phosphorus regeneration in continental margin sediments

Pergamon

Geochimica et Cosmochimica Acta, Vol. 61, No. 14, pp. 2891-2907, 1997 Copyright 0 1997 Elsevier Science Ltd Printed in the USA. All rights reserved

0016-7037/97 $17.00 + .OO

PI1 SOO16-7037( 97)00138-5

Phosphorus regeneration in continental margin sediments

JAMES MCMANUS, I** WILLIAM M. BEFLELSON, ’ KENNETH H. COALE,’ KENNETH S. JOHNSON,*-3 and TAMMY E. KILGORE’ ‘Department of Earth Sciences, University of Southern California, Los Angeles, California 90089-0740, USA

‘Moss Landing Marine Labs, P.O. Box 450, Moss Landing, California 95039-0450, USA ‘Monterey Bay Aquarium Research Institute, P.O. Box 628, Moss Landing, California 9.5039-0628, USA

(Received June 27, 1996; accepted in revised form March 20, 1997)

Abstract-Benthic incubation chambers have been deployed in a variety of geochemical environments along the California Continental Margin. These include both high and low oxygen environments and sites where the rate of organic matter oxidation on the seafloor (C,,) ranges from < 1 mmol mm2 day-’ to more than 7 mmol m-* day -I through a depth range of 100-3500 m. This range in the rate of organic matter oxidation along with variations in the concentration of bottom water oxygen allow us to elucidate the diagenetic conditions under which P regeneration may be decoupled from organic matter cycling. Under conditions where bottom water oxygen concentration is low (<50 PM), and the rate of organic matter oxidation is also low ( < 1 mmol mm2 day -’ ), P regeneration may be less than that expected from the decay of organic debris and, in some cases, there is a flux of phosphate into the sediments. At stations where bottom water oxygen is low, and the degradation rate of organic material is greater than 1 mm01 m-* day -‘, phosphate may be released at a rate exceeding the production expected from the oxidation of organic matter. At stations having high bottom water oxygen concentrations, rates of organic matter decomposition < -7 mmol me2 day-‘, and where benthic irrigation is not significant, P regeneration is consistent with that expected from the decomposition of organic debris. In addition, our data indicate that high benthic iron fluxes are observed in regions exhibiting a decoupling between organic matter and phosphate, whereas low to zero iron fluxes are observed in regions where P regeneration is either consistent with or less than that expected from the decomposition of organic material. These results support previous work suggesting a coupling between iron cycling and phosphate cycling in suboxic environments. Data presented here show that this coupling may result in either preferential phosphate burial or release relative to organic material in suboxic environments. Copyright 0 1997 Elsevier Science Ltd

1. INTRODUCTION

It has been argued that phosphorus provides the long-term limit on the magnitude of oceanic primary production (e.g., Holland, 1978; Broecker and Peng, 1982). Thus, to success- fully model changes in ocean fertility and chemistry requires an understanding of the geochemical cycling of this major nutrient. Despite the substantial body of literature on the diagenesis of phosphorus in marine systems, the controls on phosphorus burial in modem sediments are still not well understood (cf. Ruttenberg, 1993; Filippelli, 1994). Part of this lack of knowledge derives from the fact that phosphorus cycling in marine sediments is a complex process. The phosphorus cycle is influenced by the cycling of biogenic debris, Fe, and possibly Mn; as well as bottom water oxygen concentration and authigenic mineral precipitation (e.g., Froe- lich et al., 1982; Mach et al., 1987; Lucotte and d’Anglejan, 1988; Van Cappellen and Bemer, 1988; Ruttenberg and Bemer, 1993; Ingall et al., 1993; Ruttenberg, 1993; Filip- pelli, 1994; Wheat et al., 1996; and others).

Most of the phosphorus initially transported to modern sediments is incorporated into organic material (e.g., Froe- lich et al., 1982). However, although organic matter is the dominant shuttle for phosphorus transfer from the upper water column to the sediment, organically-bound phosphorus

*Present address: Oregon State University, COAS, Ocean Admin Bldg. 104, Corvallis, Oregon 97331-5503, USA.

2891

may represent less than 25% of the total reactive-P that is ultimately buried in marine sediments (Ruttenberg, 1993). The other dominate burial phases for reactive P are iron-P and authigenic-P (Ruttenberg, 1993). The disparity between the importance of organic phosphorus as a source to the sediments and inorganic phosphorus phases as a sink, reflects in part the diagenetic transfer of P from an organic phase to an inorganic phase (e.g., Bemer et al., 1993; Ruttenberg, 1993; Ruttenberg and Bemer, 1993; Filippelli, 1994).

Continental margin sediments play a major role in the geochemical cycle of this element, and it is precisely these sediments that are the most complex with respect to phosphorus geochemistry (e.g., Bemer et al., 1993; Ruttenberg, 1993; Ruttenberg and Bemer, 1993). Furthermore, given the complexity of phosphorus cycling in these environments, it has become clear that there is a paucity of continental margin data from which to accurately quantify the global phosphorus cycle (cf. Ruttenberg, 1993). For example, there is mounting evidence that the phosphorus burial efficiency in low oxygen continental margin environments is reduced relative to environments having high bottom water oxygen concentrations (e.g., Ingall and Jahnke, 1994, 1996). On the other hand, both high and low oxygen margin environments have been found to host contemporary accumulations of authigenic apatite, suggesting perhaps an enhanced burial efficiency of phosphorus in some locations, e.g., the California margin (Reimers et al., 1996), the Mexican margin (Jahnke et al., 1983; Schuffert et al., 1994), the Peru margin (Veeh and

2892 J. McManus et al.

Burnett, 1973; Burnett et al., 1982; Froelich et al., 1988; Glenn and Arthur, 1988), and the West African margin (Price and Calve& 1978; Birch, 1979). These examples, along with the facts that the mechanisms regulating authi- genie apatite precipitation and the quantitative significance of these deposits in the modern ocean are poorly constrained, illustrate the complexity of understanding the phosphorus system in the marine environment.

