Understanding the export of biogenic particles in oceanic waters: Is there consensus?

Progress in Oceanography Progress in Oceanography 72 (2007) 276–312 www.elsevier.com/locate/pocean

Review

Understanding the export of biogenic particles in oceanic waters: Is there consensus? P.W. Boyd

a,*

, T.W. Trull

b

a

b

NIWA Centre for Chemical and Physical Oceanography, Department of Chemistry, University of Otago, P.O. Box 56, Dunedin, New Zealand CSIRO Marine Research, University of Tasmania, and the Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart, Tasmania, Australia Received 31 May 2005; received in revised form 2 October 2006; accepted 3 October 2006 Available online 30 November 2006

Abstract We examine progress towards a global view of oceanic export of particulate organic carbon (POC) and other nutrient elements (P, N, Si) from the surface (upper 100 m), through the subsurface, to the deep sea (>1000 m), focusing on syntheses published since 1999 and on the Joint Global Ocean Flux Study. Food-web structure is important, and surface and subsurface processes contribute similarly to determine the fraction of net primary production (NPP) reaching the deep sea. NPP by large cells generally favours high surface export of POC. Preferential remineralization of P and N (versus C) with depth is common, as is regional variation in subsurface POC flux attenuation. The role of mineral fluxes is complex. Annual mean fluxes of POC and minerals are correlated in global deep sediment trap records, but causality and the relative importance of different minerals depends on the assumptions made. Time-series observations at single sites can oppose the geographic trends, and their large seasonal variability in the contribution of POC to total flux is at odds with mechanistic models for POC transport by minerals. Despite generally positive correlations between biogenic carbonate and POC fluxes, the overall role of carbonate export is to decrease the transfer of carbon dioxide from the atmosphere to the ocean. Both autotrophs and heterotrophs produce minerals, and progress in separating these contributions is required for the deconvolution of mineral ballast and food-web effects. Many recent models suggest global surface POC export of 10 GTC/yr, despite widely varying biological complexity. This limits the usefulness of their prediction of ecosystem and carbon cycle responses to global change. Progress requires better observations for model validation, and more efforts to relate the models to the observed complexity, rather than to overly simplified global syntheses. We advocate more time-series stations targeting under-studied biogeochemical regions, development of automated in situ tools for study of the subsurface ocean, and increased emphasis on combining ecological and biogeochemical methods. Ó 2006 Elsevier Ltd. All rights reserved. Keywords: Biogenic particles vertical carbon export; Ocean biogeochemistry

*

Corresponding author. Tel.: +64 3 479 5249; fax: +64 3 479 5248. E-mail address: [email protected] (P.W. Boyd).

0079-6611/$ - see front matter Ó 2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.pocean.2006.10.007

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

277

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Question 1. Has there been a paradigm shift following JGOFS? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Question 2. What is the relative importance of the surface and subsurface ocean in the control of the biological pump? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1. Oceanic processes affecting export. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. Vertical attenuation of sinking particle fluxes in the subsurface ocean . . . . . . . . . . . . . . . . . . . . . . . . 3.3. Sinking particle degradation and elemental stoichiometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Question 3. What is the biogeochemical role of mineral particles? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1. Minerals and the enhancement of POC export . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2. Life cycles of calcifiers and silicifiers and their influence on POC export . . . . . . . . . . . . . . . . . . . . . . . 4.3. Overall influence of carbonate export on atmosphere–ocean CO2 partitioning. . . . . . . . . . . . . . . . . . . Question 4. What insights have we gained from regional comparisons?. . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. Regional case-studies – the Southern Ocean. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2. Regional case-studies – the Northeastern Atlantic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3. Comparing global maps of export properties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Question 5. How far can we simplify the parameterization of export in global circulation models? . . . . . . 6.1. Comparison of export from models of varying biological complexity . . . . . . . . . . . . . . . . . . . . . . . . . 6.2. Regional trends in export from two models with differing biological complexity . . . . . . . . . . . . . . . . . 6.3. How robust is the validation of global export models? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Question 6. Can we predict future changes in downward biogenic export with any certainty? . . . . . . . . . . 7.1. How well do we understand the present controls on export? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.2. How are ocean properties predicted to change and how will they impact export? . . . . . . . . . . . . . . . . 7.3. How much might export fluxes change – can we set biogeochemical bounds? . . . . . . . . . . . . . . . . . . . Is there consensus?. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Recommendations for future work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

277 278

Contents 1. 2. 3.

4.

5.

6.

7.

8. 9.

281 281 282 284 285 287 291 292 293 293 294 296 297 297 298 300 301 301 302 302 303 304 305 305

1. Introduction The downward export of biogenic particles, colloquially known as ‘‘the biological pump’’, redistributes carbon and nutrients in the ocean and plays a significant role in controlling atmospheric carbon dioxide (CO2) levels (Volk and Hoffert, 1985). For example, a stronger biological pump during the last glacial maximum appears to have contributed up to a 30 latm reduction in atmospheric CO2 concentrations (Sigman and Boyle, 2000). In the modern ocean, the biological pump transfers around 5–15 GTC/yr to the deep sea (Falkowski et al., 1998), which is returned as part of the much larger overturning circulation of dissolved inorganic carbon (100 GTC/yr; (e.g., Sarmiento et al., 1998)). These transfers are large in comparison to the 6 GTC/yr of anthropogenic CO2 emissions to the atmosphere, but only alteration of the biological- and circulation-driven transfers can influence the uptake of this additional flux by the ocean. The circulation-driven transfer responds directly to the increasing atmospheric CO2 partial pressures, and removes 1/3 to 1/2 of the annual atmospheric load (Sabine et al., 2004). Modelling experiments indicate that changes in the strength and efficiency of the biological pump are also important (e.g., Sarmiento et al., 1998), but they are much less well understood (Falkowski et al., 2000; Sigman and Boyle, 2000). The identification of the factors controlling the downward export of biogenic particles has received considerable attention for three decades (Berger et al., 1989; Suess, 1980; Tre´guer et al., 2003). In the last decade there have been significant advances in biogeochemical studies, including increased temporal and spatial coverage of both the surface and deep ocean, with much of this undertaken in a co-ordinated way via the Joint Global Ocean Flux Study (JGOFS). In the same period, weekly satellite 1–4 km resolution global maps of chlorophyll (SeaWiFS and MODIS) became widely available, and oceanic General Circulation Models

278

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

(GCM’s) increasingly incorporate rudimentary biology and/or biogeochemistry (Cox et al., 2000; Matear and Hirst, 1999; Sarmiento et al., 1998; Six and Maier-Reimer, 1996). Syntheses of each of these aspects of our improved knowledge has begun and includes global (e.g., Francois et al., 2002; Lutz et al., 2002; Dunne et al., 2005) and regional (e.g., Antia et al., 2001) syntheses of particle export e.g., global estimates of processes important to export such as NPP from satellite observations (e.g., Behrenfeld and Falkowski, 1997), comparisons among global models (Orr, 2002; Doney et al., 2004), and between models and biogeochemical observations (Laws et al., 2000; Moore et al., 2002a,b). These syntheses and modelling efforts at the end of JGOFS (Doney, 1999; Doney and Sarmiento, 1999) provide a unique opportunity to assess our present understanding of the biological pump and whether there is consensus regarding the key drivers of downward export of biogenic particles. Given, the potential for climate change to impact the functioning of the biological pump (Bopp et al., 2001; Feely et al., 2004), our ability to model downward particulate biogenic fluxes will be central to determining the role of oceanic biogeochemical feedbacks in influencing climate in the future. In this review we have focused our analysis on synthesis and modelling papers published since 1999, and primarily those based on the results of JGOFS, including both regional (Antia et al., 2001; Armstrong et al., 2002; Nelson et al., 2002) and global (Francois et al., 2002; Klaas and Archer, 2002; Laws et al., 2000; Lutz et al., 2002; Moore et al., 2002a,b) studies. Our emphasis is on vertical particle export in the open ocean. Particle transfers to the deep sea are mediated by large scale circulations, such as off-shelf transport (Barth et al., 2002; Hwang et al., 2004; Verity et al., 2002; Walsh, 1991) or downwelling (Moore et al., 2002a,b; Oschlies and Kahler, 2004) were beyond our scope and are not considered. To structure our review we addressed six questions: (1) Has there been a paradigm shift following JGOFS? (2) What are the relative importance of the surface and subsurface ocean in the control of the biological pump? (3) What is the biogeochemical role of mineral particles? (4) What insights have we gained from regional comparisons? (5) How far can we simplify the parameterization of export in global circulation models? (6) Can we predict future changes in downward biogenic export with any certainty? Throughout the paper, we have used the following arbitrary definitions: the surface and subsurface ocean are the 0–100 m and 100–1000 m depth strata, respectively. Surface export and deep export are the downward transfer of particles at 100 m and 1000 m depth, and NPP is column-integrated net primary production. 2. Question 1. Has there been a paradigm shift following JGOFS? We used the syntheses on export fluxes reported during the Dahlem Conference (Berger et al., 1989) to reference our understanding prior to JGOFS. At that time there was already recognition that the processes controlling particle export are diverse and complex (Fig. 1a and Table 1). Efforts to couple the direct study of particle fluxes with the ecosystem processes that control them and with their influence on dissolved nutrient and other elemental distributions were also well advanced (Fig. 1b). In our view, the main advances in the past decade have been (a) to expand the global coverage of flux studies; (b) to address some of the limitations of the tools used in their measurement; (c) to begin to combine the direct results of the flux studies with broader issues of ecosystem function. In our view there are no new paradigms, but important advances have been made on previously reported ideas, including: (i) Validation of the role that pelagic food-web structure plays in setting the export of biogenic particles from the seasonal mixed layer (Legendre and LeFevre, 1989; Michaels and Silver, 1988), in a wide range of biogeochemical provinces (Boyd and Newton, 1995, 1999; Karl et al., 1996; Moore et al., 2002b). Prior to this, the magnitude of NPP was thought to be the key determinant of particle export (e.g., Pace et al., 1987; Suess, 1980).

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

279

Processes net primary (particle) production in euphotic zone

Resp.

Prod.

biogenic mineral production

physical and biological particle aggregation

particle sinking

particle re-packaging

de-entrainment from the mixed layer

organic matter solubilisation and remineralization to dissolved inorganic nutrients

advective particle supply

remineralization in sediments

biogenic-mineral redissolution

carbonate lysocline

zooplankton migrations, consumption

carbonate compensation

Observations

air-sea fluxes of O2,CO2

surface-tethered and neutrally-buoyant free-drifting sediment traps, 234Th efficiencies

net, filter, bottle samples food-web studies

seasonal nutrient and 234Th depletion

suspended particle distributions from cameras and pumps bottom-moored sediment traps

seasonal production of suspended barite basin-scale nutrient distributions

230Th, 231Pa trap-efficiences

centenial

sediments

Horizons and Timescales air-sea exchange: euphotic zone

seasonal

summer mixed layer

annual winter mixed layer

mode waters watermass subduction and return to surface

sediments

decadal

intermediate waters

centenial deep waters

millennial

holocene Fig. 1. Processes, observations, and depth horizons important to export estimates. (a) The biological pump begins with photosynthesis and culminates in the transfer of sinking organic matter to the deep sea. Many physical and biological processes contribute to the overall efficiency. (b) Observations of export range from direct measurement of sinking particle fluxes to inferences of particle transfer based on food-web structure, air–sea gas exchange, nutrient distributions. (c) The penetration of sinking particles to depth competes with some aspects of ocean circulation (i.e., upwelling) to remove carbon from contact with the atmosphere on a wide range of timescales.

280

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Table 1 A summary of the key processes reported to drive particle production and destruction in the surface and the subsurface ocean (a) Surface ocean processes Particle production Algal aggregation Particle aggregation – heterogenous aggregates TEP production and – biological glues Ballasting of aggregates Algal community structure – large cells Particle destruction Bacterial solubilisation Disaggregation/faecal pellet production (b) Subsurface ocean processes Particle production Ballasting Ontogenetic migration/ passive C flux Particle destruction Disaggregation/faecal pellet production Bacterial solubilisation Preferential remineralization (C/N/Si etc.) Vertical migration/active transport

Seasonality

Regional distributions

Episodic blooms – spring and fall (Billett et al., 1983) Over annual cycle (Passow, 2004)

Mesotrophic and upwelling waters

Over annual cycle (Engel, 2004)

All regions

Over annual cycle, but also some seasonality (Schiebel, 2002) Episodic blooms – spring and fall (Boyd and Newton, 1995)

Tropical to temperate waters (calcifers) – tropical to polar waters (silicifiers) Mesotrophic and upwelling waters (Buesseler, 1998)

Over annual cycle – impacted by temperature (Pomeroy et al., 1991) Over annual cycle (Dilling and Alldredge, 2000)

All regions – decreases with decreasing water temperature (Rivkin and Legendre, 2001) All regions

Seasonality

Regions

Over annual cycle – some seasonality (Schiebel, 2002) Late summer (Bradford-Grieve et al., 2001)

Tropical to temperate waters (calcifers) – tropical to polar waters (silicifiers) Subpolar waters (Kobari et al., 2003)

Over annual cycle (Dilling and Alldredge, 2000) Over annual cycle – decreases with depth (Hoppe et al., 1993) Over annual cycle (Bidle et al., 2002; Denman and Pena, 1999 Over annual cycle

All regions

All regions

All regions – all regions – decreases with decreasing water temperature (Rivkin and Legendre, 2001) All regions All regions (Longhurst and Harrison, 1988)

(ii) Elucidation of the mechanisms driving algal aggregation (Alldredge and Jackson, 1995) which can deliver autotrophic carbon directly to the deep ocean (Billett et al., 1983). This finding tempered the previously recognized importance of faecal pellets as the main vector for carbon export (Small et al., 1983). (iii) Resurgent emphasis on the role of minerals in modulating organic carbon export (Ittekkot, 1993; Ittekkot et al., 1992; Keil et al., 1994), both by enhancing sinking rates via ballasting (e.g., Francois et al., 2002) and by providing protection from microbial decomposition for a proportion of Particulate Organic Carbon (POC) (Armstrong et al., 2002; Hedges et al., 2000). (iv) Acknowledgement that the role of Particulate Inorganic Carbon (PIC) in biological carbon sequestration is complex, because its formation increases surface ocean CO2 partial pressures, thereby reducing ocean uptake of atmospheric CO2 (Frankignoulle et al., 1994), while its high specific gravity may increase the sinking rates of biogenic aggregates (e.g., Francois et al., 2002; Klaas and Archer, 2002). (v) Greater emphasis on the importance of particle transformations at depth rather than only those that occur in the surface ocean, from (a) studies of variations of particle properties with depth, including physical characteristics (e.g., Logan et al., 1995), elemental composition (e.g., Nelson et al., 2002; Schneider et al., 2003), and organic geochemistry (e.g., Hedges et al., 2000; Lee et al., 2000); (b) evidence of the importance of heterotrophic bacteria as constituents of deep-sea particles (Lampitt et al., 1993) and of

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

281

their ability to alter particle elemental stoichiometry (Bidle and Azam, 1999; Bidle et al., 2002); and (c) models seeking to explain distributions of dissolved nutrients (Howard et al., 2006; Li and Peng, 2002; Schlitzer, 2004). (vi) Increasing awareness of the limitations of the global applicability of the power-law algorithm (Martin et al., 1987; Suess, 1980) for the attenuation of POC downward fluxes with depth (Francois et al., 2002; Lutz et al., 2002), along with the need to link changes in downward particulate export to the first-order kinetics commonly encountered in particle degradation and microbial respiration experiments (Berelson, 2001a; Lutz et al., 2002). These refinements have resulted in part from the larger, though still very sparse, geographical coverage of observations achieved in the past decade, and also from the application of new or improved techniques. The application of 234Th to estimate export from the surface ocean has provided both new perspectives and a means to evaluate export based on sinking particle traps (Buesseler, 1991, 1998). Similarly 230Th and 231Pa have allowed deep-sea particle fluxes to be ‘‘calibrated’’ (Bacon, 1996; Scholten et al., 2001; Yu et al., 2001). Modelling of dissolved nutrient distributions leading to estimates of particle fluxes and their transformations has also improved (Anderson and Sarmiento, 1994; Brea et al., 2004; Li and Peng, 2002; Matear and Holloway, 1995). All these advances have led to increasing acknowledgement that assessment of ‘‘export’’, the ‘‘biological pump’’, and ‘‘carbon sequestration’’ requires very careful definition of the process under investigation. The transport and transformation of carbon and other nutrients occurs via a wide range of mechanisms, and their influence depends very much on the depth and timescales at which these processes are considered (Fig. 1c, and for additional discussion see Ducklow et al., 2001; Oschlies and Kahler, 2004). 3. Question 2. What is the relative importance of the surface and subsurface ocean in the control of the biological pump? The fraction of NPP exported from the surface ocean is generally in the range of 2–20%, with values reaching 50% in some regions (as estimated from 234Th-based measures of export (Buesseler, 1998) and food-web studies (Laws et al., 2000)). The fraction of this export that sinks beyond 1000 m ranges from 6% to 25%, as estimated from deep-moored trap fluxes (Berelson, 2001a; Francois et al., 2002; Martin et al., 1987). Thus, based on changes in the bulk fluxes (i.e., NPP, surface and deep export fluxes) the surface and subsurface ocean layers are similar in their attenuation of the NPP signal. In contrast, we understand much less, at the process level, about the controls on the attenuation of this signal in either zone, and this lack of understanding limits the usefulness of our quantitative geochemical observations. Nonetheless, there has been progress. We briefly summarize the main processes important to surface and/or subsurface export (listed in Table 1), and examine the state of understanding of the attenuation of sinking particle fluxes with depth, including changes in elemental stoichiometries. 3.1. Oceanic processes affecting export Prior to JGOFS, NPP was thought to be the main process controlling export fluxes from the surface ocean (e.g., Pace et al., 1987; Suess, 1980). However, there was initial evidence that processes such as heterotrophic bacterial activity (Cho and Azam, 1988) or direct export of phytodetritus, associated with diatom blooms, to >3 km depth (Lampitt, 1985) would also impact surface export flux. Observations during the JGOFS era clearly demonstrate that the surface ocean is the site of both particle production and destruction (Table 1). The SIGMA lab mesocosm study (Alldredge and Jackson, 1995; Logan et al., 1995) confirmed the critical role that aggregation plays in particle production, resulting in the rapid transfer of algal carbon to depth. This downward transfer of carbon was driven by diatoms, and regional studies provided evidence (Boyd and Newton, 1995, 1999; Buesseler, 1998) for the important relationship between pelagic food-web structure (in particular the proportion of diatoms among primary producers) and the export flux from the surface ocean. Alldredge et al. (1993) demonstrated that a range of particles derives from transparent exopolymers (TEP) released from algal cells that act as biological glues. These glues are thought to be essential for initiating

