C. I., and Wheat C. G. (1994) Iron anomalies at Loihi Seamount and Puna Ridge. J. Trans. Am. Geophys. Union 75, 313. Zheng X., Fu C., Xu X., Yan X., Huang Y., Han S., Hu F., and Chen G. (2002) The Asian nitrogen cycle case study. Ambio 31, 79 – 87. Zhu X. R., Prospero J. M., and Millero F. J. (1997) Diel variability of soluble Fe (II) and soluble total Fe in North African dust in the trade winds at Barbados. J. Geophys. Res. 102, 21297–21305. Zhuang G., Duce R. A., and Kester D. R. (1990) The dissolution of atmospheric iron in surface seawater of the open ocean. J. Geophys. Res. 95, 16207–16216. Zhuang G., Yi Z., Duce R. A., and Brown P. R. (1992a) Chemistry of iron in marine aerosols. Glob. Biogeochem. Cyc. 6, 161–173. Zhuang G., Yi Z., Duce R. A., and Brown P. R. (1992b) Link between iron and sulphur cycles suggested by detection of Fe(II) in remote marine aerosols. Nature 355, 537–539.