This paper focuses on the rate of phosphorus regeneration along the central and southern California continental margin. Our intention is to examine the influence of bottom water oxygen concentration, the organic carbon oxidation rate (C,,), and the cycling of Fe on the regeneration rate of dissolved phosphate. Based on data presented here and within the literature, we suggest that there are conditions currently active along the western North American margin where P is being sequestered in sediments relative to organic

carbon. The exact mechanism for this process remains enigmatic; however, we suggest that iron cycling and the depth of oxygen penetration (a combination of bottom water oxygen

concentration and C,,) may play a significant, although not exclusive, role in phosphorus burial and regeneration in this region.

2. STUDY AREA AND METHODS

This paper presents results from four cruises off the California continental margin (Fig. 1; Table 1). Two cruises, CC 1 & 2, con- sisted of transects that were perpendicular to the central California continental margin, another cruise took place in the California Bor- derland Basin area, and the final cruise, CC 3, took place in both regions. The central California sites were in a highly productive region underlying the California current system where the oxygen minimum zone impinges on the continental margin at depths ranging from approximately 500 to 1000 m (Fig. 2a). The rate of seafloor organic carbon oxidation in this region ranges from -0.5 to >7 mm01 me2 day-’ (Fig. 2b; Berelson et al., 1996). The southern California Borderlands area consists of a series of submarine basins separated from each other by a series of topographic ridges and sills. Below the sill depth of each basin, horizontal flow is restricted, and the depth of each sill dictates the hydrographic properties of each basin (e.g., Berelson et al., 1987). The combination of restricted flow and varying sill depths in these basins results in several basins becoming nearly anoxic, thus providing an opportunity to study systems having lower concentrations of bottom water oxygen than are present at comparable depths within the oxygen minimum zone (Fig. 2a).

The benthic chambers (landers) used in this study and methods for calculation of fluxes have been described extensively elsewhere (Berelson and Hammond, 1986; Berelson et al., 1987, 1994; Ham- mond et al., 1996). Briefly, each chamber has a volume of approximately 7 L, covers a sediment surface area of 720 cm’, and is stirred with a paddle. A trace metal clean CsCl spike solution is added in situ to determine chamber volume, and Cs is later measured on individual samples by flame emission techniques. Six chamber samples are drawn during incubations; each draw removes approximately 230 mL of water from the chamber, and this water is replaced by ambient bottom water. Chamber modifications were made to mini- mize trace metal contamination including all PVC and acrylic con- struction. It should be noted that chamber water oxygen concentration was monitored with a pulsed electrode every 6 min. Any chamber data that was collected after the oxygen vs. time relationship became nonlinear was not included in flux calculations (Berelson et al., 1996). There are two reasons that chamber data may become nonlinear. ( 1) In the case of phosphate, if chamber oxygen concentrations go to zero, a change in chamber redox chemistry will release Fe-bound phosphate from the sediments yielding high P-fluxes (e.g., Sundby et al., 1986). For all stations presented, oxygen concentra-

tions did not go to zero during the course of the incubations, nor do we observe significant curvature suggesting increasing P release with time. (2) Because solute concentrations within an incubation chamber can change during sediment incubation, the porewater- overlying water concentration gradient will theoretically vary during the course of the experiment, thus yielding a nonlinear concentration vs. time relationship (e.g., Bender et al., 1989). Given the relatively small increases in phosphate concentration relative to the analytical uncertainties that are observed for most incubations and the lack of apparent curvature for all but those data from Santa Monica basin (Fig. 3), all data are fit with a linear model as previously described (e.g., Hammond et al., 1996). For the data from Santa Monica basin, we only fit those data which could be approximated by a linear concentration vs. time relationship.

Cores processed for porewaters were collected using a multiple corer (Barnett et al., 1984). This device simultaneously lowered eight individual 10 cm diameter polycarbonate core tubes into the sediment. Porewaters were extracted at approximately bottom water temperatures on board ship. Cores were sectioned under a nitrogen atmosphere, typically at intervals of 0.5 cm (0- 1 cm), 1 cm ( l-5 cm), 2 cm (5-13 cm), 3 cm (13-16 cm), and 4 cm (>I6 cm). The sectioned mud was centrifuged (7- 10,000 rpm), and the water obtained was filtered through a 0.45 Frn filter under a nitrogen atmosphere and analyzed for the major nutrients.

Samples for nutrient analyses were filtered (0.45 pm) and refriger- ated at -2°C until analysis. Porewater and lander nutrient analyses were made by flow injection analysis using a Lachat QuikChem 4200 auto-analyzer. The nutrients were measured in replicate on filtered subsamples with analytical uncertainties of l%, 2.50/c, 2%, and 0.2 PM for silicic acid, nitrate + nitrite, phosphate, and ammonia, respectively. Total dissolved iron was determined by a chemilu- minescent technique (Obata et al., 1993) modified for flow injection analysis (Elrod et al., 1991). In most cases bottom water nutrient and trace element concentrations were obtained from analysis of water collected either from a Niskin sampler that was attached to the lander or from adjacent hydrocasts. In some instances we did not retrieve bottom water samples due to failure of niskin samplers, these stations are identified as those not having bottom water values in Fig. 3.