282

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

the aggregation of particles at the low particle concentrations observed over much of the annual cycle (Engel, 2004). During JGOFS, detailed regional studies provided evidence for the role of heterotrophic bacteria in determining the magnitude of export from the surface ocean. Hoppe et al. (1993) demonstrated that bacterial solubilisation rates in the NE Atlantic are significant and thus lead to particle destruction. Lab mesocosm (Smith et al., 1995) and field studies off California (Smith et al., 1992) used biochemical assays to demonstrate the key role of bacteria in remineralizing particles. These experiments on particle degradation revealed how certain elements are preferentially solubilised and provided the first estimates of microbial solubilisation rates for C and N (Smith et al., 1995), and Si and C (Bidle and Azam, 1999; Bidle et al., 2002). Bacterial solubilisation of particles, although highest in the surface ocean and decreasing with depth, was also an important factor controlling export in the subsurface ocean (Table 1). Within this stratum, processes contributing to particle destruction – (rather than production) were dominant and driven by both bacteria and zooplankton (Table 1). Over the last decade, most progress has been made in elucidating the role of heterotrophic bacteria in modifying the downward export signal. In part, this is due to the novelty of research into both bacteria and particle transformations, but it also reflects the difficulties of conducting meaningful shipboard dietary experiments on animals such as vertically migrating zooplankton. Major advances in bacterial-particle studies include regional characterization of bacterial rate processes (e.g., ecto-enzyme activity) to depths of 1 km such as in the NE Atlantic (Turley and Mackie, 1995), NE Pacific (Boyd et al., 1999) and tropical and polar waters (Christian and Karl, 1995). Secondly, there have been concerted efforts to compile global syntheses of, and subsequently report key trends within, the vertical distributions of bacterial stocks, production and respiration (Duarte and Agustı´, 1998; Rivkin and Legendre, 2001). Rivkin and Legendre (2001) report that bacterial growth efficiency is greater at high latitudes and thus that ‘‘a greater proportion of production can be exported in polar than in tropical regions’’. Thirdly, there have been advances in describing bacterial community structure (Karl, 2003) with evidence of functionally specialised microbial communities in discrete depth strata over the upper 300 m of the water column (Wright et al., 1997). Advances in mesozooplankton research include the elucidation of a further pathway, the active transport of DOC (Steinberg et al., 2000), in addition to respiration of carbon and nutrients at depth (Dam et al., 1995). It is now established that many particle transformations in both the surface and subsurface ocean set the magnitude of POC export (Azam, 1998; Boyd and Stevens, 2002), and recent models incorporate food-web structure (e.g., Laws et al., 2000) and particle transformations (e.g., Moore et al., 2002b). However, few studies, either field or modelling, have evaluated the relative importance of particle transformations (Table 1) to export flux. Boyd et al. (1999) attempted to quantify the relative contributions of mesozooplankton and bacteria to export flux over the upper kilometer of NE Pacific waters. Their analysis suggested that both groups played an important role, but emphasized that major uncertainties, such as determining bacterial respiration and growth efficiency at depth, were evident. Difficulties in ranking the many particle transformation processes may stem from some processes being specific to particular water masses and/or exhibiting seasonality, whereas others are ubiquitous (Table 1). Thus, the relative importance of each process will vary with water mass and over the annual cycle. Some modelling studies have provided insights into specific aspects of these processes such as zooplankton feeding modes in relation to particle repackaging (Jackson and Burd, 2001), or the influence of bacterial solubilisation rates on particle residence times in the mixed layer (Boyd and Stevens, 2002). In summary, further modelling studies are now required that incorporate sensitivity analysis, and that consider each of the processes listed in Table 1. 3.2. Vertical attenuation of sinking particle fluxes in the subsurface ocean Quantitative expressions for the attenuation of POC fluxes with depth were already in common use at the time of the Dahlem Conference (Berger et al., 1989). Bishop (1989) compiled seven power-law expressions for flux at depth as a function of NPP, and demonstrated that none of these predicted deep POC fluxes as well as another power-law expression from Martin et al. (1987) that scaled deep fluxes to particulate fluxes at 100 m depth:

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

283

F ðzÞ ¼ F ðz0 Þ  ðz=z0 Þb in which z0 defines the reference depth of the surface layer export flux F(z0), and b characterizes the flux attenuation with depth. For b equal to 1, the flux decreases with the reciprocal of depth. Martin et al. (1987) suggested b  0.858 ± 0.1 for the open ocean based on a composite of limited observations from the Pacific Ocean, and this formulation has since been widely adopted, especially in global biogeochemical models (e.g., in all the Ocean Carbon Model Intercomparison Project (OCMIP) II runs; Doney et al., 2004). Regional variations in ‘‘b’’ in the range of 0.5–2 have been reported (e.g., Berelson, 2001a; Francois et al., 2002). However, most of these ‘‘b’’ evaluations have been based on decreases in fluxes to moored sediment traps at depths of 1000–4000 m, where decreases in POC fluxes with depth are relatively small. ‘‘b’’ values have also been inferred from comparison of deep fluxes with the outputs from surface ocean production or export models (Berelson, 2001a; Francois et al., 2002; Lampitt and Antia, 1997; Lutz et al., 2002). There have been very few direct determinations of flux variations in the subsurface ocean where flux attenuation is strongest, because determining subsurface (and surface) fluxes is fraught with difficulties. Under- and over-trapping can result from hydrodynamic effects as has been revealed by laboratory and field studies of trap behaviour (Gardner, 2000) and by radionuclide calibrations (Buesseler, 1991; Scholten et al., 2001; Yu et al., 2001). Furthermore, distinguishing flux contributions by sinking particles from those of zooplankton entering traps (i.e., ‘swimmers’) can be very difficult (Karl and Knauer, 1989; Michaels et al., 1990; Silver and Gowing, 1991). There has been progress on these issues, including the development of neutrally buoyant sediment traps (Buesseler et al., 2000; Stanley et al., 2004; Valdes and Price, 2000) to minimize hydrodynamic problems. Sediment traps have also been designed to exclude zooplankton ‘swimmers’ (Coale, 1990; Peterson et al., 1993, 2005), and other studies have focused on the contributions by ‘‘swimmers’’ to traditional sediment trap collections (e.g., Steinberg et al., 1998). These are promising approaches, but there are not yet sufficient data to assess whether present estimates of downward export with depth within the subsurface ocean will need to be significantly revised. One of the most striking aspects of the good fit of the power-law to many flux observations is the issues it raises with respect to developing a biogeochemical understanding of flux attenuation. Many physical, chemical, and biological processes follow ‘‘first-order’’ kinetics, i.e., their rates are proportional to the amount of material present. Thus, we might expect that particle transformations such as disaggregation, solubilisation, and remineralization would cause the flux to decrease exponentially with time (e.g., Armstrong et al., 2002; Lutz et al., 2002; Suess, 1980): F ðtÞ ¼ F ðt0 Þeðtt0 Þ=t



in which t* is the characteristic time for particle loss, and t0 is the time the particle passes the reference horizon z0. If the particles are sinking at a constant rate, s, this can be reformulated in terms of depth as (e.g., Lutz et al., 2002): 

F ðzÞ ¼ F ðz0 Þeðzz0 Þ=st ¼ F ðz0 Þeðzz0 Þ=z



in which z* is the characteristic length scale for the flux decrease with depth, or the ‘‘decay length’’. As shown in Fig. 2, an expression of this form differs considerably from the power-law formulation. In particular, firstorder decay sufficient to match the power-law decay at subsurface depths results in too low particle fluxes in the deep ocean. Comparison of the power-law and exponential-law expressions makes it clear that either particle degradation rates must decrease with depth or sinking rates increase with depth. There are many possible reasons for a decrease in particle decay rates with depth in the subsurface ocean. There are fewer bacterial cells at depth (Boyd et al., 1999). Colder temperatures at depth slow biological and chemical kinetics (Hoppe et al., 1993), as does pressure (e.g., Turley, 1993). Differential solubilisation of particles (e.g., Smith et al., 1992) may lead to slower remineralization rates, if biological activity preferentially removes biologically available compounds and leaves recalcitrant materials behind. This preferential remineralization may also increase particle sinking rates if it removes relatively buoyant organic material (such as highly hydrated mucus or TEP) relative to denser organic material or mineral phases (Boyd and Stevens, 2002). There is some limited evidence for an increase in sinking rates with depth (Berelson, 2001b).

284

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Fig. 2. Comparison of algorithms for flux variations with depth. A power-law produces a sharp decrease in flux in subsurface waters but retains some flux at depth (solid line for power-law exponent suggested for the open ocean by Martin et al., 1987). A first order exponential with decay length of 150 m captures the rapid decrease in the subsurface, but negligible flux arrives at depth (dashed line). The sum of two particle types with different rates of decay (or sinking) can parameterize rapid flux attenuation in subsurface waters and slower flux attenuation in deep waters (dot-dash and dotted lines, both for a 5% initial contribution of the slowly decaying component, e.g., Lutz et al., 2002).

Alternatively, the power-law attenuation of flux with depth could have its origin in the surface ocean, if materials with a wide range of sinking rates or ease of degradation are exported. For example, representing the vertical flux variations as the result of a sum of slow (z1 ) and fast (z2 ) degrading particles: z=z

F ðzÞ ¼ F 1 ðz0 Þ  e1

z=z

þ F 2 ðz0 Þ  e2

can produce a vertical flux profile that is similar to the power-law, provided the decay timescales (or sinking rates) differ by at least an order of magnitude (Fig. 2). Sinking rates have been observed to vary by several orders of magnitude (Diercks and Asper, 1997). The presence of a slowly degrading component of POC has been inferred from sediment studies (Keil et al., 1994), and from the observation that POC fluxes in the deep sea are in roughly constant proportion to biogenic mineral fluxes at many sites (Armstrong et al., 2002; Klaas and Archer, 2002). Thus it appears that several different, and possibly co-occurring, explanations for the power-law form of the decrease in flux with depth are possible. The basic power-law decrease in flux with depth has been abundantly demonstrated and is widely accepted. However the controls on its regional variability, and whether these controls derive primarily from particle properties that are imprinted in the surface ocean or particle transformations that occur in the subsurface ocean, remain unclear. More studies of particle fluxes in the subsurface where the flux changes most rapidly with depth are needed to resolve this issue, especially studies that separate the flux into components of differing reactivity and sinking rate. To date, studies of this type have primarily focused on changes in the bulk composition with depth, especially nutrient element stoichiometry and the proportions of organic and mineral matter. We discuss progress in these two areas in the next two sections. 3.3. Sinking particle degradation and elemental stoichiometry Construction of ocean carbon budgets, food-webs, and global biogeochemical cycles has long been guided by the simplification introduced by Redfield et al. (1963) that important nutrient elements (C, N, P, O, etc.) are utilized by phytoplankton in fixed ratios. Field programs during the JGOFS era revealed a multitude of important departures of the order of 20–50% from this stoichiometric simplification, especially for community nutrient consumption ratios in the euphotic zone over periods of days to months (e.g., Bates et al., 1996; Copin-Montegut, 2000; Kahler and Koeve, 2001; Karl et al., 2001; Kortzinger et al., 2001; Lourey and Trull, 2001; Rubin, 2003; Sambrotto et al., 1993). This ecological and biological plasticity is now reasonably well accepted (see Falkowski and Davis, 2004; Geider and La Roche, 2002; Ho et al., 2003), and its importance to global biogeochemical cycles, for example in terms of increased algal carbon fixation per unit N or P,

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

285

has been recognized (e.g., Arrigo et al., 1999; Denman and Pena, 2000; Shaffer, 1996; Thomas et al., 1999; Toggweiler, 1993). Clear consensus that the remineralization of sinking or suspended particles also has variable stoichiometry has been slower to develop. Studies of suspended particles in the subsurface ocean (see Table 2) revealed gradients in nutrient ratios indicative of more rapid remineralization of N and P than organic C in particles (e.g., Boyd et al., 1999; Christian et al., 1997). Similarly, examinations of particles and brine solutions from deep sediment traps (e.g., Honjo and Manganini, 1993; Martin et al., 1987; Newton et al., 1994; Schneider et al., 2003) also revealed gradients in nutrient ratios often of the order of 20–50% but ranging to a factor of 2 or more. But early efforts to assess particle remineralization ratios from dissolved nutrient distributions suggested that long-term average C/N/P ratios were close to that of the Redfield stoichiometry for phytoplankton (Anderson and Sarmiento, 1994). This led to the perspective that particle-based observations of non-Redfield stoichiometry were temporally or spatially biased, and that on average over decadal to centennial scales and at ocean basin spatial scales fixed Redfield stoichiometry held for particle export and remineralization. Subsequent examinations of dissolved nutrient fields have found that export and remineralization estimates are easily biased by small errors in circulation estimates (Matear and Holloway, 1995). Regression analyses suggest that large deviations (e.g., 20–50%) from Redfield stoichiometry do occur with depth in the ocean interior (Shaffer et al., 1999) and with circulation patterns at oceanic basin scales (Brea et al., 2004; Li et al., 2000; Li and Peng, 2002). These and other recent studies have generally found P and N remineralization to be more rapid than that of organic C (Table 2). The implications of element-specific remineralization length scales for global biogeochemical cycles are beginning to be explored in global models. These include simulations in which the biological pump strength is increased relative to estimates based on Redfield stoichiometry (e.g., Schneider et al., 2003; Shaffer, 1996). There has also been debate about the possibility of changes in remineralization ratios over decadal timescales from examinations of deep ocean dissolved nutrient concentrations. However, no clear conclusion has emerged because of issues of data quality and the difficulty of separating nutrient remineralization changes from circulation and nutrient source region changes (Pahlow and Riebesell, 2000; Zhang et al., 2000). There is agreement that changes in whole ocean nutrient inventories can only occur on the timescales of nutrient element residence times (which for N and P are on the order of 103–104 yr (Coale et al., 2001; Falkowski and Davis, 2004; Tyrrell, 1999)). Study of the processes that control the stoichiometry of remineralization of organic matter are still in their infancy. Organic geochemical studies are able to examine changes in proportions of major compound classes (proteins, lipids, polysaccharides, etc.) and also individual compounds. However, the majority of particulate organic matter remains difficult to characterize (Hedges et al., 2000) and simple, broadly applicable links to elemental stoichiometry have yet to emerge. For this reason, and because review of organic geochemical studies is beyond our scope, we simply note a few important recent results. Study of subsurface suspended particles in the Equatorial Pacific using molecular biological techniques suggests that contributions to the particle pool from zooplankton and heterotrophic bacteria increase with depth (while that from phytoplankton detritus decreases), but also suggest that these contributions vary geographically (Sheridan et al., 2002). Incubation experiments have found that the relative degradation rates of C and N rich compound classes (sugars and proteins, respectively) vary with water mass across fronts in the Southern Ocean (Panagiotopoulos et al., 2002). The role of dissolved oxygen levels in controlling degradation rates and degradation stoichiometry of sinking particles appears to be minor (Kristensen et al., 1995; Pantoja et al., 2004), but may influence the final extent of degradation (Van Mooy et al., 2002). 4. Question 3. What is the biogeochemical role of mineral particles? Minerals make up the majority of particulate matter fluxes in the deep ocean (e.g., Francois et al., 2002; Wefer, 1989). Three types dominate: lithogenic silica from aerosols and riverine inputs, biogenic silica (opal), and biogenic carbonates. These minerals play several roles, which must be considered together to determine their function in carbon biogeochemistry. They influence POC export fluxes in two distinct ways: (i) by adding ballast that increases the specific gravity of marine aggregates and thus sinking rates (Francois et al., 2002; Ittekkot et al., 1992; Klaas and Archer, 2002), and (ii) by protecting organic matter, through association with

286

Table 2 Estimates of preferential remineralization rates of exported organic matter Study

Region

Depth (m)

Remineralizationa order

Remineralization ratios (molar)

Method and commentsb

Redfield et al. (1963)

Surface waters

C=N=P

C:N, N:P 6.6, 16

Particles, phytoplankton, i.e., production ratios

Bisset et al. (1999)

BATS

Near surface 100–200

N>C

C:N exported

1-D seasonal model for particulate and dissolved organic matter, sinking flux attenuation

N>C

100 m 6.7 200 m 7.1 C:N exported

NP

70 m 8.7, 150 m 9.9 200 m 11.8 C:N, N:P exported

Anderson and Pondaven (2003)

HOT

70–400

150–500

P>C

Abell et al. (2000)

NE Pacific

0–300

CN

Martin et al. (1987)

NE Pacific

35–2000

N>C

Boyd et al. (1999)

OSP

100–1000

N>C

Wang et al. (2003)

APFZ

0–1000

P>N

Schneider et al. (2003)

Global

100–5000

N>C

Loh and Bauer (2000)

N. Pacific and Southern Ocean

25–4000

P>N > C

Li and Peng (2002)

Global

>500

P>C

Shaffer et al. (1999)

Atlantic

750–1500

NP P>C

Anderson and Sarmiento (1994) Brea et al. (2004)

Atlantic

>500

NP NPC

S. Atlantic

0–300

P>C

150 m 8 300 m 10.3 500 m 12.1 C:N of remineralization 30 C:N of remineralization 6.2* (z/ 100)Ù 0.13 C:N exported 100 m 7.0 200 m 8.1 Z P ¼ 180 m, Z N ¼ 180–250 m C:N exported 7.1 increasing 0.2/ 1000 m C:N, N:P standing stocks DOM: 9 to 30, 16 to 60 POM: 6 to 14, 21 to 75 C:P, N:P remin. ratios 73 to 124, 15

1-D seasonal model for DIC, nitrate, and C/N of shallow sediment traps

1-D model for particulate and dissolved nutrients, high initial surface N:P from nitrogen fixation

TOC, TON,TOP distributions, hypothesized highly labile TOC produced via nitrogen fixation Free-drifting traps best-fit power-law Sediment trap flux ratios

e-folding lengthscales Z* to fit seasonal nutrient depletion Compilation of sediment traps Dissolved organic matter and suspended particulate organic matter

Dissolved nutrient distributions, C:P range across ocean basins

C:P, N:P flux across depth range 130, 16

Dissolved nutrient distributions fluxes

Redfield within uncertainties

Dissolved nutrient distributions

C:P, N:P remin. ratios 80 to 100, 17 to 18

Dissolved nutrient distributions

N > P a

The order in which a given element is transformed e.g., N > C means N is removed from particles more rapidly than C, N  P means N and P are remineralized at similar rates. b Depending on the method used the relative rates may apply to loss from the sinking flux (i.e., via particle disaggregation or solubilisation), to actual remineralization, or to some combination of these processes. See text for discussion.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Christian et al. (1997)

BATS

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

287

mineral grains, from remineralization (Hedges et al., 2000; Keil et al., 1994; Mayer, 1999). The effectiveness of minerals in these roles depends in part on their forms and origins, which are directly biologically mediated for carbonates and opal (Schiebel, 2002; Verity and Smetacek, 1996; Fig. 3), and biologically influenced via aggregation for lithogenic minerals. For biogenic carbonates there is a third effect; their precipitation elevates the CO2 (aq) concentration by 0.6 mol/mol precipitated. These alkalinity changes associated with carbonate formation offset the positive influence of carbonate minerals on the transfer of carbon from the atmosphere into the ocean interior. We examine these three biogeochemical roles in turn. 4.1. Minerals and the enhancement of POC export The role of minerals in promoting POC export has been re-evaluated using the growing global set of deepmoored sediment trap observations (e.g., Armstrong et al., 2002; Francois et al., 2002; Klaas and Archer, 2002). These studies have underlined a strong correlation of POC flux with ‘‘ballast’’ mineral fluxes (Armstrong et al., 2002) and indicate that calcite is more efficient than other minerals in promoting POC export (Francois et al., 2002; Klaas and Archer, 2002). Here, we briefly review these studies and then test if the statistical relations obtained from globally-distributed observations also hold at locations where multi-year timeseries are available. Francois et al. (2002) defined the ‘‘efficiency’’ of POC export to depth (2000 m) as a fraction of surface export flux (FluxCorg/EP in their terminology) and noted that this efficiency increased with the mineral flux (Fig. 4a). Based on multiple linear regression of this efficiency versus the corresponding fluxes of mineral components for 64 trap studies, they concluded that carbonates play a much larger role in enhancing POC export flux than either lithogenic or biogenic silicates (which were found to have negligible influences). This result is illustrated in Fig. 4a which shows a strong separation between fluxes dominated by carbonates, which exhibit high POC ‘‘efficiency’’ in comparison to fluxes dominated by silica. This conclusion depends in part on their estimates of surface export flux – which were based on SeaWiFS ocean colour measurements combined with NPP and export production (f-ratio) algorithms. The separation of fluxes based on mineral type is not nearly as strong in terms of the direct correlation of POC and mineral fluxes (Fig. 4b).