The rate of organic carbon oxidation (C,,) in the sediments studied here (Table 2) was determined from X0, fluxes into the incubation chambers and estimates of the portion of this flux that is attributable to calcium carbonate dissolution as discussed in Berelson et al. (1996). The depth of oxygen penetration was taken from Berelson et al. ( 1996).

3. RESULTS

The rate of organic carbon oxidation as a function of bottom depth decreases in a nonlinear fashion with increasing bottom depth. This observation has been reported on numerous occasions and reflects the pattern of organic carbon delivery to the seafloor (e.g., Martin et al., 1987; Devol and Christensen, 1993; Berelson et al., 1996). The rate of organic carbon oxidation does not differ significantly between the basin and the margin sites; however, there are some temporal differences among revisited sites (Table 1 and discussed in Berelson et al., 1996).

Benthic phosphate regeneration rates measured at our study sites demonstrate that phosphate may be either released or taken up at the sediment-water interface (Table 1; Figs. 3 and 4a). Phosphate uptake at the sediment-water interface is limited to those sites located at regions having low ( <50 PM) bottom water oxygen concentrations (Fig. 4a). How- ever, the oxygen minimum zone also hosts the highest phosphate fluxes observed during this study (Fig. 4a). Thus, low bottom water oxygen concentration is not the sole criteria for sedimentary phosphate uptake in this region.

Geochemistry of P in coastal marine sediment 2893

38'

36"

32"

30" I I

1180

Fig. 1. Map of study area (CC stations are marked l-n, 2-n, and 3-n respectively). The borderland basins are identified in the south where SM is Santa Monica Basin, SP San Pedro, Cat Catalina Basin, Cl San Clemente Basin, TB Tanner Basin, and PE Patton Escarpment. V and PV represent locations of sediment trap deployments VERTEX 1 and PREVTX (Martin et al., 1987)) MB denotes location of a sediment trap deployed in Monterey Canyon (Pilskaln et al., 1996), and the site identified as M is the location of traps described by Smith et al. (1994). Depth contours are in meters.

Iron fluxes are consistently greater than zero at bottom water depths shallower than -1500 m. This observation likely reflects at least some coupling of the iron cycle with the biogeochemical cycling of organic matter, i.e., higher organic carbon oxidation rates are also within this depth region (Berelson et al., 1996; Fig. 2b). The exception to this broad pattern is within the oxygen minimum zone. As with the phosphate fluxes, the highest benthic iron fluxes are within the oxygen minimum zone (Fig. 4b). Throughout most of the study area, those locations having the lowest

dissolved oxygen concentration also tend to have the highest benthic iron flux (Table I), thus suggesting an additional, but expected, relationship between bottom water oxygen concentration and benthic iron remobilization.

From the flux data in Table 1, it is apparent that many factors will affect the flux of both Fe and phosphate. These factors may operate within a single basin, for example Lander TB4 shows low phosphate (not distinguishable from zero) and low iron fluxes, whereas Lander TB7 shows high phosphate and iron fluxes (Table 1). These devices were


Table 1. Lander Flux Measurements

Station Lat. Long. LandKham (“N) (“w Date

Bottom depth [O&w (m) (H)

CC 1 (6/91)

CCl-5-R -Y

cc 1-6-R -Y

CC1-I-R -Y

CCI-12-R -B

Ccl-llA-R Ccl-llB-R

-Y

CC 2 (6/92)

CCZ-I-B -R -Y

CCZ-2-B -R -Y

CCZ-3-R -Y

ccz-4-B -R -Y

CCZ-5-B -R -Y

ccz-6-Y

Borderland basins (3/94)

PE-4-B -R -Y

PE-7-B -R -Y

TB-2-B -Y

TB-7-B -R

Cat-4B-R -Y

SP-7-B -R -Y

SM.7-R

X-7-B -R -Y

cc 3 (11/95)

CC3-2-B -R -Y

CC3-3-B -R -Y

CC3-4-B -R -Y

CC3-6-B -R -Y

CC&B -R -Y

CC3-9-B -R -Y

CC3-10-B -R -Y

CC3-12-B -R -Y

35.7 121.5 61719 1 532 48

35.4 121.1 61919 1 231 17

35.2 121.3 6/11/91 638 16

35.5 121.6 6/18/91 1010 22

35.6 121.2 6/21/91 95 100 35.6 121.2 612319 1 95 100

35.6 121.3 519192 97 138

35.2 121.3 5110192 670 18

35.5 121.6 5/l l/92 1010 24

36.1 122.4 5114192 2025 80

36.1 122.3 5/15/92 3375 113

36.2 122.4 5119192 1358 47

32.4 120.6 3/9- 13194 3707 132

32.4 120.6 319-13194 3707 132

33.0 119.7 316-8194 1514 27

33.0 119.7 316-8194 1514 27

33.3 118.6 3/22-23194 1300 19

33.5 118.4 3116-18194 896 8

33.7 118.8 3/19-21194 905 IO

32.6 118.1 3/22-24194 2053 52

36.2 122.4 1 l/2-3/95 1455 53

36.0 123.0 1 l/2-4/95 3595 133

36.1 122.5 1 l/4-5/95 2215 95

35.2 121.3 1 l/7-8/95 670 18

33.7 118.8 11/12-13/95 905 9

32.6 118.1 11/13-14195 2070 65

32.4 120.6 11/14-16195 3710

1515

137

33.0 119.8 11/17-18195 26

- Phosphate Iron

(mm01 II-* day-‘) (pmol mm2 day-‘)

0.070 + 0.020 0.065 f 0.016 0.040 + 0.009 0.127 f 0.011 0.006 5 0.009 0.031 + 0.008

-0.017 + 0.029 -0.023 + 0.029

0.052 2 0.009 0.042 2 0.016 0.055 2 0.019

0.024 t 1.04 2.81 k 15.75

(48 -t 11) (38 -t 8)

12.01 i 10.57 10.59 z 1.31

0.15 + 0.76

0.72 + 0.28 3.64 f 1.41 1.30 f 0.54

0.107 It 0.011 0.29 + 0.08 0.091 -+ 0.027 1.32 -t 0.09 0.068 2 0.009 0.42 k 0.17

-0.022 2 0.004 0.90 2 0.53 0.004 + 0.006 2.76 i- 0.26

-0.016 2 0.004 2.62 ? 0.76 0.006 2 0.006 1.46 i- 0.68

-0.002 2 0.003 1.72 +- 0.69 0.021 2 0.002 1.61 i 0.36 0.010 I! 0.002 0.22 f 0.22 0.010 t 0.002 1.02 2 0.23 0.009 t 0.002 0.50 -t 0.10 0.006 i 0.002 0.33 + 0.21 0.009 2 0.002 0.30 + 0.10 0.014 t 0.003 2.65 + 0.24

0.002 2 0.002 0.002 2 0.001 0.004 ? 0.002 0.008 2 0.004 0.003 z 0.003 0.006 2 0.004 0.002 + 0.002 0.005 + 0.004 0.022 + 0.007 0.011 + 0.003 0.005 2 0.003

-0.004 I! 0.004 0.020 2 0.008 0.027 i- 0.012

0.104 -t 0.023 0.093 t 0.021 0.006 ? 0.003 0.008 i- 0.002 0.007 t 0.003

0.18 -c 0.06 -0.26 +- 0.10 -0.19 2 0.07 -0.48 2 0.61 -0.46 ? 0.15

0.13 t 0.27 0.38 2 0.1 I

-0.22 i 0.13 6.87 2 1.60 1.14 2 0.45 1.20 -t 1.40 1.10 -i- 0.51 19.8 z 9.2 19.3 i 3.3 4.10 2 1.20 7.35 2 1.40 17.0 t 3.0 0.33 t 0.36 0.18 2 0.05 0.37 t 0.33

0.019 ? 0.005 0.015 + 0.004 0.030 2 0.004 0.008 + 0.003 0.013 + 0.003 0.010 + 0.003 0.021 + 0.006 0.020 + 0.003 0.024 + 0.004 0.011 + 0.006 0.019 + 0.004 0.015 c 0.004 0.159 -c 0.032 0.192 2 0.056 0.166 + 0.034 0.010 lr 0.003 0.008 -t 0.002 0.009 + 0.002 0.004 2 0.002 0.006 k 0.002 0.004 2 0.002 0.024 + 0.004 0.007 t 0.003 0.035 2 0.005

2.53 -c 1.26 -0.21 2 0.14

3.68 -c 0.96 -0.15 t 0.13

0.04 2 0.10 -0.20 t 0.26 -0.25 2 0.13

0.99 ? 0.28 -0.27 ? 0.66 13.15 t 4.41

3.29 t 2.03 16.47 i- 4.74 19.36 t 14.4 21.81 2 4.62 12.74 i- 5.04

0.02 2 0.06 0.15 t 0.04 0.08 2 0.02 0.62 t 0.16

PO.37 ? 0.09 -0.06 i- 0.07

0.70 5 0.32 1.11 i- 0.52 0.09 -t 0.33

Uncertainties are +lc where positive values are out of sediments and negative values are into sediments. Note, stations CC3-8, 9, IO, 12 are the Santa Monica Basin, San Clemente Basin, Patton Escarpment, and Tanner Basin, respectively.


Dissolved oxygen (PM) 50 100 150 200 250

. Basin Sites -Patton Escarpment

Co , (mmol me ’ day’ ‘)

1 2 3 4 5 6 7 8

1 1500 1 .%O 0 1 ?? Basin sites

0 Central Cslif. sites

3500 rn8

4000 ““‘.“‘I..-,“““‘I”“I”“I””

Fig. 2. (a) Dissolved oxygen as a function of depth. Circles and squares represent bottom water values from the study sites and the line is a water column profile taken at the Patton Escarpment. (b) Benthic organic carbon oxidation rate (C,,) as a function of depth from the study sites (from Berelson et al., 1996).

deployed within 2 km of each other. Likewise both the magnitude and direction of the phosphate flux for Catalina Basin are uncertain, but there does appear to be a significant flux of Fe from the sediments in that basin. Despite the variability between chambers of a lander and between landers at a given station, the variability among stations is distinctive.

4. DISCUSSION

It has been suggested that bottom water oxygen concentration plays a role in determining the organic carbon to organic phosphorus ( Corg :Porg) ratio in marine sediments (e.g., Ingall and Van Cappellen, 1990; Ingall et al., 1993). For example, low sedimentary Corg :P,, ratios have been ascribed to the ability of certain bacteria to preferentially sequester P relative to C in high oxygen environments and to the presence of some highly refractory organo-phosphorus compounds (Ingall et al., 1990; Ingall and Van Cappellen, 1990). High C,,:P,, ratios found in sediments that were likely covered by anoxic waters have been ascribed to the inhibition of bacterial sequestering of P, and in these environments it is proposed that C is preserved preferentially over P (e.g., In- gall and Jahnke, 1994; 1996; and references therein).