Fig. 3. Schematic overlaying ballast producers onto a simplified open-ocean pelagic food-web (from Michaels and Silver, 1988). The grey arrows represent grazer-mediated pathways of energy transfer between food-web components, and the black arrows denote the main pathways through which carbon is exported to depth. a denotes the contribution of calcite liths to biogenic particles (Passow, 2004), b is the direct export of opal via algal aggregation and c represents the sinking flux associated with grazers (i.e., faecal pellets, exuvia and passive flux of organisms). The black-framed boxes denote sources of opal for particle ballasting, while the hatched framed boxes represent sources of calcite ballast for sinking particles. Note that the production of mineral ballast occurs at multiple trophic levels – from small phytoplankton such as coccolithophores to omnivorous grazers such as foraminifera (Anderson, 1993; Caron and Swanberg, 1990). Opal production – from diatoms to radiolarians also spans several trophic levels (Scharek et al., 1999). Another example of the implications of food-web complexity for biogeochemical transformations is that the dissolution of opal has been linked to the destruction of organic (carbon containing) surface membranes by bacteria (Bidle et al., 2002).

288

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312 0.20

a

0.18

FCorg / EP

0.16 0.14 0.12 0.10 0.08 0.06 0.04 0.02 0.00 6

b

5

FCorg

4 3 2 1 0 0

20

40

60

80

100

120

Fmin >30% C (Nordics) > 25% lithics > 50% opal, < 30% calcite > 30% calcite

Fig. 4. Global correlations of POC and mineral fluxes. (a) A plot of the annual POC flux (FCorg) normalized to annual export production (EP) versus the corresponding annual mineral flux (Fmin) in deep traps (>2000 m depth) at a range of sites as reproduced from Fig. 2 in Francois et al. (2002). The units for FCorg, EP, and Fmin are g m2 yr1, and the EP data were derived from the modelling study of Laws et al. (2000). (b) A plot of the annual POC flux (as presented in (a) but with the EP normalization removed) versus the corresponding annual mineral flux (all data courtesy of R. Francois). The range of symbols reflects the classification of trap sites employed by Francois et al. (2002). The term Nordics represents data from sites north of 65°N, and between 1 and 7°W.

To quantify the role of the different ballasts directly, we repeated the multiple linear regression analysis of Francois et al. (2002) using only the POC and mineral flux data (rather than the efficiency estimated from the FluxCorg/EP ratio). That is, we evaluated the ‘‘carrying coefficients’’ of different minerals that best express the POC flux as the sum of contributions from carbonate, opal, and lithogenic particles: F POC ¼ aF opal þ bF CaCO3 þ cF litho This re-analysis confirms an important ‘‘POC carrying coefficient’’ for carbonates (i.e., the least-squares maximum likelihood estimate of the weight percentage of POC associated with PIC flux as determined from their flux correlations), but also shows a similar capacity for lithogenic material, and a lower capacity for biogenic silicates (Table 3). The calculated carrying capacities are very similar to those estimated for an overlapping set of deep trap results from 78 locations (Klaas and Archer, 2002). Thus, there appears to be consensus about the globally-averaged relationships among deep POC and deep mineral fluxes. All minerals correlate strongly with POC flux, with biogenic silicate perhaps less important than carbonates or lithogenic silicates. We also applied this multiple linear regression approach to time-series results from deep traps (on a cup-bycup basis representing weekly to monthly changes over several years) at Ocean Station Papa (OSP) in the subarctic NE Pacific (Wong et al., 1999) and the Bermuda SCIFF site in the subtropical N. Atlantic (Conte and Ralph, 1999). This analysis also suggests that all minerals are important for ballasting, but without a clear consensus for a greater role for carbonates than for opal or lithogenic material. For example, carbonate had a lower carrying coefficient than opal or lithogenic material at Bermuda (Table 3). The estimated relative importance of the different minerals is sensitive to small changes in the methodology. Because many sediment trap programs do not measure lithogenic fluxes, these are often estimated from total

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

289

Table 3a Ballast mineral ‘‘POC carrying-coefficients’’ (weight % POC in dry mineral) Data source Klaas and Archer (2002)e 78 traps deeper than 2000 m annual means

Best estimatef: 95% rangeg:

Biogenica,b CaCO3

Biogenica,c SiO2

Lithogenica,d material

Correlation coefficient

7.5 6.4–8.6

2.9 2.0–3.7

5.2 3.4–7.0

–

7.4 5.1–9.7

0.93

POM massd = 2.2 * POC mass Francois et al. (2002)h 62 traps deeper than 2000 m annual means

Best estimate 95% range

7.4 6.4–8.4

1.5 0.8–2.2

POM mass = 1.87 * POC mass Wong et al. (1999) Ocean Station P 3800 m depth 200  bi-weekly cups for 1982–1993

Best estimate 95% range

2.1 1.3 0.2–3.9 0.8–3.4 POM mass = 1.87 * POC mass

23.3 17.0–29.7

0.69

Wong et al. (1999) Ocean Station P 3800 m depth 200  bi-weekly ups for 1982–1993

Best estimate 95% range

2.5 3.3 0.5–4.5 1.1–5.5 POM mass = 2.3 * POC mass

13.8 6.7–20.9

0.56

Conte et al. (2001) Bermuda SCIFF site 3200 m depth 31  bi-monthly cups for 1978–1984

Best estimate 95% range

4.5 6.3 3.8–5.3 2.4–10.2 POM mass = 1.87 * POC mass

6.5 3.4–9.6

0.98

a

Most results for <1 mm fraction, except <0.125 mm for Bermuda. Measured for most studies, expressed as mass of CaCO3. c Measured for most studies, expressed as mass of SiO2. d Measured at some sites, but mainly estimated from total mass by difference, using: lithogenic mass = total mass  CaCO3– SiO2  A * POC. A is given for each set of results – see text. e Klaas and Archer (2002) results taken from their Table 2, all others calculated here from data in references. Sites were globally distributed. f The best estimate is the maximum likelihood value from the multiple linear regression. g Bounds for 95% of the distribution of carrying coefficients. h Sites were globally distributed (see Fig. 6b). b

Table 3b Observed proportions of each component in the total flux (weight % in drymass) Data source

Biogenic CaCO3

Biogenic SiO2

Lithogenic material

Biogenic POC

Francois et al. (2002)a 62 annual means

Maximum Minimum Median

78.6 11.3 53.6

82.4 5.2 22.8

60.7 3.7 10.8

7.4 1.6 4.8

Ocean Station P 3800 m 200 cups

Maximum Minimum Median

72.6 11.0 42.6

71.1 11.4 39.8

55.9 5.0 9.4

14.1 1.3 2.8

Bermuda SCIFF site 3200 m 31 cups

Maximum Minimum Median

65.5 7.3 18.3

17.0 1.8 4.4

13.3 2.4 5.8

4.5 0.6 1.5

a

Sites were globally distributed (see Fig. 6b).

mass by subtracting opal, carbonate, and organic matter masses, which first requires an estimate of organic matter mass from POC measurements. The choice of the factor for converting POC to organic matter can affect the subsequent correlations. For example, for the OSP record as shown in Table 3 use of relatively low mass for organic matter (1.87 * POC; Anderson, 1995) to estimate lithogenic fluxes suggests carbonates carry more POC than biogenic silicates, but the opposite conclusion is reached if a higher organic mass is used (2.3 * POC; Conte et al., 2001).

290

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Examination of time-series observations is also useful for confronting mechanistic models proposed to explain the observed correlations of POC and mineral fluxes. For example, it is clear that the ‘‘POC quantitatively associated with minerals’’ model of Armstrong et al. (2002) is over-simplified. The observational range of POC contents on a cup-to-cup basis (e.g., 1.3–14.1% of dry mass at OSP, Table 3) far exceeds the small variation among annual mean values from regional data sets that they used to justify a constant, mineral-associated percentage for deep POC export. It is also worth emphasizing that the mean correlations of deep POC export to mineral fluxes may not be a good guide to potential climate-change mediated changes in POC export. These relationships do not appear to hold in years of anomalously high POC fluxes (e.g., results presented from ENSO years in Conte et al., 2001, Wong et al., 1999). There are not many multi-year records of export, but those that do exist call into question the representativeness of short-term records, as well as the reliability of any models of export that do not explicitly address seasonal variability. For example, the longest existing record (at OSP in the Northeast Pacific) exhibits variability of annual mean fluxes of at least a factor of 3 (Fig. 5a; Wong and Matear, 1999), and the next-longest record (at the SCIFF and OFP sites off Bermuda) exhibits large interannual variability in seasonality (Fig. 5b; Conte et al., 2001). Overall, the geochemical statistical relationships summarized in Table 3 and in previous work (Armstrong et al., 2002; Francois et al., 2002; Klaas and Archer, 2002) provide an overview of global flux variations and suggest a correlation with mineral fluxes. However, they do little to identify the mechanisms responsible for controlling POC flux, much less its probable response to climate change. This is in part because the geochemical measurements provide overly simplified proxies for biogeochemical processes (Verity and Smetacek,

a

2.5

-2

-1

POC flux (mmol m d )

3.0

2.0 1.5 1.0 0.5 0.0 1982

1984

1986

1988

1990

1992

1994

Year

MASS flux (mg m-2 d-1 )

100

b

80

60

40

20

0 1990

1992

1994

1996

1998

Year

Fig. 5. Interannual variability of the annual magnitude and seasonal cycles of deep export fluxes (a) annual mass fluxes at 3800 m depth at Ocean Station Papa (50N 145W) varied 3-fold over a 20 year record as shown by the white circles for each year (Wong and Matear, 1999). The highest flux year (1983) had unusually high POC and opal contents, in comparison to other years in which POC correlated well with PIC; (b) the seasonal structure of mass fluxes at 3200 m depth off Bermuda varied considerably over a 10 year record, in particular in terms of the presence and magnitude of a spring maximum, yet annual mean fluxes with constant within 15% (Conte et al., 2001).

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

291

1996). For example, lithogenic materials have a wide range of provenance and transport mechanisms (Jickells and Spokes, 2001), while biogenic carbonates and silicates have important sources in both the autotrophic and heterotrophic communities (Fig. 3). Observations on carbonate and opal (Deuser et al., 1995; Schiebel, 2002) indicate that they derive in similar proportions from phytoplankton in the surface ocean (primarily coccolithophores and diatoms), and zooplankton in both surface and subsurface realms (pteropods, foraminifera, and radiolarians). This range of provenances preclude a simple separation of the relative roles of food-web structure and ballast minerals in driving particle export (Fig. 3). 4.2. Life cycles of calcifiers and silicifiers and their influence on POC export Contributors to the oceanic carbonate budget include coccolithophores, pteropods, dinoflagellates with calcareous tests (e.g., Thoracosphaera) and foraminifera (US-JGOFS, 2001). The relative contributions to carbonate production were recently summarized by Schiebel (2002), who reports that 32–80% of the deep ocean carbonate budget is due to planktonic foraminifera, with 12% from coccolithophores, 10% from pteropods and 3.5% from dinophytes. However, due to wide regional and temporal variations in the contribution of these calcareous components, and the perennial issue of under-sampling the ocean, uncertainties remain as to their relative contributions to the global calcite budget (Schiebel, 2002). In contrast, no partitioning of the global opal budget (Tre´guer et al., 1995) into the main groups of silicifiers has so far been attempted, and so regional studies (Takahashi et al., 1990, 1991) provide the best indicators of the main biogenic opal producers. At OSP, diatoms, silicoflagellates, and radiolarians were evident in trap samples at 3800 m (Takahashi, 1991). Diatoms dominated POC fluxes (based on a biovolume:carbon algorithm) both over the annual cycle, and the seasonal POC flux maxima. Radiolarians contributed around a third that of the diatom flux to the POC export, and silicoflagellates made an insignificant contribution (Takahashi, 1991). In the Tropical Atlantic, N. Central Pacific and Panama Basin, radiolarian fluxes contribute 40%, 51% and 27%, respectively to biogenic opal fluxes at around 4000 m depth, with their contribution increasing with depth relative to that of diatoms (Rau et al., 1991). Thus, biogenic minerals derived from both phytoplankton and zooplankton contribute to global carbonate and opal budgets (Fig. 3). Their different life cycles mean that they contribute to mineral budgets at different depth horizons (Schiebel, 2002) and exhibit different seasonal signatures (e.g., for coccolithophores see Brown and Yoder, 1994). Schiebel (2002) provides evidence of pronounced temporal and spatial variations in both the abundances and export of foraminifera. Examples include marked seasonality with peak abundances of live foraminifera in the upper 0–150 m in April/May at the BIOTRANS site (N. Atlantic), and a deeper range during this period for the empty tests (upper 500 m). At the WAST site (Arabian Sea), the 0–100 m horizon had the highest abundances of living foraminifera (>200 animals m3) but with intermediate numbers (30– 100 m3) down to 500 m (Schiebel, 2002). There was evidence at both sites of episodic (within one month) carbonate pulses to depth, pulses representing 31–77% of the annual carbonate export, and that would impact strongly the ballasting or ‘POC protection’ of sinking particles. Fewer data were available on the seasonality and depth distributions of the silicifiers. In tropical and subtropical waters, the majority of radiolarian populations are observed in the upper at 50–100 m, often with two discrete maxima (near surface and 50–100 m depth) (Abelmann and Gowing, 1997; Kling and Boltovskoy, 1995), whereas in the Southern Ocean, the highest abundances are mainly observed at 200–400 m depth (Abelmann and Gowing, 1997; Boltovskoy and Alder, 1992). These temporal and spatial signals are likely to have marked implications for the influence of biogenic minerals on POC export fluxes, as are the details of their specific forms. For example, the role of biogenic carbonates in ballasting POC for export in aggregates will be reduced in situations where peak seasonal production of carbonates does not coincide with peak POC production, and reduced also where carbonate production occurs at greater depths than POC production (e.g., in the Sub-Antarctic Zone south of Australia; Trull et al., 2001a). Ballasting efficiency will also be affected by the size and geometry of different calcifiers, which range from coccolithophores (5–30 lm) to foraminifera and pteropods (300–3000 lm). For example, pteropods have very fast sinking rates (Schiebel, 2002) and will probably exit the upper ocean with few interactions with other particles, but they may scavenge slower sinking/suspended particles in the subsurface via differential sinking (Kepkay, 1994). In contrast, coccolithophore carbonate may be mainly present as suspended mineral fragments whose

292

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

removal is controlled by the formation of organic matter dominated aggregates (Passow, 2004; Passow and De La Rocha, 2006). The detailed properties of biogenic minerals may also impact their ability to protect POC from remineralization. Organic geochemical studies of sediments (Keil et al., 1994; Mayer, 1994, 1999) suggest that incorporation of POC into the mineral matrix provides protection (pores of <10 nm probably exclude bacterial ectoenzymes). Thus, mineral surface area and the distribution of organic matter on this surface is an important determinant of POC protection (Mayer, 1999). It is difficult to obtain such mineral surface area estimates for oceanic particles, but the range of sizes (and hence surface area:volume ratios) observed for silicifiers and calcifiers suggest that large variations in mineral surface area are likely. Observations on biogenic mineral abundances and forms in the subsurface ocean are particularly sparse, but again it appears that the details of mineral form and provenance are important. Export flux of planktic foraminiferal tests decreases with depth due to processes over the upper kilometer, including microbially-mediated decomposition of cytoplasm and decreasing pH within the tests (Schiebel, 2002). For the silicifiers, the contribution of radiolarians to bulk opal fluxes tripled from 15% at 300 m depth to 45% at 4000 m in the Tropical Atlantic, N Central Pacific and Panama Basin. This trend is indicative of greater dissolution of components of the opal flux such as diatom spicules (Bidle et al., 2002) than of intact radiolarian skeletons (Rau et al., 1991). The relative importance of biogenic minerals may also change dramatically in the subsurface. Dissolution of biogenic opal commences as shallow as 100 m where significant dissolution of diatoms (i.e., up to 50%) has been recorded (Boyd et al., 2004; Tre´guer et al., 1995). In contrast, carbonates will not undergo significant dissolution rates above the lysocline depth (2000 m for most of the ocean), although recent studies suggest dissolution also occurs above this depth (Feely et al., 2004). Finally, we note that the role of calcifying plankton in biogeochemical cycles may already be changing in response to climate change (see Question 6). Lab experiments suggest that coccolithophore carbonate precipitation is reduced by elevated CO2 (Riebesell et al., 2000). Paleoceanographic observations find lower foraminiferal shell weights associated with altered N. Atlantic circulation (Barker et al., 2004), and there appears to be recent, temperature-mediated, geographical expansion of coccolithophore blooms into the Bering Sea (Napp and Hunt, 2001). 4.3. Overall influence of carbonate export on atmosphere–ocean CO2 partitioning Carbonate minerals increase the downward export of POC, but this positive influence on the transfer of carbon from the atmosphere into the ocean interior is offset by the decreased alkalinity and increased pCO2 that accompanies carbonate formation. For each mole of carbonate precipitated, molecular CO2 increases by 0.6 mol (ranging from 0.5 to 0.8 mol in surface waters with larger increases occurring in colder waters; Frankignoulle et al., 1994; Koeve, 2002). Thus, for carbonate precipitation to assist in the sequestration of CO2, each mole of carbonate exported must increase the transfer of organic carbon by 0.6 mol, i.e., a POC/PIC export ratio of 0.6 or higher must be achieved. The carrying coefficients for POC by PIC found from deep trap correlations are less than 10% by weight (Klaas and Archer, 2002; Table 3), which is equivalent to less than 0.01 on a molar basis and thus insufficient to offset the effect of decreased alkalinity. Thus, even if we assume that no POC would have been exported without the concurrent export of PIC, the dominant effect of PIC formation is to reduce rather than accelerate CO2 uptake from the atmosphere, at least on the century to millenial timescales that apply to fluxes to the deep sea. On much shorter timescales (months), primary production by carbonate-documented communities has been observed to reduce surface water pCO2 (Buitenhuis et al., 2001). This reflects the high POC/PIC ratio of these communities of 4 or more (Balch et al., 1996; Koeve, 2002). However, there is no clear indication that PIC production promotes POC export at shallow depths beyond that which would occur without the PIC production. Indeed the POC/PIC export ratio (or ‘‘rain ratio’’) estimated from seasonal alkalinity and nitrate depletion tends to decrease rather than increase with the extent of PIC export from surface waters (Koeve, 2002). In summary, despite the strong correlation of deep carbonate export with POC export observed at global scales (Francois et al., 2002; Klaas and Archer, 2002), the dominant effect of carbonate formation is to weaken the biological pump. This provides a positive feedback to increasing atmospheric CO2 levels (Ridgewell et al., in press), as has been previously emphasized from a conceptual standpoint by Frankignoulle et al.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