In addition to variations in organic phosphorus preservation, other diagenetic factors will also influence reactive phosphorus burial efficiency (e.g., Van Cappellen and Ingall, 1994; Ingall and Jahnke, 1996). The interactions of phosphate with iron oxide phases can alter the burial efficiency of total sedimentary phosphorus (e.g., Ruttenberg, 1993; Van Cappellen and Ingall, 1996; Wheat et al., 1996; Slomp et al., 1996). Because P adsorbs onto Fe and possibly manganese- oxide phases (e.g., Krom and Berner, 1981; Klump and

Martens, 1981; Froelich et al., 1982; Lucotte and d’Anglejan, 1983, 1988; Martens, 1993; Ingall and Jahnke, 1994; and others), this phase could trap or release phosphate in a near- surface layer of marine sediments. Iron chemistry could thus produce two possible relationships between bottom water oxygen concentration and the flux of P from marine sediments. For example, if iron oxidation is occurring sufficiently close to the sediment-water interface, then P may be sequestered onto near-surface oxides as it is released from organic matter during early diagenesis. Alternatively, if this boundary is above the sediment-water interface, then phosphate released from organic matter will diffuse into the overlying water column. Further, there is the potential that any phosphate that has entered the sediments associated with iron oxide phases will be released in addition to the phosphate from organic matter oxidation. The magnitude of this potential source of regenerated phosphate may thus also de- pend on bottom water oxygen concentration or oxygen penetration depth.

Data presented here as well as elsewhere (Ingall and Jahnke, 1996) suggest that although, in general, the C,,:P regeneration ratio along the continental margin is similar to that expected from the regeneration of Redfield organic matter, there are places where phosphate is clearly being taken up at the sediment surface and where phosphate is being released in excess of that expected from the degradation of organic matter (Fig. 5). We hypothesize that the deviations from Redfield C,,:P regeneration ratios observed along the California margin derive, in part, from interactions between the cycling of Fe and phosphate. For the discussion that fol- lows, we group the data into three categories: those regions having low bottom water oxygen concentrations and low C,, ; regions having low bottom water oxygen concentrations and high C,,; and regions having high bottom water oxygen concentrations. Furthermore, we assume that the C:P ratio of particulate organic material is - 106. We recognize that some variations in the C:P regeneration ratio could be the result of variations in the C:P ratio of the organic material; however, the preponderance of data from this region suggest that the organic matter being recycled has a C:P ratio that is indistin- guishable from Redfield matter (Fig. 5 and Table 2). For our entire dataset the C,, vs. P relationship yields a slope of 79 ( r2 = 0.58); however, by eliminating those data with negative or near-zero phosphate fluxes plus those stations having very low C,, :P ratios (Santa Monica and San Pedro basin stations) the regression yields a C,,:P ratio of 92 (r2 = 0.86), a value that is not significantly different than 106. We recognize that the large uncertainties in the C,,:P ratio preclude a detailed discussion of our data in terms of this ratio, thus we will refer to those sites having the largest deviations from C,, :P = 106, and, more specifically the pattern of phosphate fluxes in low oxygen environments.

4.1. Patterns of Phosphate Remobilization

Where bottom water oxygen is <50 PM and the rate of carbon oxidation is -1 mmol mm2 day-’ or less, the P regeneration rate can be less than that predicted from the decomposition of Redfield-type organic matter, and in some cases, there is a flux of phosphate into the sediments. Thus,


CClSta.1 3.8~,

2.8t- 0 5 10 15 20

Time (hours)

CClSta.7

0 5 10 15 20 25 Time (hours)

0 5 101520253035 Time (hours)

CC 1 Sta. 17A

4.5 CClSta.6

2.ou 0 5 10 15 20 25 30

Time (hours)

0 10 20 30 40 50 60 Time (hours)

CC 1 Sta. 17B

0 5 10 15 20 2s 30 0 5 10152025303540 Time (hours) Time (hours)

Fig. 3. Phosphate as a function of incubation time from all the study sites. Note that the different symbols represent different chambers from the same lander deployment. Error bars are analytical tlo. Arrows indicate bottom water values.

under these conditions, P appears to be preferentially sequestered in the sediments relative to organic carbon. Stations from this study demonstrating this process most strikingly include Ccl-12, CC2-2, CC2-3, TB-4, and Cat-4B; all stations having bottom water oxygen concentrations falling within a fairly narrow range between 15 and 30 PM and oxygen penetration depths of 0.6-0.1 cm (Table 1).

Where bottom water oxygen concentration is low and the rate of carbon oxidation is > 1 mmol m -’ day ’ , we observe P regeneration rates that are similar to or greater than those

expected from the regeneration of organic matter decomposition. Under these conditions it appears that P may be remo- bilized in excess of that available from organic matter alone. Examples of stations exhibiting these characteristics include SM-7, CC3-8, and CC l-5.

When bottom water oxygen concentrations are high and carbon oxidation rates are <7 mmol mm2 day -’ (and where benthic irrigation is not significant), we observe P regeneration rates that are consistent with those expected from the decomposition of Redfield organic matter.


1.0- 0 5 10 15 20

Time (hours)

CC2Sta.3

2.6- 0 10 20 30 40 50

Time (hours)

0 20 40 60 80 10 Time (hours)

2 2.8

S

4 2.6 c 2 0 % 2.4

22. '0 10 20 30 40 50

Time (hours)

CC2Sta.4 3.2

2 3.0 S

2 2.8 c 3 c 2.6

2.4- 0 10203040506070

Time (hours)

CC2Sta.6 3.0

2.7- 0 10 20 30 40 50

Time (hours)

Fig. 3. (Continued)

4.2. Iron-Oxygen-Phosphorus Interactions: The Paradigm

We propose that where bottom water oxygen is low and the rate of carbon oxidation is also low, a layer of iron oxyhydroxides may form at or immediately below the sediment-water interface. Thus, the low phosphate remobilization rates are caused in part by phosphate adsorption onto oxide surfaces as depicted in Fig. 6a.