293

(1994) and quantified from examination of deep-moored sediment traps for the Sub-Antarctic Zone (Trull et al., 2001a) and the Atlantic Ocean (Koeve, 2002). 5. Question 4. What insights have we gained from regional comparisons? In addition to the examination of food-web structure and mineral flux controls on export, other researchers have asked whether major ocean basins or oceanographically defined regions contribute disparately to global export. Compilations of trap data have found more than 10-fold variations in annual average POC export to the deep sea (Antia et al., 2001; Berelson, 2001b; Fischer et al., 2000; Francois et al., 2002; Lampitt and Antia, 1997; Lutz et al., 2002) but a clear consensus on regional differences and their origins remains elusive. Because the global coverage remains sparse, we focus our review on two regions which have received considerable attention in recent years – the Southern Ocean and the NE Atlantic. We then return to the issue of the reliable estimation of global export patterns in light of the regional results. 5.1. Regional case-studies – the Southern Ocean The open Southern Ocean has deep surface mixed layers, only moderate seasonal warming or stratification, and large areas that are characterized by the High-Nitrate-Low-Chlorophyll (HNLC) condition (e.g., Tre´guer and Jacques, 1992; Trull et al., 2001b). Interest in export in this region is high, because of its possible role in moderating atmospheric CO2 (Sigman and Boyle, 2000), but for logistical reasons observations remain limited. There is strong consensus for relatively high POC surface export flux (as a fraction of NPP) in the Southern Ocean in comparison to lower latitudes. Based on 234Th deficits, several studies (summarized in Buesseler, 1998, Buesseler et al., 2001) suggest that surface export flux is on the order of 30–50% of NPP (based on shipboard 14C incubations). This range is corroborated by other approaches, including 15N-uptake experiments (e.g., Sambrotto and Mace, 2000; Savoye et al., 2004), mixed layer nitrate depletion (Nelson et al., 2002), and modelling of seasonal nutrient cycles (Pondaven et al., 2000). This relatively high surface export, as a fraction of NPP, results in ‘‘annual POC export that is similar to that in more productive systems at lower latitudes, in spite of the lower primary productivity’’ (Nelson et al., 2002). Views on the magnitude of deep POC export fluxes in the Southern Ocean, and their relation to NPP and surface export, have been more disparate and are evolving. Lampitt and Antia (1997) examined export fluxes from three trap sites in the S. Atlantic and suggested that POC export to 2000 m depth was ‘‘very low’’ as a fraction of NPP as derived from satellite (CZCS) ocean colour (Longhurst, 1998). This result depended in part on low POC fluxes in these traps, and more recent studies (in the Indian and Pacific sectors of the Southern Ocean) have found higher POC fluxes. Trull et al. (2001a) compiled results from 18 traps at sites in open-ocean ice-free waters and suggested that POC export fluxes (>1000 m) were indistinguishable from the global median of 1.0 g C m2 yr1 at 2000 m estimated by Lampitt and Antia (1997). Based on a similar compilation, Nelson et al. (2002) suggested that deep fluxes were close to twice the global average of 0.96 g C m2 yr1 at 2000 m depth reported by Honjo et al. (2000). Fischer et al. (2000) examined six sites in the S. Atlantic and reported that POC export to 2000 m depth was ‘‘low to moderate’’ as a fraction of NPP (as estimated from CZCS Ocean Colour by Antoine et al., 1996). The global study of Francois et al. (2002) focused on depths >2000 m and included only one Southern Ocean sediment trap datum. However, based on that and sediment pore water properties at four additional sites along the S.W. Pacific AESOPS transect (Sayles et al., 2001), they tentatively suggested ‘‘low’’ deep export efficiency (expressed as a fraction of their surface export flux based on NPP and f-ratio estimates from 15N uptake experiments). In contrast, Lutz et al. (2002) included 10 sites, but focused mainly on 5 from the S. Atlantic because the AESOPS sites did not yield data from multiple trap depths. They concluded that deep export fluxes in the S. Ocean were ‘‘above average’’, in comparison to both NPP (based on CZCS Ocean Colour estimates of Arrigo et al. (1998)) and surface export fluxes (from NPP and f-ratio estimates from 15N-uptake experiments of Sambrotto and Mace, 2000). These varying estimates of deep POC export, as a fraction of surface NPP or surface POC export estimates, have led to suggestions that the POC remineralization length scale may differ in the Southern Ocean from those in other regions. Berelson (2001a) combined SW Pacific AESOPS results from 4 (>2000 m) sediment

294

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

traps with sediment pore water results and NPP datasets. He suggested that Southern Ocean surface POC export is rapidly attenuated with depth in comparison to flux attenuation in the Arabian Sea and Equatorial Pacific (based on JGOFS data). This view was also expressed by Francois et al. (2002) based on a different selection of AESOPS data. In contrast, Nelson et al. (2002) compared the AESOPS SW Pacific new production estimates with trap export fluxes at around 1000 m depth, and concluded that ‘‘remineralization efficiency is low’’. As noted above, Lutz et al. (2002) also inferred low remineralization efficiencies from their examination of traps moored at different depths (>2000 m) in the S. Atlantic. Some evidence for relatively rapid and shallow remineralization of POC has been provided from studies of subsurface particulate barium distributions in polar water south of Australia (Cardinal et al., 2004); and from subsurface suspended particle distributions south of New Zealand (Lam and Bishop, in press). In summary, over time and with the inclusion of additional observations it appears that the assessment of Southern Ocean POC export to the deep ocean has changed from ‘‘below average’’ to ‘‘near or above average’’ in comparison to other regions of the global ocean. This conclusion must be tempered by recognition of several aspects of these analyses: (i) The assessment of export efficiency to the deep ocean is highly dependent on estimates of NPP or surface export flux to which deep export fluxes are compared. For example the ‘‘above average’’ export efficiency from (Lutz et al., 2002) is based on application of an f-ratio of 0.15 to estimate surface export flux, whereas a higher estimate, e.g., 0.4 as applied by Nelson et al. (2002), would reduce this to an ‘‘average’’ efficiency. Similarly, if satellite estimates of NPP are too low in the Southern Ocean (as suggested by Schlitzer, 2002), then reports of the ‘‘above average’’ export efficiency as a fraction of NPP could be an overestimate. (ii) Trap fluxes are subject to biases, including variable trapping efficiency, and modification by captured zooplankton (e.g., Gardner, 2000; Usbeck et al., 2003). There is so far no clear way to correct for these biases. Efforts to calibrate collection efficiency using radionuclides suggest that traps at <1000 m depth may be particularly prone to bias. However, greater than 2-fold variations also occur in deeper traps, and there is not yet a defined path for the correction of carbon and other component fluxes based on the radionuclide results (e.g., Scholten et al., 2001; Yu et al., 2001). Notably, the only study to date to compare radionuclide calibrated with un-calibrated trap fluxes in the Southern Ocean found that such calibration did not alter their conclusions regarding regional variations in export efficiency (Lutz et al., 2002). (iii) The Southern Ocean exhibits regional variations in deep POC export among different circumpolar zones and also among different sectors. There is also substantial temporal variation at individual locations where multi-year records have been obtained (e.g., Honjo et al., 2000). These variations contribute to (and in part their detection is confounded by) the seeming discrepancies among the global and regional reviews summarized above. Indeed, Lampitt and Antia (1997) suggested that high seasonal and interannual variability may be the most important characteristic of polar environments, rather than any particular bias toward high or low export. There have not yet been enough multi-year experiments to objectively assess this view. 5.2. Regional case-studies – the Northeastern Atlantic The NE Atlantic exhibits greater seasonality in stratification, nutrient depletion, and NPP than the Southern Ocean, and it has a strong spring bloom in the region along 20°W from 35°N to 50°N, e.g., as documented by the North Atlantic Bloom Experiment (Ducklow and Harris, 1993). Antia et al. (2001) published a major review of deep trap observations including 23 sites north of the equator. Annual deep export fluxes vary considerably from <0.3 to >7.0 g POC m2 yr1; with virtually all of this range occurring within particular regions, e.g., the Nordic Seas and Fram Strait. Here, we focus on the relatively well-studied temperate NE Atlantic from 33°N, 20°W to 55°N, 20°W (Antia et al., 2001). Based on a sub-set of annual or longer trap collections from 1000 m or deeper, for which normalization to 230Th fluxes was possible, Antia et al. (2001) concluded that POC fluxes to >3000 m are relatively constant across this latitudinal range, despite the trend of increasing surface export flux with increasing latitude. To explain this observation, they invoked a more active subsurface biological community at higher latitudes to remineralize a larger fraction of the elevated surface POC export (but see Question 2).

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

295

This postulated latitudinal variation in remineralization length scales depends in part on the 230Th normalization for the deep trap fluxes, and it is also sensitive to the depth range of the trap data selected. Without the 230 Th correction and using only data from traps deeper than 2000 m, Francois et al. (2002) found a slightly greater remineralization length scale at 34°N 20°W than at 48°N 20°W (‘‘b’’ values of 0.93 versus 0.82, see Fig. 6b). However, they obtained spatially invariant ‘‘b’’ values after 230Th correction. Berelson (2001a)

Fig. 6. Global maps of production and export characteristics: (a) annual NPP for 1997–1998, based on the Behrenfeld and Falkowski algorithm (http://web.science.oregonstate.edu/ocean.productivity/); (b) remineralization power-law length scales (‘‘b’’ values) as estimated from the vertical attenuation of POC export flux with depth, superimposed onto a contour plot of sea floor O2 flux (Jahnke, 1996). The colour bar represents oxygen flux (mol O2 m2 y1). Blue symbols denote a b value within the range of 0.6 to 0.8, green 0.9 to 1.0, orange 1.1 to 1.2 and purple <1.2. Red stars denote the O2 flux sites, circles denote the b values from Berelson (2001a), and squares represent the b values from Francois et al. (2002). See text for further discussion. It was not possible to graphically represent other broadscale regional estimates of ‘‘b’’ values. These include: (Usbeck et al., 2002) S. Atlantic PFZ > 2, Weddell Sea 1–2; (Schlitzer, 2002) S. Ocean south of 50°S average 1.04, std deviation 0.12; (Antia et al., 2001) N. Atlantic, average for sites 33°N–54°N 0.68. Note nordic stations (i.e., N of 65°N) presented in Francois et al. (2002) are not presented on this map, their ‘‘b’’ values range from 1.11 to 2.01 (n = 7).

296

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

included the fluxes derived from shallower traps (as shallow as 150 m) and suggested stronger remineralization (i.e., ‘‘b’’ of 1.28, Fig. 6b) for the 48°N 20°W site. Lutz et al. (2002) included the 34°N and 48°N 20°W trap data in their global examination of regional variations in export. Their analysis was based on fitting a sum of non-labile and labile (first order with depth) particle components and suggested that POC remineralization was relatively high at subsurface depths in the NE Atlantic, and deep export flux to abyssal depths was correspondingly low. Francois et al. (2002) found the NE Atlantic data to be similar to the global median in terms of POC export to >2000 m depth expressed as a fraction of surface POC export flux. To a large degree these studies offer consensus regarding deep ocean POC fluxes in the NE Atlantic – they are close to global median values and do not vary more than 2-fold along the 20°W meridian from 30 to 55°N. The studies also suggest that export to depth (>2000 m) as a fraction of surface export flux may decrease at higher latitudes in this region. However, this decrease is small in comparison to uncertainties in trap collection efficiencies and in surface export flux estimates, and to interannual variability. The apparent decrease in efficiency of deep (>3000 m) POC export, as a fraction of NPP, with increasing latitude appears to be equally well correlated with either (a) the northward increase in the degree of seasonality in the NPP signal (i.e., the low efficiency is driven by increasing surface NPP); or (b) the northward increase in the opal/carbonate ratio of the deep export fluxes (i.e the low efficiency is driven by the lower POC ‘‘carrying capacity’’ of opal-rich in comparison to carbonate-rich materials; Antia et al., 2001; Francois et al., 2002). Interestingly, the opal/carbonate ratio of deep fluxes also shows interannual variations; Antia et al. (2001) demonstrated the ratio decreased between 1989 and 1996 in deep (>3000 m) trap data from 34°N and 48°N. The decrease in opal was accompanied by a decrease in POC flux. Thus, the temporal opal/carbonate flux ratio variations at each site appear to correlate with deep POC flux in the opposite sense to that suggested by the latitudinal variations. This conclusion is also supported from analysis of the multi-year time-series of deep trap fluxes at OSP in the NE Pacific in comparison to global geographic variations (Table 3). It is useful to keep in mind that in the NE Atlantic temporal variations in opal/carbonate ratios have been observed over only periods of only a few years, and that methodological differences can make their confirmation difficult. In this regard, the apparent temporal decrease in opal/carbonate ratios near Bermuda, suggested by Deuser et al. (1995) and cited as similar to NABE and German JGOFS results at 34°N and 48°N by Antia et al. (2001), has since been called into question following re-evaluation of the Bermuda results by Conte et al. (2001). 5.3. Comparing global maps of export properties Over the last decade the there have efforts to produce global maps of a range of properties relevant to export flux, including NPP (e.g., Behrenfeld and Falkowski, 1997) nutrient concentrations (Louanchi and Najjar, 2000), PIC distributions (Milliman, 1993), modelled f ratio and surface export fluxes (e.g., Laws et al., 2000), deep ocean export flux at 2000 m (e.g., Lampitt and Antia, 1997) and benthic oxygen fluxes (Jahnke, 1996). Trap fluxes have also been combined with the surface export estimates to examine the geographic distributions of export flux attenuation with depth in terms of remineralization power-law ‘‘b’’ values (Francois et al., 2002; Lutz et al., 2002). Sampling coverage in the surface ocean is greatest, with satellite-derived measurements providing estimates on weekly or shorter timescales (Fig. 6a). Nutrient distributions are the next best sampled, but estimates of particle fluxes from them represent averages over at least annual timescales (Schlitzer, 2002) and are subject to considerable uncertainties related to circulation models (Matear and Holloway, 1995). Benthic sedimentary oxygen flux measurements offer the next best coverage (Jahnke, 1996; Fig. 6b), followed by direct flux measurements from traps in the deep sea (Lampitt and Antia, 1997; Antia et al., 2001). Measurements in the mesopelagic are sparse (Berelson, 2001a), and surface export (upper 100 m) is the most poorly sampled of all, with very few time-series of surface export flux that span the annual cycle in sufficient detail to provide reliable annual estimates for global mapping (see Buesseler, 2004). Comparison of the global NPP map (Fig. 6a) with that of benthic oxygen flux (Fig. 6b) shows smaller spatial gradients at depth, with the NPP 10-fold variations reduced to 4-fold variations in the oxygen demand. The maps also show different patterns, with an emphasis on gradients from north to south for NPP and greater east–west trends in the oxygen demand map often paralleling ocean bathymetry. These differences suggest that flux attenuation over the water column must vary from place to place to explain the differences.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

297

However, no distinct regional trends in attenuation have so far emerged from power-law attenuation coefficients estimated from deep trap studies (Fig. 6b), e.g., Equatorial Pacific mean ‘‘b’’ = 0.81 (n = 12); Indian Ocean mean ‘‘b’’ = 0.76 (n = 5). Thus, the observed 2- to 3-fold damping of regional gradients in benthic oxygen fluxes in comparison to NPP estimates may reflect errors in the estimation of surface export fluxes, or other localized influences on deep-sea oxygen demand. In that regard, Lampitt and Antia (1997) reported good correlation of deep ocean carbon fluxes to the benthic O2 fluxes of Jahnke (1996) for the Pacific, but markedly poorer correlation for the Indian Ocean. Lampitt and Antia suggested the pronounced influence of riverine input on sediments in the Indian Ocean as the likely explanation for these differences. Based on these maps, and on the insights obtained from our regional case studies, it seems that separating the influences of near surface (export production), mesopelagic (flux attenuation) and deep-sea (sediment redistributions) influences on the estimation of carbon export to depth remains problematic. Thus, all global maps should be viewed with great caution. This perspective is further reinforced by consideration of the few available decadal time-series of deep trap fluxes. Trends of POC flux with biogenic mineral composition differ from trends extracted from geographic variations (Table 3, Question 3 above). Also, there are occasional years with fluxes that vary nearly as far from the mean (e.g., 3-fold at OSP; Wong et al., 1999) as the global geographic range in sedimentary oxygen demand (4-fold; Jahnke, 1996). The seasonally resolved records from these and other time-series sites also emphasize the shortcomings of the models based on the characteristics of annual mean fluxes. For example, the appearance of abundant phytodetritus (i.e., rapidly sinking, fresh material with little bacterial remineralization at depths >3000 m) following the spring bloom (Lampitt, 1985; Passow, 2004) is at odds with models of mineral control of POC flux to depth based on annual averages (Armstrong et al., 2002; Hedges et al., 2000). In summary, progress in defining global export requires attention to regional differences, including the seasonality of export processes. Time-series sites are essential to this effort, and we suggest that in light of interannual variability there is a need to commit to more long-term time-series. Our review of the global maps and their uncertainties suggests that models of export should aim less at the reproduction of generalized global characteristics from data syntheses, and more at the simulation of key regional and seasonal characteristics. For example, these might include high f-ratios in the Antarctic Circumpolar Current in the South Pacific Basin (Sambrotto and Mace, 2000), low seasonality in surface POC export in the eastern subtropical Pacific despite seasonal variations in N sources from upwelled nitrate and from nitrogen-fixation (Dore et al., 2002), and large interannual variability in POC export to depth dominated by silicifiers at OSP (Wong et al., 1999). This approach, rather than fitting models to overly simplified syntheses of field data (such as the correlations of mean annual POC flux with mineral fluxes) will also ensure that ocean biogeochemical models make the most of the field observations in developing their extrapolations to past and future environmental conditions. 6. Question 5. How far can we simplify the parameterization of export in global circulation models? This section focuses on a key question raised during the US-JGOFS Synthesis and Modelling phase: how far can we simplify the parameterization of export in GCM’s so that they ‘‘capture the essence of C cycle processes compactly enough to be useable in large-scale models’’ (Sarmiento and Armstrong, 1997)? The answer clearly depends on the tasks and timescales we want to address. Here, we provide a brief overview of global models of surface export flux and compare their estimates of global magnitudes and regional patterns of export flux. We also consider whether there are sufficient data to validate these estimates. The formulation of export in these models is clearly important to our ability to predict ocean export in the future (see Question 6). 6.1. Comparison of export from models of varying biological complexity A wide range of models of surface export flux have been developed over the past decade, including (i) diagnostic ‘‘nutrient-inversion’’ models that seek best-fits to the physico-chemical property distributions obtained during the World Ocean Circulation Experiment (WOCE) and other hydrographic programs (Ganachaud and Wunsch, 2002; Howard et al., 2006; Schlitzer, 2002); (ii) semi-prognostic models that combine estimates of NPP from satellite remote sensing with parameterizations of its fractional export (f- or e-ratios) from observations or

298

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

models (e.g., Dunne et al., 2005; Laws et al., 2000). Such models can in turn be coupled to circulation models (e.g., Gnanadesikan et al., 2004); and (iii) fully prognostic models that estimate export from parameterizations of NPP coupled with either export estimates from f-ratio parameterizations, simple ‘‘NPZ’’ biological models (e.g., Cox et al., 2000), or more complex models with representations of food-web and particle dynamics (e.g., Moore et al., 2002a,b). Depending on their complexity, these models can be run ‘‘on-line’’ within GCM’s, or ‘‘off-line’’ using transports derived separately from the GCM’s. Here, rather than comprehensively reviewing these three approaches, we restrict ourselves to illustrating some of the characteristics and uncertainties important to applying such models within carbon cycle simulations. Table 4 lists an illustrative selection of predictions of surface export estimates. The range in global export to below 100 m is relatively small (7.9–13.1 GTC/yr). A similar range of ±40% was obtained from 13 models with the same biogeochemical parameterization, but differing physical circulation (Doney et al., 2004). The similarity of the export estimates is somewhat surprising given the diversity of approaches presented in Table 4, and suggests that either the problem has a relatively unique solution, or all models are making similar approximations. Significantly, in all cases considered, modelled estimates of NPP were within <30% of one another, and surface export flux is generally parameterized as either a constant or at least a monotonically proportional fraction of NPP. This is one probable contribution to the observed convergence in global surface export fluxes. Another is that model export estimates are generally either directly determined by surface and deep nutrient distributions (e.g., Schlitzer, 2002, 2004), ‘‘relaxed’’ to reproduce them (e.g., Sarmiento and LeQuere, 1996), or include parameters selected so that ‘reasonable’ nutrient distributions occur (see discussion in Doney et al., 2004 on the OCMIP-II intercomparison). The effect on export estimates of circulation-driven variations among the models in nutrient supply to surface waters is smoothed out to some degree by the general assumption that all nutrients supplied to surface waters are exported (except in the relatively small regions of mode, intermediate and deep water formation). Thus, global export estimates can be relatively similar despite much larger variations in the range of surface export among different regions (e.g., Gnanadesikan et al., 2004). 6.2. Regional trends in export from two models with differing biological complexity A further useful check of model performance is to compare regional trends in predicted surface export flux. Summaries of modelling intercomparisons such as OCMIP (Doney et al., 2004; Orr, 1999, 2002) and the IPCC (2001) reveal that GCM’s display significant disparities at regional scales, for many of reasons ranging from circulation effects to differing degrees of biological complexity. Here we focus instead on the similarity of regional trends in surface export flux for the two most biologically complex models in Table 4, neither of which is run ‘‘on-line’’ in a GCM. We explore whether the models of Laws et al. (2000) and Moore et al. (2002a,b) obtain consensus on the essential variables for the parameterization of surface export fluxes that could then be used in GCM’s. A prominent example of the similarity of surface export fluxes for the two models is the simulation of relatively high surface export in the open Southern Ocean. This occurs despite significant differences in model parameterization. Both models have compartments for large and small phytoplankton and are structured to produce greater surface export flux when the fractional abundance of large phytoplankton increases. However, the controls on NPP, the proportion of large phytoplankton, and their export are quite different. Importantly, this difference includes the role of temperature in the control of surface export flux. Laws et al. (2000) developed a 10 compartment ecosystem model which produces small and large phytoplankton. It routes the biomass of small phytoplankton through more trophic levels than for large phytoplankton, before export of both as detritus. Increased NPP (driven by increased nutrient loading) leads to greater total export (because steady-state is assumed), and also to a larger fraction of NPP passing through the shorter large phytoplankton pathway as the result of the optimization of a criterion for food-web resistance to perturbation. All production and grazing rates in the model are temperature-dependent, with grazing set to decrease at low temperatures more rapidly than production, so that higher export (as a fraction of NPP) occurs at lower temperature, and thus at high latitudes such as the Southern Ocean. The Moore et al. (2002b) model is more complex than that of Laws et al. (2000). For example, it divides its large phytoplankton compartment into three, uses multi-element rather than single element nutrient limitation

Table 4 Global surface export flux estimates for a range of models with differing complexity Model

Reference

Complexity (particle production, limiting nutrient)

Complexity (particle export)

NPP (Gt C yr1)

Nutrient Inversion (of P, O2, DIC, etc.)