When bottom water oxygen concentration is low and the rate of carbon oxidation is > 1 mmol mm* day-‘, oxygen penetration depth may be too shallow for the formation of a near-surface iron layer, as discussed for the previous case (Fig. 6b). Thus, phosphate initially associated with iron oxyhydroxides may escape from sediments along with the P released from the oxidation of organic matter, yielding rela-

tively high regeneration rates of phosphate. From these predictions, we would expect a flux of Fe from sediments, where there is available reducible sedimentary iron. This latter point is not subtle. Without an adequate supply of reactive iron, this scenario will not be important to P cycling.

Under conditions where bottom water oxygen concentrations are high and carbon oxidation rates are <7 mm01 mm2 day -I, we hypothesize that oxygen penetration is sufficiently deep that the iron diagenetic front will not play a major role in phosphorus cycling at the sediment-water interface (Fig. 6~). At sites where carbon oxidation rates are >7 mmol mm* day-’ and bottom water O2 > 50 PM, benthic irrigation may dominate benthic exchange. Under this condition, variations in the P regeneration rate will be influenced by the availability of sedimentary iron within the sediment irrigation zone as well as the sedimentary redox conditions within this region.

2898 J. McManus et al

Basins PE 4

0 20 40 60 80 100 Time (hours) Basins TB 4

3*8Frr”m”r”“m”“l

3.0- 0 10 20 30 40 50

Time (hours) Basins SP 7

f 3.4 2

+ 3.2 AZ 3 % 3.0

0 5 10152025303540 Time (hours)

Basins XI 7 3.3y,

Basins PE 7

2.8Y--L-

2.4’ 0 20 40 60 80 100

Time (hours) Basins TB 7

3.8

2 3.6 t

% 3.4 r:

z %-_

3.oh 0 10 20 30 40 50

Time (hours) Basins SM i

0 10 20'30 40 50 60 Time (hours)

Basins Cat 4b

X-----Y

2.8- 0 10 20 30 40 50 60

Time (hours)

Fig. 3. (Continued)

3.0t 0 5 101520253035

Time (hours)

For example, in sedimentary environments where there is a significant iron oxide component, and where the carbon oxidation rate is high and sediment irrigation is significant, P regeneration could be enhanced because of irrigation of anoxic

pore fluids that contain dissolved phosphorus that was initially associated with oxide phases. For purposes of the present study we restrict our discussion to those sites where diffusion is the primary control on sediment-water exchange.


-.e

0 5 10152025303540 Time (hours)

cc3sta.4

3.3p-.,....,.

2.7- 0 5 10 15 202530 35

Time (hours) CC 3 Sta. 8

5.5

3.0 0 5 10 15 202530 35

Time (hours) cc 3 sta. 10

2.6- 0 20 40 60 80 100

Time (hours)

0 20 40 60 80 100120 Time (hours) CC3Sta.6

0 Time (hours) cc3sta.9

3.4

r^ 3.3 a

$ 3.2 B-f %

g 3.1

3.0 0 10 20 30 40 50 60

Time (hours) cc 3 sta. 12

4.2 . . . . . . . . . . . . . . . . . . . . . . . _’

0 10 20 30 40 50 Time (hours)

Fig. 3. (Continued)

4.3. Combined Effects of Dissolved Oxygen Concentrations and Iron Cycling

It has been shown that bottom water dissolved oxygen concentrations play some, albeit indirect, role in influencing

phosphate cycling (e.g., Ingall and Jahnke, 1994, 1996; and others). For example, Ingall and Jahnke ( 1994) present data that clearly indicate that benthic regeneration of reactive P is more extensive when the sediments are overlain by waters low in dissolved oxygen. Recognizing that the phosphate


Table 2. Lander Flux Station Averages”

Station Phosphate C0: Iron Ironb (mmol m-* day-‘) (mmol me2 day-‘) (pm01 mm’ day-‘)

G:P O2 pen.’ (cm)

cc I

ccl-5 0.068 2 0.013 Ccl-6 0.084 + 0.062 ccl-7 0.019 + 0.018 ccl-12 -0.020 t 0.021 ccl-17 0.050 k 0.008

cc2 cc2-1 0.089 2 0.014 cc2-2 -0.011 + 0.010 CC2-3 0.002 ? 0.006 CC2-4 0.014 k 0.004 CC2-5 0.008 2 0.002 CC2-6 0.014 2 0.003

7.3 z 1.0 0.68 5 0.56 0.71 k 0.06 1.1 +- 0.3 2.09 2 1.03 2.42 k 0.22 0.9 2 0.8 1.59 k 0.18 1.59 t 0.48 1.1 lr 0.3 0.95 k 0.70 0.77 k 0.15 0.7 k 0.3 0.38 2 0.11 0.39 k 0.07 1.8 i 0.6 2.65 z 0.24 2.65 z 0.24

450 t- 1400 79 + 22 88 -+ 43

1142 49

Borderland Basins

PE 0.004 k 0.001 0.4 t 0.1 -0.18 2 0.28 -0.04 -+ 0.04 100 k 74 TB 0.010 k 0.005 0.9 t 0.6 1.94 t 2.82 0.38 k 0.08 90 ir 75 Cat 0.001 2 0.006 1.2 2 0.9 1.15 ? 0.07 1.11 t 0.48 1200 k 7300 SP 0.024 2 0.007 1.8 2 0.4 11.75 i 10.82 5.89 5 1.13 75 2 28 SM 0.099 + 0.015 1.7 2 0.4 12.18 2 6.82 9.08 k 1.27 17 t 5 SC1 0.007 + 0.002 1.1 z 0.3 0.29 t 0.10 0.19 i 0.05 157 2 62

cc 3d CC3-2 0.025 k 0.008 cc3-3 0.010 2 0.001 cc3-4 0.022 ? 0.002 CC3-6 0.015 + 0.004 cc3-8 0.179 2 0.007 cc3-9 0.009 2 0.001 cc3-10 0.005 k 0.001 CC3-12 0.022 i 0.014

2.4 + 0.9 0.7 I 0.4 1.6 -f 0.7 2.2 k 0.8 2.0 5 0.8 0.9 + 0.4 0.3 t 0.4 1.3 ? 0.6

2.00 k 2.00 -0.10 k 0.14 114 -’ 48 0.54 -0.10 F 0.13 -0.05 5 0.08 70 k 41 2.91

0.16 k 0.72 -0.04 -+ 0.12 73 F 32 1.10 10.97 -+ 6.86 6.52 + 1.72 147 t 57 0.2 1 17.97 + 4.69 17.76 k 3.31 12 + 5 0.26 0.08 k 0.07 0.09 2 0.02 100 + 46 I .28 0.06 k 0.51 -0.09 k 0.05 60 -t 81 3.66 0.90 2 0.21 0.85 2 0.21 59 i 35 0.57

3.8 5 2.9 6.0 k 1.5 1.8 k 1.8

6.1 5 1.3

1.53 + 1.82 43 2 7

11.3 2 1.00 0.15 i- 0.76 1.89 k 1.55

0.25 2 1.04 41 -t 6

10.61 + 1.30

56 2 44 71 256 95 -’ 131

0.93 2 0.24 122 k 33

82 r 17

0.22 0.27 0.07 0.16

0.51 0.39 1.19 0.94 0.48

2.87 0.4 1 0.61 -

0.16 1.28

a Averages are calculated by weighting each chamber flux equally; uncertainties are 1~. b Averages are weighted based on the uncertainty of each individual chamber flux. Note that, in general, there is no significant difference

between the weighted and unweighted averages, but that the weighted uncertainties tend to be smaller. ’ C,, and oxygen penetration depths are taken from Berelson et al. (1996). d Stations CC3-8-12 correspond to Santa Monica Basin, San Clemente Basin, Patton Escarpment, and Tanner Basin, respectively.

system has a variety of other complicating factors acting upon it (e.g., Van Cappellen and Ingall, 1994, 1996; Ingall and Van Cappellen, 1990), we evaluate the evidence in our dataset for this paradigm.

Our data do indeed show that, for some environments, at low bottom water oxygen concentrations, the phosphate flux is elevated; in point of fact the highest benthic phosphate fluxes and the lowest C,,:P regeneration ratios occur in the low-oxygen Santa Monica basin (Tables 1, 2). The San Pedro basin likewise displays somewhat elevated phosphate fluxes for this region relative to that expected for organic matter decay. Our data also suggest, however, that at slightly higher but still low bottom water oxygen concentrations surface sediments may be sequestering phosphate. Furthermore, from the compiled dataset of Ingall and Jahnke ( 1996) and that presented here, both low (zero) and high phosphate fluxes can be observed at higher bottom water oxygen concentrations.

the eastern Pacific boundary. Their compiled dataset is from the California and Washington margins. One of the more notable features of the dataset presented in their work is the low phosphate fluxes relative to organic carbon oxidation at the higher values of C,,. Note also that these high C,,, and low P flux values occur in relatively shallow continental shelf/slope regions. The significance of the addition of their data to ours is that these high organic carbon/low phosphate remineralization sites have high bottom water oxygen concentrations. This observation is consistent with our paradigm described above (Fig. 6) that the depth of oxygen penetration and iron cycling will be major influences on phosphate recycling. Essentially, the sites in this region having high C,, and low P fluxes also have shallow oxygen penetration depths (<0.5 cm; Archer, 1990; Archer and Devol, 1992). Furthermore, because these locations are near the mouth of the Columbia River, these sediments are likely to have a significant sedimentary iron component.

This latter point is better illustrated by the relationship For the sites studied here it would appear that for oxygen between phosphate remobilization and organic carbon oxida- penetration depths less than 0.3 cm, at least some Fe may tion (Fig. 7). By adding the data of Ingall and Jahnke ( 1996) escape from the pore fluids (Fig. 8a), thus at these oxygen we can get a more complete picture of this relationship along penetration depths we might expect P regeneration rates to be


Phosphate flux (mmol m’ ’ day’ ‘) 0.00 0.05 0.10 0.15 0.20

Fig. 4. (a) Phosphate and (b) iron regeneration rate as a function of depth for our study area. Note both the high and low phosphate fluxes within the oxygen minimum zone (500- 1000 m) and the high iron fluxes within the oxygen minimum zone.

greater than those expected from the degradation of organic material alone. Although Santa Monica basin is one station where this expectation is clearly met, all the stations having high phosphate fluxes relative to the organic matter degradation rate (Ccl-6, Ccl-7, SP-7, SM-7, TB-7, CC3-8) also have high benthic iron fluxes, as compared to the rest of our data set. However, as shown in Fig. 8b shallow oxygen penetration depths do not necessarily yield high P regeneration rates. Although it is clear that the highest P fluxes are indeed at the shallow oxygen penetration depths, the negative

0.15

0.10

0.05

0.00

-0.05

I ‘I ” I ’ ” x I ” ” I, I, I ‘, ” I. ‘I, I

\ Fe’*

i

Fig. 6. (a-c) Schematic diagrams demonstrating proposed geochemical conditions that will allow enhanced phosphate burial.