(Schlitzer, 2002) (Laws et al., 2000)

Redfield stoichiometry

Regionally-varying remineralization power-law f-ratio

n.a.

9.6

52.1f

20.9f

(Laws et al., 2000) (Laws et al., 2000)

NPP, SST mapsa

52.1

12.9

52.1

11.1

COAM-AG (OPAICE-ARPEGE)g Prognostic (COAM (NCAR) and offline ecosystem model) Prognostic (COAM (NCAR) and offline ecosystem model) 13 models in OCMIP-II

(Bopp et al., 2001) (Bopp et al., 2001) (Bopp et al., 2001) (Moore et al., 2002a,b) (Moore et al., 2002a,b) (Doney et al., 2004)

Remote sensing NPP mapsa

NPP, SST maps, 10-component food-web,b N is limiting HAMOCC3 (Maier-Reimer, 1993) no explicit biology; nutrient restoration (P) PZDN (Aumont, 1998) one phyto. class; P is limiting HAMOCC3 (Maier-Reimer, 1993) no explicit biology; nutrient restoration (P) 11 compartment model,c three algal groups, potential N, P, Si, Fe limitation of phytoplankton growth Ditto See Doney et al. (2004)

Regression of temperature versus observed ‘‘ef’’ ratio Food-web model temperature dependence of grazing, remineralization, etc. Remineralization based on powerlaw Remineralization power-law

n.a.

13.1

64.7

11.1

Remineralization power-law

n.a.

9.5

Two detrital components, aggregation, remineralization. temperature dependence Ditto

45.2

7.9d

45.2

12.0e

Power-law with fixed length scale

n.a.

±40% range

n.a. Denotes not available. See the cited references for model acronyms and further details. a NPP from (Behrenfeld and Falkowski, 1997), SST from AVHRR. b Algal growth driven by N, and/or temperature (after Eppley, 1972). c Algal growth driven by irradiance, nutrients, trace metals and/or temperature (after Geider et al., 1998). d Sinking particulate C export only and does not include circulation or zooplankton mediated export. e Total biogenic export, including circulation and zooplankton mediated export. f This elevated estimate of surface export from Eppley and Peterson is from a parameterization based on lab studies and is no longer favoured (Laws et al., 2000). g OAICE denotes the oceanic model Ocean Parellelise; ARPEGE and LMD5 represent two alternate atmospheric components of the model (see Bopp et al., 2001) for more details.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Semi-prognostic (Eppley and Peterson f ratio versus temperature relationship) Semi-prognostic (Temperature and Export ratio) Semi-prognostic (Export ratio, NPP and temperature – scaled using remote sensing dataa) Coupled Ocean Atmosphere Model (COAM)-LG (OPAICE – LMDS) COAM-LB (OPAICE – LMDS)g

Surface export (Gt C yr1)

299

300

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

and parameterizes algal aggregation and particle sinking. Large phytoplankton still play a dominant role in export, as a result of direct parameterization of both lower grazing on larger phytoplankton and greater export of sinking detrital material from large phytoplankton. However, the dependence of surface export flux on temperature arises from a directly parameterized decrease in the remineralization rate of detritus at colder temperatures, rather than from a decrease in the efficacy of grazing. The overall dependence of the surface export flux (as a fraction of NPP) on temperature is much weaker than that in the Laws et al. (2000) model (see Fig. 13 of Moore et al., 2002a). Nonetheless, the Moore et al. (2002a) model also obtains high surface export flux (as a fraction of NPP) in the Southern Ocean, because the lower temperature effect is offset by greater growth of large phytoplankton (in particular diatoms) in response to high silicic acid concentrations in Southern Ocean waters. High surface export flux in the Southern Ocean is also achieved in a third model (Armstrong et al., 2002; Dunne et al., 2005) that again parameterizes lower grazing pressure on large phytoplankton but also assigns all mineral ballast production to large phytoplankton and assumes that this retards remineralization and enhances POC export. There is a marked difference in the classification of algal functional groups among these models, e.g., calcifying coccolithophores are treated as ‘‘small’’ phytoplankton by Moore et al. (2002a,b) but as ‘large’ phytoplankton by Dunne et al. (2005). In summary, the complexity of these models differs dramatically, yet they produce similar estimates of surface export flux (as a fraction of NPP) globally and in the Southern Ocean (although regional differences occur elsewhere: see discussion in Moore et al., 2002a). In simple terms, the Laws et al. (2000) model explains regional variations in surface export flux as the result of temperature controls. In contrast, Moore et al. (2002a,b) favours an important role for multi-element nutrient limitation, which then depends on the decoupling of oceanic cycles of biogenic C and Si (and also Fe). Thus, while simulations of global surface export by these two models suggest consensus, they are driven by different controlling factors. This has pronounced implications for predicting how climate change will alter surface export fluxes (see Question 6). This brief discussion illustrates that different hypotheses can be successfully parameterized and included in global ocean carbon cycle models to reproduce broad-scale characteristics such as NPP, surface export fluxes, nutrient and chlorophyll distributions. Choosing the most appropriate parameterization requires more detailed comparisons to observations. For example, the Laws et al. (2000) model has been shown to have difficulty in reproducing seasonal cycles at the Hawaii Ocean Time-series (HOT) site (Laws, 2004), and for this reason very detailed models are perhaps more amenable to robust testing than are more generalized models. Finally, we note that no models have yet included sufficient complexity to capture the observed variability of export fluxes. Determining which additional factors, beyond those of temperature, chlorophyll and NPP, are, most critical is a high priority task. 6.3. How robust is the validation of global export models? To answer the question posed by Sarmiento and Armstrong (1997) as to how far we can simplify model parameterization, we require model validation. Both Laws et al. (2000) and Moore et al. (2002a,b) used JGOFS regional data to validate model outputs such as NPP, chlorophyll and nutrient concentrations. Laws et al. (2000) report good agreement between modelled surface export flux (as a fraction of NPP) and observed mean f-ratios (from 14C and 15N, mainly from short-term studies such as the NABE) for regions ranging from polar to tropical waters. Recently, Laws (2004) conducted a more detailed comparison of new production over an annual cycle at the HOT site (i.e., one of nine sites used in the validation of Laws et al., 2000). The Laws et al. (2000) model under-estimated new production at values of NPP > 70 mmol C m2 d1, and Laws (2004) suggested that this ‘discrepancy reflects the inability of a steady-state model to describe a system that is at times highly dynamic’’. Moore et al. (2002a) compared their model outputs with observations that span seasonal cycles from nine JGOFS time-series sites. The simulated seasonal cycles generally compared favourably for properties including mixed layer depth, chlorophyll concentrations, and NPP. They also compared simulated surface POC export with the few observations that were available. Two key points emerge from the effort to validate these models that are parameterized to estimate surface export flux globally using JGOFS regional datasets. First, a synthesis of surface export fluxes (from 234Th)

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

301

from the JGOFS (Buesseler, 2004), reveals insufficient seasonal coverage (with the exception of HOT) to compare with modelled estimates of surface export. Thus, the annual cycle is insufficiently characterized to be fully useful in model evaluation. As previously stated, this paucity of data stems from the difficulties in measuring surface export flux for time periods longer than days due to the need for a shipboard platform for drifting trap or 234Th studies. Measurements of flux in the subsurface ocean are even more sparse, as are data on the drivers of particle remineralization (see Question 2). The deep ocean moored trap record is somewhat more extensive, at least in terms of providing full seasonal cycles, though geographic coverage is still sparse and uneven (Fig. 6b). Second, the selection of representative observations for model validation plays an influential role in determining the conclusions from modelling studies. For example, Moore et al. (2002a) comment that the high latitude sites (NE Water Polyna, and Ross Sea) used by Laws et al. (2000) were both characterized by high surface export and high nutrient concentrations and are not broadly representative of high latitude waters. This role of site selection extends to the relative influence of spatial variations and temporal variations on model parameterizations. As demonstrated under Question 3, estimation of the influence of mineral ballasts on deep POC export differs between regressions based on globally distributed data and regressions based on interannual variability at selected sites. Similarly, it has been demonstrated that there is no strong relationship between NPP and surface export fluxes over interannual timescales at many individual sites (Bishop, 1989; Boyd and Newton, 1999; Dunne et al., 2005), including time-series such as HOT (Karl et al., 1996). Yet, NPP does appear to exert a strong influence in setting surface export fluxes when globally distributed sites are compared (Dunne et al., 2005; Laws et al., 2000). Ideally, models should be able to reproduce export signatures on a range of timescales, including seasonal, interannual, decadal (shifts associated with atmospheric-ocean ‘‘oscillations’’), and centennial to millennial changes in overturning circulation. However, sufficient data to calibrate models on all these scales are badly lacking. In addition to designing models to address problems, their calibration against data needs careful thought. For example, the relationships between surface export flux and biogeochemical variables (e.g., chlorophyll or mineral ballasts) differ across versus within geographic regions. This suggests that an approach based on a mosaic of regional responses for short time-scales (years) coupled with estimation of the migration of regional boundaries on longer timescales (decades) may be a fruitful alternative. The definition of regional ecosystem characteristics important to export has progressed considerably in the past decade (e.g., Boyd and Doney, 2002; Longhurst, 1998), including efforts to define regions in terms of physical properties alone (Sarmiento et al., 2004). However, more studies of the controls on regional boundaries are urgently needed, and perhaps have been disadvantaged by the conflicting need to establish time-series at sites far from regional boundaries. 7. Question 6. Can we predict future changes in downward biogenic export with any certainty? 7.1. How well do we understand the present controls on export? Over the last decade, our appreciation of the myriad processes that influence export fluxes has increased considerably. We now view both the surface and subsurface ocean as playing key roles in modifying sinking particles. However, there is not yet sufficient knowledge even to rank the specific controls on export (e.g., grazer transformations of particles in the surface ocean, subsurface remineralization rates, or particle sinking rates) in order of importance (Boyd and Stevens, 2002). In contrast, a comparison of global predictions of surface export flux from the subset of models we considered (Table 4) shows good agreement, concurrence that seems to be independent of model complexity. Further analysis is required to understand the reasons for these contrasting trends in process-oriented versus broader-scale global syntheses and modelling studies. However, we can use these two categories of information on surface export to provide different perspectives and insights on how export fluxes might be modified by climate change. Process studies will provide a check-list to assess how the changes to ocean properties predicted by modelling (e.g., Cox et al., 2000; Matear and Hirst, 1999; Sarmiento et al., 1998) will influence individual processes, e.g., by altering calcification (Heinze, 2004; Riebesell et al., 2000; Schulz et al., 2004) or silicification (Milligan et al., 2003) with subsequent impacts on ballasting. In contrast, global syntheses provide estimates

302

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

of key integrative indices such as the ratio of surface export flux to NPP, ‘‘b’’ values, etc, that can be explored with respect to interannual variability within each region, to regional differences, or to extreme years (see Fig. 5) that are often driven by climate variability such as ENSO (Behrenfeld et al., 2001). These will help constrain the bounds on export fluxes and how they might be altered. 7.2. How are ocean properties predicted to change and how will they impact export? Projections of how properties will alter are provided both from global modelling experiments and by extrapolation of ongoing time-series observations, such as increased pCO2 (Dore et al., 2003), and decreased pH (Sabine et al., 2004). There is consensus from modelling experiments (Joos et al., 1999; Matear and Hirst, 1999; Sarmiento et al., 1998) that climate change will lead to warming, increased stratification, and shoaling of the surface mixed-layer over the coming decades. Given that these concurrent changes will alter underwater light climate, the supply of nutrients and trace metals and the ocean’s carbonate chemistry, it is difficult to predict how phytoplankton – which produce the ‘seed stocks’ of most sinking particles – will respond. But clearly changes to global NPP and/or floristic shifts between taxa, such as from diatoms to coccolithophores, would alter surface export fluxes globally. Only one modelling study has so far attempted to forecast regional changes in surface export due to climate change (Bopp et al., 2001). On the basis of changes to upper ocean light climate and nutrient fields, Bopp et al. predict regional variations in surface export in the future ranging from 15% (tropics – zonal average) to +10% for the Southern Ocean. However, large uncertainties are evident as to which biogeochemical provinces would be most influenced, or their boundaries most altered, by climate change. The IPCC (2001) review of global modelling experiments emphasizes that the largest disparities between model predictions and future changes are likely to occur at regional scale. Thus, until there are more modelling studies on the effects of climate change on export, and in particular on potential biotic feedbacks (Boyd and Doney, 2003), we are unable to predict how surface export fluxes might alter in different provinces over the coming decades. The results from modelling experiments do indicate that the surface ocean is likely to see larger changes in physico-chemical properties than the subsurface ocean. This points to the upper ocean as a key focus for future research, although changes there may be either damped or amplified by subsurface processes. 7.3. How much might export fluxes change – can we set biogeochemical bounds? Despite the existence of uncertainties at the process-level as to how climate change might alter surface export, it is possible to determine some probable biogeochemical bounds from: (a) the geological record (e.g., Anderson et al., 2002; Chase et al., 2001; Kumar et al., 1995); (b) studies of extreme years in trap records (e.g., the 1984 ENSO event at OSP, Fig. 5a); (c) consideration of how nutrient utilisation in the ocean might change (Bopp et al., 2001; Gnanadesikan et al., 2004; Matear and Elliot, 2004; Sarmiento and LeQuere, 1996). Although there are assumptions included in the use of paleo-proxies to define export fluxes, paleo-oceanographic studies of the Southern Ocean suggest that up to 2- to 4-fold higher export fluxes have occurred there (e.g., Francois et al., 1997; Kumar et al., 1995; Lourey et al., 2004). The deep export fluxes during the extreme year at OSP were 2-fold higher than the typical annual flux (Fig. 5a). Regardless of the timescales of change for the above examples (years to centuries), altered export fluxes were probably driven by either enhanced nutrient supply (Rothlisberger et al., 2004), increased ability to take up nutrients (e.g., due to the alleviation of iron limitation), or a change in nutrient uptake stoichiometry (see Denman et al., 1996). Modelling experiments indicate climate change in the future will enhance water column stratification, leading to reduced nutrient supply. This in turn will decrease new production and hence reduce surface export flux by 6% globally (Bopp et al., 2001). Other modelling experiments, in which surface wind-stresses are reduced, indicate the importance of upwelled nutrient supply on setting both NPP and surface export flux (Palmer and Totterdell, 2001). If climate change results in enhanced regional nutrient or trace metal supply, some model simulations point to increased calcite production (Matear and Elliot, 2004) based on an increased macronutrient inventory via ‘fertilisation’. Increased iron supply may result in enhanced diatom blooms with higher surface export flux (Gnanadesikan et al., 2003), but only for a limited, decadal timescale based on simulations for the Equatorial Pacific.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

303

Thus, these approaches suggest that up to a halving or doubling of surface export fluxes might be geochemically possible on decadal timescales i.e., that are relevant to climate change. However, interannual variations in atmospheric carbon dioxide and oxygen concentrations and isotopic compositions suggest that the total variability in ocean–atmosphere gas exchange is also of this order. Most of this variability has been attributed to thermal and circulation effects (Battle et al., 2000; Keeling and Shertz, 1992; Keeling et al., 1996), perhaps with countervailing changes among different regions (Le Que´re´ et al., 2003; Peylin et al., 2005). Whether the limits on export variability inferred from such short records can be applied to future predictions is a subject of active investigation. Clearly, there are many unknowns at the process level, which will impact the magnitude and sign of feedbacks to climate change and hence the timescales of change. These include changes in the areal extent of biomes (Boyd and Doney, 2002), regional shifts in the production of mineral ballast (opal:calcite) in the surface ocean, responses of zooplankton (Richardson and Shoemann, 2004), and shifts in the length of the algal growth season (Reid et al., 1998). 8. Is there consensus? There is general agreement that the main processes important to export have been identified (Table 1), though of course surprises could still occur. To a large degree these processes were known at the time of the Dahlem Conference (Berger et al., 1989) though their details have now been refined. The major advances of the past decade have been the expansion of the geographical coverage and the lengthening of times-series of observations relevant to export assessment. Synthesis of these results has provided an unprecedented view of the links between POC export and other aspects of marine ecosystems. There is consensus on many aspects of the phenomenology of export. For example, that on global scales POC export to the deep sea correlates with mineral fluxes, that POC export from the surface ocean correlates with significant production by large phytoplankton, and that flux attenuation is faster in the subsurface ocean than at great depth. Other broad-brush views are widely held and probably correct: that the composition of export varies geographically from opal to carbonate to organic matter domination, and that this variation correlates with food-web structures. There is also growing concern, as emphasized in this review, that the generality of these correlations at global spatial scales and mean annual timescales does not hold for smaller variations that occur within individual oceans, between years, or over the course of seasonal cycles. There is much less consensus about the causality of the correlations, or of high POC export itself. For example, opinions differ on whether the correlation of deep POC export with mineral fluxes is driven by the ballast effect of the minerals, or the ability of the organic matter to aggregate the minerals (Armstrong et al., 2002, cf. Passow, 2004; Passow and De La Rocha, 2006). These differing perspectives in part reflect the scope of applicability of models of mineral control of POC export – formulated from annual mean fluxes in the deep sea, but unable to accommodate seasonal variations or shallower observations. Some have argued that the correlations and associated models still offer predictive value if applied within their narrow scope, but there is concern that the statistics of these correlations will not be static in a changing ocean. Such multiple explanations for the observed correlations are common, e.g., the association of large phytoplankton with high export can be explained by escape from grazing pressure, and/or the formation of organic rich phytodetrital aggregates that sink rapidly. We believe that a global view of export sufficient to offer predictive capabilities will require a regionalisation of the parameterization of export processes, and additional complexity in the processes simulated. The need for regionalisation is evident in the recognition that subsurface flux attenuation varies geographically, beyond that which a single power-law parameterization can capture (Berelson, 2001a; Francois et al., 2002; Howard et al., 2006, this review). The process of regionalisation has begun (e.g., Boyd and Doney, 2002; Le Que´re´ et al., 2005; Moore et al., 2002a) echoing the more advanced state of understanding within terrestrial ecology. There has been some resistance toward more complex models, and reasonably so, because of the lack of observations to adequately test and validate them (Doney, 1999), and because of the need for computational economy (Sarmiento and Armstrong, 1997). After the last decade of advances, computing restrictions have eased, and paucity of observations is the largest limitation on progress. This is the greatest consensus to emerge from

304

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

our review – the ocean remains dramatically under-sampled in space and time. To address this daunting task we offer a final section of recommendations on key areas requiring progress. 9. Recommendations for future work (1) Despite improvements in spatial and temporal coverage of export observations, the paucity of results from high latitude regions and the S Hemisphere (such as the Southern Ocean and SW Pacific) remains evident. We recommend an effort to address this imbalance in geographical coverage, guided by the identification of a minimum set of ‘export biogeochemical provinces’ that contains the lower or upper bounds of key parameters, such as NPP, temperature, seasonality, mineral flux, aerosol supply, diatoms versus coccolithophores, subsurface temperature, zooplankton characteristics. Establishment of time-series programs within these regions and their continuation elsewhere are essential. We also need to elucidate better the physical and chemical characteristics that define the boundaries of these regions. (2) There is a need for development, validation, and widespread application of new technologies to overcome under-sampling issues. For example, the C-Explorer profiling float – which provides POC profiles (via transmissivity and telemetry, Bishop et al., 2002) and can be used as an optical sediment trap (Bishop et al., 2004) could greatly increase the number of time-series of surface export flux over the annual cycle, and provide comparisons with both NPP (satellite-derived) and deep export fluxes. (3) There is a need to focus on export from surface waters and through the subsurface ocean. This is where most flux attenuation occurs and represents the depth range of importance for prediction of carbon cycle changes over the coming decades and centuries. Export studies have been badly hampered by the difficulty of obtaining full annual time-series at the surface and by methodological problems with the collection of sinking particles in the presence of strong currents and actively migrating and feeding zooplankton. Promising developments include neutrally buoyant sediment traps (NBSTs, e.g., Buesseler et al., 2000), traps with swimmer exclusion valves (Peterson et al., 1993), and methods to determine particle sinking velocities in situ (Peterson et al., 2005). (4) Research into the biochemistry and ecology of the processes responsible for remineralization is in its infancy, but progress is accelerating on several fronts. These include bacterial community structure, remineralization experiments utilising radio- and isotope-labelling approaches, organic geochemical studies of changes in particle composition, and studies of the correlation of chemical properties with particle sinking rates. Major limitations of this work are imposed by the compromises made in collecting and bringing the samples to the surface. To make further progress, there is a critical need to devise new means to examine remineralization processes in situ. (5) Remote sensing from satellites and in situ sensors on profiling floats also offer great promise to overcome poor sampling resolution in the upper ocean via efforts to retrieve POC and PIC (ballast) fields (Balch et al., 2001; Bishop et al., 2002; Stramski, 1999). For example the relative role of coccolithophores in global calcite production could be estimated better using PIC fields in conjunction with algorithms to derive coccolithophore concentrations (Brown and Yoder, 1994). These remote observational approaches require co-ordinated large-scale calibration programs. (6) Progress in all these areas presents the combined challenge of an increasing need for resolution of biogeochemical transformations in terms of the multitude of biological processes that mediate them. Thus, there is the need for a wide range of specialised expertise, with the need to bring this diversity together in coherent field programs. An obvious example is the separation of carbonate into phytoplankton and zooplankton sources. Large field programs are needed, but perhaps at fewer sites than during the broad scoping years of the JGOFS program, but with perhaps more measurements that integrate physics, biology, biochemistry, geochemistry and ecology. Finally, the scope of export studies required for predictions of global change extends beyond the open ocean processes reviewed here. Clearly, the next generation of programs must also address the role of the ocean shelf in carbon delivery to the deep sea, and begin to address their historically convenient but increasingly artificial separation.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

305

Acknowledgements We thank A. Antia, D. Caron, M. Conte, J.P. Dunne, M. Lutz, R. Matear, D. Montagnes, J.K. Moore, S. Strom and C.S. Wong for discussions and/or access to data or work in progress. We are grateful to R. Francois and R. Jahnke for supplying us with the original data/figures, from publications on global trends in deep export fluxes and benthic oxygen flux, respectively. Erika McKay was responsible for the graphics in this review. This review was significantly improved by the insights and comments from four reviewers. We dedicate this brief illustration of the dim nature of our understanding of ocean biogeochemistry to two people who sought to illuminate it throughout their lives – Jack Dymond and John Hedges. References Abell, J., Emerson, S., Renaud, P., 2000. Distributions of TOP,TON, and TOC in the North Pacific subtropical gyre: implications for nutrient supply in the surface ocean and remineralization in the upper thermocline. Journal of Marine Research 58, 203–222. Abelmann, A., Gowing, M.M., 1997. Spatial distribution of living polycystine radiolarian taxa – baseline study for paleoenvironmental reconstructions in the Southern Ocean (Atlantic sector). Marine Micropaleontology 30, 3–28. Alldredge, A.L., Jackson, G.A., 1995. Aggregation in marine systems. Deep-Sea Research II 42, 1–7. Alldredge, A., Passow, U., Logan, B., 1993. The abundance and significance of a class of large, transparent organic particles in the ocean. Deep-Sea Research I 40, 1131–1140. Anderson, O.R., 1993. The trophic role of planktonic Foraminifera and Radiolaria. Marine Microbial Food Webs 7, 31–51. Anderson, L., 1995. On the hydrogen and oxygen content of marine phytoplankton. Deep-Sea Research I 42, 1675–1680. Anderson, T.R., Pondaven, P., 2003. Non-Redfield carbon and nitrogen cycling in the Sargasso Sea: pelagic imbalances and export flux. Deep-Sea Research I 50, 573–591. Anderson, L., Sarmiento, J., 1994. Redfield ratios of remineralization determined by nutrient data analysis. Global Biogeochemical Cycles 8, 65–80. Anderson, R.F., Chase, Z., Fleischer, M.Q., Sachs, J., 2002. The Southern Ocean’s biological pump during the Last Glacial Maximum. Deep-Sea Research II 49, 1909–1938. Antia, A.N., Koeve, W., Fischer, G., Blanz, T., Schulz-Bull, D., Scholten, J., Neuer, S., Kremling, K., Kuss, J., Peinert, R., Hebbeln, D., Bathmann, U., Conte, M., Fehner, U., Zeitzschel, B., 2001. Basin-wide particulate carbon flux in the Atlantic Ocean: regional export patterns and potential for atmospheric CO2 sequestration. Global Biogeochemical Cycles 15, 845–862. Antoine, D., Andre´, J.-M., Morel, A.A., 1996. Oceanic primary production 2. Estimation at global scale from satellite (coastal zone color scanner) chlorophyll. Global Biogeochemical Cycles 10, 57–70. doi:10.1029/1095GB0283. Armstrong, R.A., Lee, C., Hedges, J.I., Honjo, S., Wakeham, S.G., 2002. A new, mechanistic model for organic carbon fluxes in the ocean based on the quantitative association of POC with ballast minerals. Deep-Sea Research II 49, 219–236. Arrigo, K.R., Worthen, D., Schnell, A., Lizotte, M.P., 1998. Primary production in Southern Ocean waters. Journal of Geophysical Research 103, 15,587–15,600. Arrigo, K.R., Robinson, D.H., Worthen, D.L., Dunbar, R.B., DiTullio, G.R., VanWoert, M., Lizotte, M.P., 1999. Phytoplankton community structure and the drawdown of nutrients and CO2 in the Southern Ocean. Science 283, 365–367. Aumont, O., 1998. Etude du cycle naturel du carbone dans un modele 3D l’ocean mondial, PhD. Thesis, Univ. Pierre et Marie Curie, Paric. Azam, F., 1998. Microbial control of oceanic carbon flux: the plot thickens. Science 280, 694–695. Bacon, M., 1996. Evaluation of sediment traps with naturally occurring radionuclides. In: Ittekkot, V., Schafer, P., Honjo, S., Depetris, P.J. (Eds.), Particle Flux in the Ocean, vol. 57. J. Wiley and Sons, Chichester, pp. 85–90. Balch, W.M., Kilpatrick, K.A., Holligan, P.M., Trees, C., 1996. The 1991 coccolithophore bloom in the central north Atlantic: optical properties and factors affecting their distribution. Limnology and Oceanography 41, 1669–1683. Balch, W.M., Drapeau, D.T., Fritz, J.J., 2001. Optical backscattering in the Arabian Sea – continuous underway measurements of particulate inorganic and organic carbon. Deep Sea Research I, 413–431. Barker, S., Kiefer, T., Elderfield, H., 2004. Temporal changes in North Atlantic circulation constrained by planktonic foraminiferal shell weights. Paleoceanography 19 (PA3008). doi:10.1029/2004PA00100. Barth, J.A., Cowles, T.J., Kosro, P.M., Shearmean, R.K., Huyer, A., Smith, R.L., 2002. Injection of carbon from the shelf to offshore beneath the euphotic zone in the California Current. Journal of Geophysical Research 107. doi:10.1929/2001JC00095. Bates, N.R., Michaels, A.F., Knap, A.H., 1996. Seasonal and interannual variability of oceanic carbon dioxide species at the US JGOFS Bermuda Atlantic Time-series Study (BATS) site. Deep-Sea Research II 43, 347–383. Battle, M., Bender, M.L., Tans, P.P., White, J.W.C., Ellis, J.T., Conway, T., Francey, R.J., 2000. Global carbon sinks and their variability inferred from atmospheric O2 and d13C. Science 287, 2467–2470. Behrenfeld, M.J., Falkowski, P.G., 1997. A consumer’s guide to phytoplankton primary productivity models. Limnology and Oceanography 42, 1479–1491. Behrenfeld, M.J., Randerson, J.T., McClain, C.R., Feldman, G.C., Los, S.O., Tucker, C.J., Falkowski, P.G., Field, C.B., Frouin, R., Esaias, W.E., Kolber, D.D., Pollack, N.H., 2001. Biospheric primary production during an ENSO transition. Science 291, 2594–2597. Berelson, W.M., 2001a. The flux of particulate organic carbon into the ocean interior: a comparison of four US JGOFS regional studies. Oceanography 14, 59–67.

306

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Berelson, W.M., 2001b. Particle settling rates increase with depth in the ocean. Deep-Sea Research II 49, 237–251. Berger, W.H., Smetacek, V.S., Wefer, G., 1989. Ocean productivity and paleoproductivity – an overview. In: Berger, W.H., Smetacek, V.S., Wefer, G. (Eds.), Dahlem Workshop Life Sciences Report – Productivity of the oceans: Present and Past. J. Wiley and Sons, New York, p. 355. Bidle, K.D., Azam, F., 1999. Accelerated dissolution of diatom silica by marine bacterial assemblages. Nature 397, 508–512. Bidle, K.D., Manganelli, M., Azam, F., 2002. Regulation of oceanic silicon and carbon preservation by temperature control on bacteria. Science 298, 1980–1984. Billett, D.S.M., Lampitt, R.S., Rice, A.L., Mantoura, R.F.C., 1983. Seasonal sedimentation of phytoplankton to the deep-sea benthos. Nature 302, 520–522. Bishop, J.K.B., 1989. Regional extremes in particulate matter composition and flux: effects on the chemistry of the ocean interior. In: Berger, W.H., Smetacek, V.S., Wefer, G. (Eds.), Productivity of the Ocean: Present and Past. J. Wiley and Sons, New York, pp. 117– 138. Bishop, J.K.B., Davis, R.E., Sherman, J.T., 2002. Robotic observations of dust storm enhancement of carbon biomass in the North Pacific. Science 298, 817–821. Bishop, J.K.B., Wood, T.J., Davis, R.E., Sherman, J.T., 2004. Robotic observations of enhanced carbon biomass and export at 55°S during SOFeX. Science 304, 417–420. Bisset, W.P., Walsh, J.J., Dieterle, D.A., Carder, K.L., 1999. Carbon cycling in the upper waters of the Sargasso Sea: I. Numerical simulation of differential carbon and nitrogen fluxes. Deep-Sea Research I 46, 205–269. Boltovskoy, D., Alder, V.A., 1992. Paleoecological implications of radiolarian distribution and standing stocks versus accumulation rates in the Weddell Sea. In: Kennett, J.P., Warnke, D.A. (Eds.), The Antarctic Environment: A Perspective on Global Change, vol. 56. American Geophysical Union, pp. 377–384. Bopp, L., Monfray, P., Aumont, O., Dufresne, J.-L., Le Treut, H., Madec, G., Terray, L., Orr, J.C., 2001. Potential impact of climate change on marine export production. Global Biogeochemical Cycles 15, 81–99. Boyd, P.W., Doney, S.C., 2002. Modelling regional responses by marine pelagic ecosystems to global climate change. Geophysical Research Letters. doi:10.1029/2001GL014130. Boyd, P.W., Doney, S.C., 2003. The impact of climate change and feedback processes on the Ocean Carbon Cycle. In: Fasham, M.J.R. (Ed.), Ocean Biogeochemistry – The Role of the Ocean Carbon Cycle in Global Change. Springer-Verlag, Berlin. Boyd, P.W., Newton, P., 1995. Evidence of the potential influence of planktonic community structure on the interannual variability of particulate carbon flux. Deep-Sea Research I 42, 619–639. Boyd, P.W., Newton, P., 1999. Does planktonic community structure determine downward particulate organic carbon flux in different oceanic provinces. Deep-Sea Research I 46, 63–91. Boyd, P.W., Stevens, C.L., 2002. A coupled physical–biological dynamics model to assess factors controlling the transfer of material to the deep ocean. Progress in Oceanography 52, 1–29. Boyd, P.W., Sherry, N.D., Berges, J.A., Bishop, J.K.B., Calvert, S.E., Charette, M.A., Giovannoni, S.J., Goldblatt, R., Harrison, P.J., Moran, S.B., Roy, S., Soon, M., Strom, S., Thibault, D., Vergin, K.L., Whitney, F.A., Wong, C.S., 1999. Transformations of biogenic particulates from the pelagic to the deep ocean realm. Deep-Sea Research II 46, 2761–2792. Boyd, P.W., Law, C.S., Wong, C.S., Nojiri, Y., Tusda, A., Levasseur, M., Takeda, S., Rivkin, R., Harrison, P.J., Strzepek, R., Gower, J., McKay, R.M., Abraham, E., Arychuk, M., Barwell-Clarke, J., Crawford, W., Hale, M., Harada, K., Johnson, K., Kiyosawa, H., Kudo, I., Marchetti, A., Miller, W., Needoba, J., Nishioka, J., Ogawa, H., Page, J., Robert, M., Saito, H., Sastri, A., Sherry, N., Soutar, T., Sutherland, N., Taira, Y., Whitney, F., Wong, S.E., Yoshimura, T., 2004. The decline and fate of an iron-induced subarctic phytoplankton bloom. Nature 428, 549–553. Bradford-Grieve, J.M., Nodder, S.D., Jillett, J.B., Currie, K., Lassey, K.R., 2001. Potential contribution that the copepod Neocalanus tonsus makes to downward carbon flux in the Southern Ocean. Journal of Plankton Research 23, 963–975. ´ lvarez-Salgado, X.A., A ´ lvarez, M., Pe´rez, F.F., Me´mery, L., Mercier, H., Messias, M.J., 2004. Nutrient mineralization rates Brea, S., A and ratios in the eastern South Atlantic. Journal of Geophysical Research—Oceans 109. doi:10.1029/2003JC00205. Brown, C.W., Yoder, J.A., 1994. Coccolithophore blooms in the global ocean. Journal of Geophysical Research C 104, 1541–1558. Buesseler, K.O., 1991. Do upper-ocean sediment traps provide an accurate record of particle flux. Nature 353, 420–423. Buesseler, K.O., 1998. The decoupling of production and particulate export in the surface ocean. Global Biogeochemical Cycles 12, 297– 310. Buesseler, K.O., 2004. US-JGOFS SMP on line report, http://usjgofs.whoi.edu/mzweb/smppi/buesseler.html#PROJECT%20DATA. Buesseler, K.O., Steinberg, D.K., Michaels, A.F., Johnson, R.J., Andrews, J.E., Valdes, J.R., Price, J.F., 2000. A comparison of the quantity and composition of material caught in a neutrally buoyant versus surface-tethered sediment trap. Deep-Sea Research I 47, 277–294. Buesseler, K.O., Ball, L., Andrews, J.E., Cochran, J.K., Hirschberg, D.J., Bacon, M.P., Fleer, A., Brzezinski, M., 2001. Upper ocean export of particulate organic carbon and biogenic silica in the Southern Ocean along 170°W. Deep-Sea Research II 48, 4275–4297. Buitenhuis, E., van der Wal, P., de Baar, H.J.W., 2001. Blooms of Emiliania huxleyi are sinks of atmospheric carbon dioxide: a field and mesocosm study derived simulation. Global Biogeochemical Cycles 15, 577–587. Cardinal, D.B., Savoye, N., Trull, T.W., Andre´, L., Kopczynska, E.E., Dehairs, F., 2004. Variations of carbon remineralisation in the Southern Ocean illustrated by the Baxs proxy. Deep-Sea Research I 52, 355–370. Caron, D.A., Swanberg, N.R., 1990. The ecology of planktonic sarcodines. Aquatic Sciences 3, 147–180. Chase, Z., Anderson, R.F., Fleisher, M.Q., 2001. Evidence from authigenic uranium for increased productivity of the glacial Subantarctic Ocean. Paleoceanography 16, 468–478.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

307

Cho, B.C., Azam, F., 1988. Major role of bacteria in biogeochemical fluxes in the oceans interior. Nature 332, 441–443. Christian, J.R., Karl, D.M., 1995. Bacterial ectoenzymes in marine waters: activity ratios and temperature responses in three oceanographic provinces. Limnology and Oceanography 40, 1042–1049. Christian, J.R., Lewis, M.R., Karl, D.M., 1997. Vertical fluxes of carbon, nitrogen, and phosphorous in the North Pacific Subtropical Gyre near Hawaii. Journal of Geophysical Research—Oceans 102, 15667–15677. Coale, K.H., 1990. Labyrinth of doom: a device to minimise the ‘swimmer’ component in sediment trap collections. Limnology and Oceanography 35, 1376–1380. Coale, K.H., Johnson, K.S., Fitzwater, S.E., Gordon, R.M., Tanner, S., Hunter, C.N., Codispoti, L.A., 2001. The oceanic nitrogen cycle: a double-edged agent of environmental change? In: Bendell-Young, L., Gallagher, P. (Eds.), Waters in Peril. Kluwer Academic Publishers, pp. 73–101. Conte, M., Ralph, N., 1999. Biological reprocessing of particle flux within meso and bathypelagic waters of the Sargasso Sea. Eos Transactions, American Geophysical Union Ocean Sciences Meeting Supplement 80, 142. Conte, M.H., Ralph, N., Ross, E.H., 2001. Seasonal and interannual variability in deep ocean particle fluxes at the Oceanic Flux Program (OFP)/Bermuda Atlantic Time Series (BATS) site in the western Sargasso Sea near Bermuda. Deep-Sea Research II 48, 1471–1505. Copin-Montegut, C., 2000. Consumption and production on scales of a few days of inorganic carbon, nitrate and oxygen by the planktonic community: results of continuous measurements at the Dyfamed Station in the northwestern Mediterranean Sea (May 1995). Deep-Sea Research I 47, 447–477. Cox, P.M., Betts, R.A., Jones, C.D., Spall, S.A., Totterdell, I.J., 2000. Acceleration of global warming due to carbon-cycle feedbacks in a coupled climate model. Nature 408, 184–187. Dam, H.G., Roman, M.R., Youngbluth, M.J., 1995. Downward export of respiratory carbon and dissolved inorganic nitrogen by dielmigrant mesozooplankton at the JGOFS Bermuda time-series station. Deep-Sea Research I 42, 187–1197. Denman, K.L., Pena, M.A., 1999. Beyond JGOFS. In: Hanson, R.B., Ducklow, H.W., Field, J.G. (Eds.), The Changing Ocean Carbon Cycle, Cambridge University Press, pp. 469–492. Denman, K.L., Pena, M.A., 2000. Beyond JGOFS. In: Hanson, R.B., Ducklow, H.W., Field, J.G. (Eds.), The Changing Ocean Carbon Cycle: A Midterm Synthesis of the Joint Global Ocean Flux Study, vol. 5. Cambridge University Press, Cambridge, pp. 469–490. Denman, K., Hofmann, E., Marchant, H., 1996. Marine biotic responses and feedbacks to environmental change and feedbacks to climate. In: Houghton, J.T., Meira Filho, L.G., Callander, B.A., Harris, N., Kattenberg, A., Maskell, K. (Eds.), Climate Change 1995, The Science of Climate Change. Cambridge University Press. Deuser, W.G., Jickells, T.D., King, P., 1995. Decadal and annual changes in biogenic opal and carbonate fluxes to the deep Sargasso Sea. Deep-Sea Research I 42, 1923–1995. Diercks, A.R., Asper, V.L., 1997. In situ settling speeds of marine snow aggregates below the mixed layer: Black Sea and Gulf of Mexico. Deep-Sea Research I 44, 385–398. Dilling, L., Alldredge, A.L., 2000. Fragmentation of marine snow by swimming macrozooplankton: a new process impacting carbon cycling in the sea. Deep-Sea Research I 47, 1227–1245. Doney, S.C., 1999. Major challenges confronting marine biogeochemical modelling. Global Biogeochemical Cycles 13, 705–714. Doney, S.C., Sarmiento, J.L., 1999. Ocean Biogeochemical Response to Climate Change. US JGOFS, Synthesis and Modeling Project Planning Workshop Report. Woods Hole Oceanographic Institution, Woods Hole. Doney, S.C., Lindsay, K., Caldeira, K., Campin, J.-M., Drange, H., Dutay, J.-C., Follows, M., Gao, Y., Gnanadesikan, A., Gruber, N., Ishida, A., Joos, F., Madec, G., Maier-Reimer, E., Marshall, J.C., Matear, R.J., Monfray, P., Mouchet, A., Najjar, R., Orr, J.C., Plattner, G.-K., Sarmiento, J., Schlitzer, R., Slater, R., Totterdell, I.J., Weirig, M.-F., Yamanaka, Y., Yool, A., 2004. Evaluating global ocean carbon models: the importance of realistic physics. Global Biogeochemical Cycles 18. doi:10.1029/2003GB00215. Dore, J.E., Brum, J.R., Tupas, L.M., Karl, D.M., 2002. Seasonal and interannual variability in sources of nitrogen supporting export in the oligotrophic subtropical North Pacific Ocean. Limnology and Oceanography 47, 1595–1607. Dore, J.E., Lukas, R., Sadler, D.W., Karl, D.M., 2003. Climate-driven changes to the atmospheric CO2 sink in the subtropical North Pacific Ocean. Nature 424, 745–757. Duarte, C.M., Agustı´, S., 1998. The CO2 balance of unproductive aquatic ecosystems. Science 281, 234–236. Ducklow, H.W., Harris, R.P., 1993. Introduction to the JGOFS North Atlantic Bloom Experiment. Deep-Sea Research II 40, 1–8. Ducklow, H.W., Steinberg, D.K., Buesseler, K.O., 2001. Upper ocean carbon export and the biological pump. Oceanography 14, 50–58. Dunne, J.P., Armstrong, R.A., Gnanadesikan, A., Sarmiento, J.L., 2005. Empirical and mechanistic models for the particle export ratio. Global Biogeochemical Cycles 19, GB4026. Engel, A., 2004. Distribution of transparent exopolymer particles (TEP) in the northeast Atlantic ocean and their potential significance for aggregation. Deep-Sea Research I 51, 83–92. Eppley, R.W., 1972. Temperature and phytoplankton growth in the sea. Fisheries Bulletin 70, 1063–1085. Falkowski, P.G., Davis, C.S., 2004. Natural proportions. Nature 431, 131. Falkowski, P.G., Barber, R.T., Smetacek, V., 1998. Biogeochemical controls and feedbacks on ocean primary production. Science 281, 200–206. Falkowski, P., Scholes, R.J., Boyle, E.A., Canadell, J., Canfield, D., Elser, J., Gruber, N., Hibbard, K., Hogberg, P., Linder, S., Mackenzie, F.T., Moore, B.I., Pedersen, T., Rosenthal, Y., Seitzinger, S., Smetacek, V., Steffen, W., 2000. The global carbon cycle: a test of our knowledge of Earth as a system. Science 290, 291–296. Feely, R.A., Sabine, C.L., Lee, K., Berelson, W., Kleypas, J., Fabry, V.J., Millero, F.J., 2004. Impact of Anthropogenic CO2 on the CaCO3 System in the Oceans. Science 305, 362–366.

308

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Fischer, G., Ratmeyer, V., Wefer, W., 2000. Organic carbon fluxes in the Atlantic and the Southern Ocean: relationship to primary production compiled from satellite radiometer data. Deep-Sea Research II 47, 1961–1997. Francois, R.F., Altabet, M.A., Yu, E.-F., Sigman, D.M., Bacon, M.P., Franck, M., Bohrman, G., Bareille, G., Labeyrie, L.D., 1997. Water column stratification in the Southern Ocean contributed to the lowering of glacial atmospheric CO2. Nature 1997, 929–935. Francois, R., Honjo, S., Krishfield, R., Manganini, S., 2002. Factors controlling the flux of organic carbon to the bathypelagic zone of the ocean. Global Biogeochemical Cycles 16. doi:10.1029/2001GB00172. Frankignoulle, M., Canon, C., Gattuso, J.-P., 1994. Marine calcification as a source of carbon dioxide: positive feedback of increasing atmospheric CO2. Limnology and Oceanography 39, 458–462. Ganachaud, A., Wunsch, C., 2002. Oceanic nutrient and oxygen transports and bounds on export production during the World Ocean Circulation Experiment. Global Biogeochemical Cycles 16, 1057. Gardner, W.D., 2000. Sediment trap sampling in surface waters. In: Hanson, R.B., Ducklow, H.W., Field, J.G. (Eds.), The Changing Ocean Carbon Cycle: a Midterm Synthesis of the Joint Global Ocean Flux Study, vol. 5. Cambridge University Press, Cambridge, pp. 240–281. Geider, R.J., La Roche, J., 2002. Redfield revisited: variability of C:N:P in marine microalgae and its biochemical basis. European Journal of Phycology 37, 1–17. Geider, R.J., McIntyre, H.L., Kana, T.M., 1998. A dynamic regulatory model of phytoplankton acclimation to light, nutrients and temperature. Limnology and Oceanography 43, 679–694. Gnanadesikan, A., Sarmiento, J.L., Slater, R.D., 2003. Effects of patchy ocean fertilization on atmospheric carbon dioxide and biological production. Global Biogeochemical Cycles 17, 1050. Gnanadesikan, A., Dunne, J.P., Key, R.M., Matsumoto, K., Sarmiento, J.L., Slater, R.D., Swathi, P.S., 2004. Oceanic ventilation and biogeochemical cycling: understanding the physical mechanisms that produce realistic distributions of tracers and productivity. Global Biogeochemical Cycles 18. doi:10.1029/2003GB00209. Hedges, J.I., Eglinton, G., Hatcher, P.G., Kirchman, D.L., Arnosti, C., Derenne, S., Evershed, R.P., ogel-Knabner, I.K., de Leeuw, J.W., Littke, R., Michaelis, W., Rullkotter, J., 2000. The molecularly uncharacterized component of nonliving organic matter in natural environments. Organic Geochemistry 31, 945–958. Heinze, C., 2004. Simulating oceanic CaCO3 export production in the greenhouse. Geophysical Research Letters 31 (L16308). doi:10.1029/ 2004GL02061. Ho, T.-Y., Quigg, A., Finkel, Z.V., Milligan, A.J., Wyman, K., Falkowski, P.G., Morel, F.M.M., 2003. The elemental composition of some marine phytoplankton. Journal of Phycology 39, 1145–1159. Honjo, S., Manganini, S.J., 1993. Annual biogenic particle fluxes to the interior of the North Atlantic Ocean; studied at 34°N 21°W and 48°N 21°W stations. Deep-Sea Research II 40, 587–607. Honjo, S., Francois, R., Manganini, S., Dymond, J., Collier, R., 2000. Particle fluxes to the interior of the Southern Ocean in the Western Pacific sector along 170°W. Deep-Sea Research II 47, 3521–3548. Hoppe, H.G., Ducklow, H., Karrasch, B., 1993. Evidence for the dependency of bacterial growth on enzymatic hydrolysis of particulate organic matter in the mesopelagic Ocean. Marine Ecology Progress Series 93, 277–283. Howard, M.T., Winguth, A.M.E., Klaas, C., Maier-Reimer, E., 2006. Sensitivity of ocean carbon tracer distributions to particulate organic flux parameterisations. Global Biogeochemical Cycles 20, GB3011. doi:10.1029/2005GB002499. Hwang, J., Druffel, E.R.M., Griffin, S., Smith, K.L.J., Baldwin, R.J., Bauer, J.E., 2004. Temporal variability of d14C, d13C, and C/N in sinking particulate organic matter at a deep time series station in the northeast Pacific Ocean. Global Biogeochemical Cycles 18, GB4015. doi:10.1029/2004GB00222. IPCC, 2001. In: Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., Linden, P.J.V.D., Dai, X., Maskell, K., Johnson, C.A. (Eds.), Intergovernmental Panel on Climate Change, Climate Change 2001. Cambridge University Press, New York. Ittekkot, V., 1993. The abiotically driven biological pump in the ocean and short-term fluctuations in atmospheric CO2 contents. Global Planetary Change 8, 17–25. Ittekkot, V., Haake, V., Bartsch, M., Nair, R.R., Ramaswamy, V., 1992. Organic carbon removal in the sea: the continental connection. In: Summerhayes, C.P., Prell, W.L., Emeis, K.C. (Eds.), Upwelling Systems: Evolution since the Early Miocene. Special Publication 64 Geological Society of America, pp. 167–176. Jackson, G., Burd, A.B., 2001. A model for the distribution of particle flux in the mid-water column controlled by subsurface biotic interactions. Deep Sea Research II 49, 193–217. Jahnke, R.A., 1996. The global ocean flux of particulate organic carbon: areal distribution and magnitude. Global Biogeochemical Cycles 10, 71–88. Jickells, T.D., Spokes, L.J., 2001. Atmospheric iron inputs to the oceans. In: Turner, D., Hunter, K.A. (Eds.), The Biogeochemistry of Iron in Seawater. Wiley, pp. 85–122. Joos, F., Plattner, G.-K., Stocker, T.F., Marchal, O., Schmittner, A., 1999. Global warming and marine carbon cycle feedbacks on future atmospheric CO2. Science 284, 464–467. Kahler, P., Koeve, W., 2001. Marine dissolved organic matter: can its C:N ratio explain carbon overconsumption? Deep-Sea Research I 48, 49–62. Karl, D.M., 2003. Microbiological oceanography – hidden in a sea of microbes. Nature 415, 590–591. Karl, D.M., Knauer, G.A., 1989. Swimmers: a recapitulation of the problem and a potential solution. Oceanography 2, 32–35. Karl, D.M., Dore, J.E., Christian, J.R., Letelier, R.M., Hebel, D.V., Tupas, L.M., Winn, C.D., 1996. Seasonal and interannual variability in primary production and particle flux at Station ALOHA. Deep-Sea Research II 43, 539–568. Karl, D.M., Bjo¨rkman, K.M., Dore, J.E., Fujieki, L., Hebel, D.V., Houlihan, T., Letelier, R.M., Tupas, L.M., 2001. Ecological nitrogento-phosphorus stoichiometry at Station ALOHA. Deep-Sea Research II 48, 1529–1566.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

309

Keeling, R.F., Shertz, S.R., 1992. Seasonal and interannual variations in atmospheric oxygen and implications for the global carbon cycle. Nature 358, 723–727. Keeling, R.F., Piper, S.C., Heimann, M., 1996. Global and hemispheric CO2 sinks deduced from changes in atmospheric O2 concentration. Nature 381, 218–221. Keil, R.G., Montluc¸on, D.B., Prahl, F.G., Hedges, J.I., 1994. Sorptive preservation of labile organic matter in marine sediments. Nature 370, 549–552. Kepkay, P.E., 1994. Particle aggregation and the biological reactivity of colloids. Marine Ecology Progress Series 109, 293–304. Klaas, C., Archer, D.E., 2002. Association of sinking organic matter with various types of mineral ballast in the deep sea: implications for the rain ratio. Global Biogeochemical Cycles 16. doi:10.1029/2001GB00176. Kling, S.A., Boltovskoy, D., 1995. Radiolarian vertical distribution patterns across the southern California Current. Deep-Sea Research 42, 191–231. Kobari, T., Shinada, A., Tsuda, A., 2003. Functional roles of interzonal migrating mesozooplankton in the western subarctic Pacific. Progress in Oceanography 57, 279–298. Koeve, W., 2002. Upper ocean carbon fluxes in the Atlantic Ocean: the importance of the POC:PIC ratio. Global Biogeochemical Cycles 16, #1056. Kortzinger, A., Koeve, W., Kahler, P., Mintrop, L., 2001. C:N ratios in the mixed layer during the productive season in the northeast Atlantic Ocean. Deep-Sea Research I 48, 661–688. Kristensen, E., Ahmed, S.I., Devol, A.H., 1995. Aerobic and anaerobic decomposition of organic matter in marine sediment: which is fastest? Limnology and Oceanography 40, 1430–1437. Kumar, N., Anderson, R.F., Mortlock, R.A., Froelich, P.N., Kubik, P., Dittrich-Hannen, B., Suter, M., 1995. Increased biological productivity and export production in the glacial Southern Ocean. Nature 378, 675–680. Lam, P.J., Bishop, J.K.B., in press. High biomass low export regimes in the Southern Ocean, Deep-Sea Research II. Lampitt, R.S., 1985. Evidence for the seasonal deposition of detritus for the deep-sea floor and its subsequent resuspension. Deep-Sea Research 32, 885–897. Lampitt, R.S., Antia, A.N., 1997. Particle flux in deep seas: regional characteristics and temporal variability. Deep-Sea Research I 44, 1377–1403. Lampitt, R.S., Wishner, K.F., Turley, C.M., Angel, M.V., 1993. Marine snow studies in the Northeast Atlantic Ocean: distribution, composition and role as a food source for migrating plankton. Marine Biology 116, 689–702. Laws, E.A., 2004. New production in the equatorial Pacific: a comparison of field data with estimates derived from empirical and theoretical models. Deep-Sea Research I 51, 205–211. Laws, E.A., Falkowski, P.G., Smith, W.O.J., Ducklow, H., McCarthy, J.J., 2000. Temperature effects on export production in the open ocean. Global Biogeochemical Cycles 14, 1231–1246. Lee, C., Wakeham, S.G., Hedges, J.I., 2000. Composition and flux of particulate amino acids and chloropigments in equatorial Pacific seawater and sediments. Deep-Sea Research I 47, 153–1568. Legendre, L., LeFevre, J., 1989. Hydrodynamical singularities as controls of recycled vs. export production in the oceans. In: Berger, W.H., Smetacek, V.S., Wefer, G. (Eds.), Dahlem Workshop Life Sciences Report – Productivity of the oceans: Present and Past. J. Wiley and Sons, pp. 155–173. Le Que´re´, C., Aumont, O., Monfray, P., 2003. Propagation of climatic events on ocean stratification, marine biology, and CO2: case studies over the 1979–1999 period. Journal of Geophysical Research 108 (C12). doi:10.1029/2001JC00092. Le Que´re´, C., Harrison, S.P., Prentice, I.C., Buitenhuis, E.T., Aumont, O., Bopp, L., Claustre, H., Cotrim da Cunha, L., Geider, R., Giraud, X., Klaas, C., Kohfeld, K.E., Legendre, L., Manizza, M., Platt, T., Rivkin, R.B., Sathyendranath, S., Uitz, J., Watson, A.J., Wolf-Gladrow, D., 2005. Ecosystem dynamics based on phytoplankton functional types for global ocean biogeochemistry models. Global Change Biology 11, 2016–2040. Li, Y.-H., Peng, T.-H., 2002. Latitudinal change of remineralization ratios in the oceans and its implication for nutrient cycles. Global Biogeochemical Cycles 16. doi:10.1029/2001GB00182. Li, Y.-H., Karl, D.M., Winn, C.D., Mackenzie, F.T., Gans, K., 2000. Remineralization ratios in the subtropical north Pacific gyre. Aquatic Geochemistry 6, 65–86. Logan, B., Passow, U., Alldredge, A., Grossart, H.-P., Simon, M., 1995. Rapid formation and sedimentation of large aggregates is predictable from coagulation rates (half-lives) of transparent exopolymer particles (TEP). Deep-Sea Research 42, 203–214. Loh, A.N., Bauer, J.E., 2000. Distribution, partitioning and fluxes of dissolved and particulate organic C, N and P in the eastern North Pacific and Southern Oceans. Deep-Sea Research I 47, 2287–2316. Longhurst, A.R., 1998. Ecological geography of the sea. Academic Press, San Diego. Longhurst, A.R., Harrison, W.G., 1988. vertical nitrogen flux from the oceanic photic zone by diel migrant zooplankton and nekton. DeeSea Research 35, 881–889. Louanchi, F., Najjar, R.G., 2000. A global monthly climatology of the upper ocean: Spring-summer export production and shallow remineralization. Global Biogeochemical Cycles 14, 957–977. Lourey, M.J., Trull, T.W., 2001. Seasonal nutrient depletion and carbon export in the Subantarctic and Polar Frontal Zones of the Southern Ocean south of Australia. Journal of Geophysical Research 106, 31463–31488. Lourey, M.J., Trull, T.W., Tilbrook, B., 2004. Sensitivity of d13C of Southern Ocean suspended and sinking organic matter to temperature, nutrient utilisation and atmospheric CO2. Deep-Sea Research I 51, 281–305. Lutz, M., Dunbar, R., Caldeira, K., 2002. Regional variability in the vertical flux of particulate organic carbon in the ocean interior. Global Biogeochemical Cycles 16, 11-11–11-18. doi:10.1029/2000GB001383.

310

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Maier-Reimer, E., 1993. Geochemical cycles in an ocean general circulation model: preindustrial tracer distributions. Global Biogeochemical Cycles 7, 645–677. Martin, J.H., Knauer, G.A., Karl, D.M., Broenkow, W.W., 1987. VERTEX: carbon cycling in the Northeast Pacific. Deep-Sea Research 34, 267–285. Matear, R.J., Holloway, G., 1995. Modeling the inorganic phosphorus cycle of the north pacific using an adjoint data assimilation model to assess the role of dissolved organic phosphorus. Global Biogeochemical Cycles 9, 101–119. Matear, R.J., Elliot, B., 2004. Enhancement of oceanic uptake of anthropogenic CO2 by macronutrient fertilization. Journal of Geophysical Research 109 (C04001). doi:10.1029/2000JC00032. Matear, R., Hirst, A.C., 1999. Climate change feedback on the future oceanic CO2 uptake. Tellus 51, 722–733. Mayer, L.M., 1994. Surface area control of organic carbon accumulation in continental shelf sediment. Geochimica Cosmochimica Acta 58, 1271–1284. Mayer, L.M., 1999. Extent of coverage of mineral surfaces by organic matter in marine sediments. Geochimica Cosmochimica Acta 63, 207–215. Michaels, A.F., Silver, M.W., 1988. Primary production, sinking fluxes and the microbial food web. Deep-Sea Research 35, 473–490. Michaels, A.F., Silver, M.W., Gowing, M.M., Knauer, G.A., 1990. Cryptic zooplankton ‘‘swimmers’’ in the upper ocean sediment traps. Deep-Sea Research 37, 1285–1296. Milligan, A.J., Varela, D.E., Brzezinski, M.A., Morel, F.M.M., 2003. Dynamics of silicon metabolism and silicon isotopic discrimination in a marine diatom as a function of pCO2 . Limnology and Oceanography 49, 322–329. Milliman, J.D., 1993. Production and accumulation of calcium carbonate in the ocean: budget of a non-steady (short-term) state. Global Biogeochemical Cycles 7, 927–957. Moore, J.K., Doney, S.C., Glover, D.M., Fung, I.Y., 2002a. Iron cycling and nutrient-limitation patterns in surface waters of the World Ocean. Deep-Sea Research II 49, 463–507. Moore, J.K., Doney, S.C., Kleypas, J.A., Glover, D.M., Fung, I.Y., 2002b. An intermediate complexity marine ecosystem model for the global domain. Deep-Sea Research II 49, 403–462. Napp, J.M., Hunt, G.L.J., 2001. Anomalous conditions in the South-Eastern Bering Sea 1997: linkages among climate, weather, ocean and biology. Fisheries Oceanography 10, 61–69. Nelson, D.M., Anderson, R.F., Barber, R.T., Brzezinski, M.A., Buesseler, K.O., Chase, Z., Collier, R.W., Dickson, M.-L., Franc¸ois, R., Hiscock, M.R., Honjo, S., Marra, J., Martin, W.R., Sambrotto, R.N., Sayles, F.L., Sigmon, D.E., 2002. Vertical budgets for organic carbon and biogenic silica in the Pacific sector of the Southern Ocean, 1996–1998. Deep Sea Research II 49, 1645–1674. Newton, P.P., Lampitt, R.S., Jickells, T.D., King, P., Boutle, C., 1994. Temporal and mesoscale variability of biogenic particle fluxes in the context of the JGOFS North-East Atlantic process studies at 47N 20W (1989–1990). Deep-Sea Research I 41, 1617–1642. Orr, J.C., 1999. On ocean carbon-cycle modal comparison. Tellus B 51, 509–510. Orr, J.C., 2002. EC Environment Climate Programme Global ocean storage of anthropogenic carbon (GOSAC) final report, Institute Pierre-Simon Laplace, Paris. Oschlies, A., Kahler, P., 2004. Biotic contribution to air–sea fluxes of CO2 and O2 and its relation to new production, export production, and net community production. Global Biogeochemical Cycles 18. doi:10.1029/2003GB00209. Pace, M.L., Knauer, G.A., Karl, D.M., Martin, J.H., 1987. Primary production, new production and vertical flux in the Eastern Pacific. Nature 325, 803–804. Pahlow, M., Riebesell, U., 2000. Temporal trends in deep ocean Redfield ratios. Science 287, 831–833. Palmer, J.R., Totterdell, I.J., 2001. Production and export in a global ocean ecosystem model. Deep-Sea Research I 48, 1169–1198. Panagiotopoulos, C., Sempere, R., Obernosterer, I., Striby, L., Goutx, M., Van Wambeke, F., Gautier, S., Lafont, R., 2002. Bacterial degradation of large particles in the southern Indian Ocean using in vitro incubation experiments. Organic Geochemistry 33, 985– 1000. Pantoja, S., Sepulveda, J., Gonzalez, H.E., 2004. Decomposition of sinking proteinaceous material during fall in the oxygen minimum zone off northern Chile. Deep-Sea Research I 51, 55–70. Passow, U., 2004. Switching perspectives: do mineral fluxes determine particulate organic carbon fluxes or vice versa? Geochemistry, Geophysics, Geosystems 5, Q04002. doi:10.1029/2003GC000670. Passow, U., De La Rocha, C.L., 2006. Accumulation of mineral ballast on organic aggregates. Global Biogeochemical Cycles 20, GB1013. doi:10.1029/2005GB00257. Peterson, M.L., Hernes, P.J., Thoreson, D.S., Hedges, J.I., Lee, C., Wakeham, S.G., 1993. Field evaluation of a valved sediment trap. Limnology and Oceanography 38, 1741–1761. Peterson, M.L., Wakeham, S.G., Lee, C., Askea, M.A., Miquel, J.C., 2005. Novel techniques for collection of sinking particles in the ocean and determining their settling rates. Limnology and Oceanography 3, 520–532. Peylin, P., Bousquet, P., Le Que´re´, C., Sitch, S., Friedlingstein, P., McKinley, G., Gruber, N., Rayner, P., Ciais, P., 2005. Multiple constraints on regional CO2 flux variations over land and oceans. Global Biogeochemical Cycles 19 (GB1011). doi:10.1029/ 2003GB00221. Pomeroy, L.R., Wiebe, W.J., Diebel, D., Thompson, R.J., Rowe, G.T., Pakulski, J.D., 1991. Bacterial response to temperature and substrate concentration during the Newfoundland spring bloom. Marine Ecology Progress Series 75, 143–159. Pondaven, P., Ruiz-Pino, D., Fravalo, C., Treguer, P., Jeandel, C., 2000. Interannual variability of Si and N cycles at the time-series station KERFIX between 1990 and 1995 – a 1-D modelling study. Deep-Sea Research I 47, 223–257. Rau, G.H., Takahashi, T., Des Marais, D.J., Sullivan, C.W., 1991. Particulate organic matter d13C variations across the Drake Passage. Journal of Geophysical Research 96, 15131–15135.

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

311

Redfield, A.C., Ketchum, B.H., Richards, F.H., 1963. The influence of organisms on the composition of seawater. In: Hill, M.N. (Ed.), The Sea, vol. 2. Inter-Science, New York, pp. 26–77. Reid, P.C., Edwards, M., Hunt, H.G., Warner, A.J., 1998. Phytoplankton change in the North Atlantic. Nature 391, 546. Richardson, A.J., Shoemann, D.S., 2004. Climate impact on Plankton ecosystems in the Northeast Atlantic Science. Science 305, 1609– 1612. Ridgewell, A., Zondervan,I., Hargreaves, J., Bijma, J., Lenton, T., in press. Significant long-term increase of fossil fuel CO2 uptake from reduced marine calcification. Geophysical Research Letters. Riebesell, U., Zondervan, I., Rost, B., Tortell, P.D., Zeebe, R.E., Morel, F.M.M., 2000. Reduced calcification of marine plankton in response to increased atmospheric CO2. Nature 407, 364–367. Rivkin, L., Legendre, L., 2001. Biogenic carbon cycling in the upper ocean: effects of microbial respiration. Science 291, 2398–2400. Rothlisberger, R., Bigler, M., Wolff, E.W., Joos, F., Monnin, E., Hutterli, M.A., 2004. Ice core evidence for the extent of past atmospheric CO2 change due to iron fertilisation. Geophysical Research Letters 31 (L16207). doi:10.1029/2004GL02033. Rubin, S.I., 2003. Carbon and nutrient cycling in the upper water column across the Polar Frontal Zone and Antarctic Circumpolar Current along 170°W. Global Biogeochemical Cycles 17, 1087. doi:10.1029/2002GB00190. Sabine, C.L., Feely, R.A., Gruber, N., Key, R.M., Lee, K., Bullister, J.L., Wanninkhof, R., Wong, C.S., Wallace, D., Tilbrook, B., Millero, F.J., Peng, T.H., Kozyr, A., Ono, T., Rios, A.F., 2004. The oceanic sink for anthropogenic CO2. Science 305, 367–371. Sambrotto, R.N., Mace, B.J., 2000. Coupling of biological and physical regimes across the Antarctic Polar Front as reflected by nitrogen production and recycling. Deep Sea Research II 47, 3339–3367. Sambrotto, R.N., Savidge, G., Robinson, C., Boyd, P., Takahashi, T., Karl, D.M., Langdon, C., Chipman, D., Marra, J., Codispoti, L., 1993. Elevated consumption of carbon relative to nitrogen in the surface ocean. Nature 363, 248–250. Sarmiento, J.L., Armstrong, R.A., 1997. The role of oceanic processes in the global carbon cycle. US JGOFS Synthesis and Modeling Project Implementation Plan. Woods Hole Oceanographic Institution, Woods Hole. Sarmiento, J.L., LeQuere, C., 1996. Oceanic carbon dioxide uptake in a model of century-scale global warming. Science 274, 1346–1350. Sarmiento, J.L., Hughes, T.M.C., Stouffer, R.J., Manabe, S., 1998. Simulated response of the ocean carbon cycle to anthropogenic climate warming. Nature 393, 245–249. Sarmiento, J.L., Gruber, N., Brzezinski, M.A., Dunne, J.P., 2004. High-latitude controls of thermocline nutrients and low latitude biological productivity. Nature 427, 56–60. Savoye, N., Dehairs, F., Elskens, M., Cardinal, D., Kopczynska, E.E., Trull, T.W., Wright, S., Baeyens, W., Griffiths, F.B., 2004. Regional variation of spring N-uptake and new production in the Southern Ocean. Geophysical Research Letters 31, L03301. doi:10.1029/2003GL018946. Sayles, F.L., Martin, W.R., Chase, Z., Anderson, R.F., 2001. Benthic remineralization and burial of biogenic SiO2 CaCO3, organic carbon and detrital material in the Southern Ocean along a transect at 170°W. Deep-Sea Research II 48, 4323–4383. Scharek, R., Tupas, L.M., Karl, D.M., 1999. Diatom fluxes to the deep sea in the oligotrophic North Pacific gyre at Station ALOHA. Marine Ecology Progress Series 182, 55–67. Schiebel, R., 2002. Planktic foraminiferal sedimentation and the marine calcite budget. Global Biogeochemical Cycles 16, 1065. doi:10.1029/2001GB00145. Schlitzer, R., 2002. Carbon export fluxes in the Southern Ocean: results from inverse modeling and comparison with satellite-based estimates. Deep-Sea Research II 49, 1623–1644. Schlitzer, R., 2004. Export production in the Equatorial and North Pacific derived from dissolved oxygen, nutrient, and carbon data. Journal of Oceanography 60, 53–62. Schneider, B., Schlitzer, R., Fischer, G., Nothig, E.-M., 2003. Depth dependent elemental compositions of particulate organic matter (POM) in the ocean. Global Biogeochemical Cycles 17 (2). doi:10.1029/2002GB00187. Scholten, J.C., Fietzke, F., Vogler, S., Rutgers van der Loeff, M., Mangini, A., Koeve, W., Waniek, J., Stoffers, P., Antia, A., Kuss, J., 2001. Trapping efficiency of sediment traps from the deep eastern North Atlantic: 230Th calibration. Deep-Sea Research II 48, 243–268. Schulz, K.G., Zondervan, I., Gerringa, L.J.A., Timmermans, K.R., Veldhuis, M.J.W., Riebesell, U., 2004. Effect of trace metal availability on coccolithophorid calcification. Nature 430, 673–676. Shaffer, G., 1996. Biogeochemical cycling in the global ocean, 2, New production, Redfield ratios, and remineralization in the organic pump. Journal of Geophysical Research 101, 3723–3745. Shaffer, G., Bendtsen, J., Ulloa, O., 1999. Fractionation during remineralization of organic matter in the ocean. Deep-Sea Research I 46, 185–204. Sheridan, C.C., Lee, C., Wakeham, S.G., Bishop, J.K.B., 2002. Suspended particle organic composition and cycling in surface and midwaters of the equatorial Pacific Ocean. Deep-Sea Research I 49, 1983–2008. Sigman, D.M., Boyle, E.A., 2000. Glacial/Interglacial variations in atmospheric carbon dioxide. Nature 407, 859–869. Silver, M.W., Gowing, M.M., 1991. The ‘‘particle’’ flux: origins and biological components. Progress in Oceanography 26, 75–113. Six, K.D., Maier-Reimer, E., 1996. Effects of plankton dynamics on seasonal carbon fluxes in an ocean general circulation model. Global Biogeochemical Cycles 10, 559–583. Small, L.F., Fowler, S.W., Moore, S.A., La Rosa, J., 1983. Dissolved and fecal pellet carbon and nitrogen release by zooplankton in tropical waters. Deep-Sea Research 30, 1199–1220. Smith, D.C., Simon, M., Alldredge, A.L., Azam, F., 1992. Intense hydrolytic enzyme activity on marine aggregates and implications for rapid particle dissolution. Nature 359, 139–142. Smith, D.C., Steward, G.F., Long, R.A., Azam, F., 1995. Bacterial mediation of carbon fluxes during a diatom bloom in a mesocosm. Deep-Sea Research II 42, 75–98.

312

P.W. Boyd, T.W. Trull / Progress in Oceanography 72 (2007) 276–312

Stanley, R.H.R., Buesseler, K.O., Manganini, S.J., Steinberg, D.K., Valdes, J.R., 2004. A comparison of major and minor elemental fluxes collected in neutrally buoyant and surface-tethered sediment traps. Deep-Sea Research I 51, 1387–1395. Steinberg, D.K., Pilskaln, C.H., Silver, M.W., 1998. Contribution of zooplankton associated with detritus to sediment trap ‘swimmer’ carbon in Monterey Bay, California, USA. Marine Ecology Progress Series 164, 157–166. Steinberg, D.K., Carlson, C.A., Bates, N.R., Goldthwait, S.A., Madin, L.P., Michaels, A.F., 2000. Zooplankton vertical migration and the active transport of dissolved organic and inorganic carbon in the Sargasso Sea. Deep-Sea Research I 47, 137–158. Stramski, D., 1999. Estimation of particulate organic carbon in the ocean from satellite remote sensing. Science 285, 239–241. Suess, E., 1980. Particulate organic carbon flux in the oceans-surface productivity and oxygen utilization. Nature 288, 260–263. Takahashi, K., 1991. Radiolaria: flux, ecology and taxonomy in the Pacific and Atlantic. In: Honjo, S. (Ed.), Ocean Biocoenosis Series. Woods Hole Oceanographic Institution, Woods Hole, p. 303. Takahashi, K., Billings, J.D., Morgan, J.K., 1990. Oceanic province: assessment from the time-series diatom fluxes in the North Eastern Pacific. Limnology and Oceanography 35, 154–165. Thomas, H., Ittekkot, V., Osterroht, C., Schneider, B., 1999. Preferential recycling of nutrients – the ocean’s way to increase new production and to pass nutrient limitation. Limnology and Oceanography 44, 1999–2004. Toggweiler, J.R., 1993. Carbon overconsumption. Nature 363, 248–250. Tre´guer, P., Jacques, G., 1992. Dynamics of nutrients and phytoplankton, and fluxes of carbon, nitrogen, and silicon in the Antarctic ocean. Polar Biology 12, 149–162. Tre´guer, P., Nelson, D.M., van Bennekom, A.J., DeMaster, D.J., Leynaert, A., Que´giner, B., 1995. The silica balance in the world ocean: a re-estimate. Science 268, 375–379. Tre´guer, P., Legendre, L., Rivkin, R.T., Ragueneau, O., Dittert, N., 2003. Water column biogeochemistry below the euphotic zone. In: Fasham, M.J.R. (Ed.), Ocean Biogeochemistry – the Role of the Ocean Carbon Cycle in Global Change. Springer-Verlag, Berlin, pp. 145–156. Trull, T.W., Bray, S.G., Manganini, S.J., Honjo, S., Franc¸ois, R., 2001a. Moored sediment trap measurements of carbon export in the Subantarctic and Polar Frontal Zones of the Southern Ocean, south of Australia. Journal of Geophysical Research 106, 31489–31510. Trull, T.W., Rintoul, S.R., Hadfield, M., Abraham, E.R., 2001b. Circulation and seasonal evolution of Polar waters south of Australia: implications for iron fertilisation of the Southern Ocean. Deep Sea Research II 48, 2439–2466. Turley, C.M., 1993. The effect of pressure on leucine and thymidine incorporation by free-living bacteria attached to sinking oceanic particles. Deep-Sea Research I 40, 2193–2206. Turley, C.M., Mackie, P.J., 1995. Bacterial and cyanobacterial flux to the deep NE Atlantic on sedimenting particles. Deep-Sea Research I 42, 1453–1474. Tyrrell, T., 1999. The relative influences of nitrogen and phosphorus on oceanic primary production. Nature 400, 525–531. Usbeck, R., van der Loeff, M.R., Hoppema, M., Schlitzer, R., 2002. Shallow mineralization in the Weddell Gyre. Geochemistry, Geophysics, Geosystems 3. doi:10.1029/2001GC000182. Usbeck, R., Schlitzer, R., Fischer, G., Wefer, G., 2003. Particle fluxes in the ocean: comparison of sediment trap data with results from inverse modelling. Journal of Marine Systems 39, 167–183. US-JGOFS, 2001. http://usjgofs.whoi.edu/mzweb/caco3_rpt.html/. Valdes, J.R., Price, J.F., 2000. A neutrally buoyant, upper ocean sediment trap. Atmospheric and Oceanographic Technology 17, 62–68. Van Mooy, B.A.S., Keil, R.G., Devol, A.H., 2002. Impact of suboxia on sinking particulate organic carbon: enhanced carbon flux and preferential degradation of amino acids via denitrification. Geochimica et Cosmochimica Acta 66, 457–465. Verity, P., Smetacek, V.S., 1996. Organism life cycles, predation, and the structure of marine pelagic ecosystems. Marine Ecology Progress Series 130, 277–293. Verity, P.G., Bauer, J.E., Flagg, C.N., DeMaster, D.J., Repeta, D.J., 2002. The Ocean Margins Program: an interdisciplinary study of carbon sources, transformations, and sinks in a temperate continental margin system. Deep Sea Research II 49, 4273–4295. Volk, T., Hoffert, M.I., 1985. Ocean carbon pumps: analysis of relative strengths and efficiencies in ocean-driven atmospheric CO2 changes. In: Sundquist, E., Broecker, W.S. (Eds.), The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present. AGU, Washington, DC, pp. 99–110. Walsh, J.J., 1991. Importance of continental margins in the marine biogeochemical cycling of carbon and nitrogen. Nature 350, 53–55. Wang, X., Matear, R.J., Trull, T.W., 2003. Nutrient utilization ratios in the Polar Frontal Zone in the Australian sector of the Southern Ocean: a model. Global Biogeochemical Cycles 17, 1009. doi:10.1029/2002GB00193. Wefer, G., 1989. Particle flux in the ocean; effects of episodic production. In: Berger, W.H., Smetacek, V.S., Wefer, G. (Eds.), Productivity of the Ocean: Present and Past. Dahlem Life Science Research Report 44. J. Wiley and Sons, New York, pp. 139–154. Wong, C.S., Matear, R.J., 1999. Sporadic silicate limitation of phytoplankton productivity in the subarctic NE Pacific. Deep-Sea Research II 46, 2539–2556. Wong, C.S., Whitney, F.A., Crawford, D.W., Iseki, K., Matear, R.J., Johnson, W.K., Page, J.S., Timothy, D., 1999. Seasonal and interannual variability in particle fluxes of carbon, nitrogen and silicon from time series of sediment traps at Ocean Station P, 1982– 1993: relationship to changes in subarctic primary productivity. Deep-Sea Research II 46, 2735–2760. Wright, T.D., Vergin, K.L., Boyd, P.W., Giovannoni, S.J., 1997. A novel d-subdivision proteobacterial lineage from the lower ocean surface layer. Applied and Environmental Microbiology 63, 1441–1448. Yu, E.-F., Francois, R., Bacon, M.P., Honjo, S., Fleer, A.P., Manganini, S.J., Rutgers van der Loeff, M.M., Ittekot, V., 2001. Trapping efficiency of bottom-tethered sediment traps estimated from the intercepted fluxes of 230Th and 231Pa. Deep-Sea Research 48, 865–889. Zhang, J.Z., Mordy, C.W., Gordon, L.I., Ross, A., Garcia, H.E., 2000. Temporal trends in deep ocean Redfield ratios. Science 289, 1839– 1841.