P fluxes are also found at sites having oxygen penetration depths that range between 0.1 and 0.6 cm (Stations CCl-

0 1 2 CO x &noI’ me ’ Lay* I,”

7 8 12, CC2-2, CC2-3, TB-4, Cat 4B; Fig. 8b). Although there is overlap in the range of oxygen penetration depths dis-

60 _ [Concentration]------+ Phosphate uptake?

Depth (en;) I

(W -[Concentration]- Fe flux

, *Sediment-water interfxc

Conditions 1. High carbon oxidation rate 2. Low bottom water oxygen

Depth Cc

I I Predictions I Geochemical conditions prohibit

the formation of a near-surface iron layer and result in a flux of iron from sediments. Also, phosphate associated with oxyhydroxides will escape from sediments producing an excess phosphate flux over that predicted from the oxidation of organic carbon alone..

Conditions 1. No significant irrigation 2. Low carbon oxidation rate 3. LAW bottom water oxygen

gredictiom Formation of iron layer at or near the sediient-water interface, thus providing sites for phosphate scavenging.

03 ~ [Concentration]----r

I Sediment-water interface

Conditions 1. Carbon oxidation rates < 7 mm01 m”day-I 2. High bottom water oxygen concentration 3. Oxygen penetration depths greater than in previous two cases.

PredictiQgs C.,:P regeneration ratios consistent with degradation of Redfield organic matter

Fig. 5. Phosphate flux as a function of the organic carbon regeneration rate. The solid line represents a C,, to P regeneration ratio of 106.

playing both high and low phosphate fluxes, sites having high P fluxes do have an average oxygen penetration depth that is shallower than stations showing low P fluxes (e.g.,


i a -o.ost,~‘~““““‘~‘I”‘~“‘I’~‘I”,i 0 2 4 6 8 10 12 14 16

Cm, (mmol m" day“)

Fig. 7. Phosphate flux as a function of organic carbon respiration rate for our study sites as well as for those of Ingall and Jahnke (1996).

0.23 ? 0.13 cm vs. 0.42 2 0.17 cm for the high vs. low P fluxes).

4.4. Porewater Data

It is generally recognized that porewater phosphate profiles have an artifact that renders these profiles not useful for quantitative calculations of the diffusive flux across the sediment-water interface (e.g., Jahnke et al., 1982). Many of the artifact observations to date are for deep-sea sediments: both carbonate rich and carbonate poor. The reasons for the artifact are largely speculative but may be associated with decompression (Jahnke et al., 1982) or adsorption during centrifugation. In accord with these observations, porewater phosphate concentrations of the top one or two points for nearly all our study sites are below bottom water concentrations, which, if not an artifact, would suggest large phosphate uptake fluxes, which we do not observe. Even if porewater profiles were unaffected by artifacts, where profiles are collected from suboxic-anoxic regions adsorption-desorption processes associated with iron cycling very near the sediment-water interface will alter phosphate mobility across the iron redox boundary (e.g., Bemer, 1977) thus making quantifying phosphate transport across the sediment-water interface impossible. An example of this latter processes is from Santa Monica basin. Here the predicted diffusive flux from Santa Monica basin is larger than that measured with the lander (-0.3 mmol mm2 day-‘). Despite these uncertainties in interpreting porewater phosphate profiles, phosphate porewater profiles may provide qualitative to semiquantita- tive information on phosphate diagenesis.

Porewater data from the borderland basins reveal patterns consistent with the idea that phosphate recycling is influenced by iron cycling in this region. Porewater profiles from Santa Monica and San Pedro Basins demonstrate high concentrations of Fe within the top 1 cm of sediment (Fig. 9). This general observation is consistent with a significant flux of Fe from these sediments, and, in the case of Santa Monica, phosphate that is being released in excess of that predicted

from the decomposition of organic material. In San Pedro Basin, it appears that less Fe is escaping the sediments as compared to Santa Monica, as indicated by the lower near- surface iron concentration in San Pedro Basin porewaters. Furthermore, in San Pedro Basin both porewater profiles exhibit large excesses of P relative to that expected from the decomposition of organic matter as modeled by the ammonia profile (Fig. IOa,b). Also the core 2 porewater phosphate profile exhibits a down core decrease in phosphate concentration suggesting down core removal of phosphate. The companion profile shows a P increase between 5 and 25 cm; however, this increase is still less than that expected from the degradation of organic material (Fig. lob). These results for San Pedro basin suggest an additional near-surface source of P other than just the decomposition of organic material, along with a down-core sink for dissolved P. Thus, both Santa Monica Basin and San Pedro Basin are locations where excess phosphate relative to that expected from the decomposition of organic matter is coming from the sediments.

The Patton Escarpment profile shows little porewater iron, and we don’t expect much interaction between Fe and phosphate cycling at this location. Any solid phase iron arriving at the seafloor will probably not exchange (adsorption/desorption) a significant amount of phosphate in the upper sediment column.

San Clemente Basin exhibits a porewater iron profile consistent with the hypothesis that the iron redox boundary is

2 50- ‘h 4 40 1

0 (a) 1

8

A 4.0

(cm)

Fig. 8. (a) Iron and (b) phosphate flux as a function of oxygen penetration depth.


Catalina Basin

Phosphate (pM) San Pedro

Phosphate (PM)

0 5 10 15 20 25 30

0 50 100 150 0 10 20 30 40 50

Fe WM) Fe WM)

Santa Monica

Phosphate (FM)

0 50 100 150

Core 1, P 35 III.III~ll~..~~I~(~.I~...

50 100 150 200 250 300

Fe 0.W

San Clemente

Phosphate (PM)

2 4 6 8 10 12 14

Tanner Basin

Phosphate (PM)

I\ 0 5 10 15 20 25 30

Patton Escarpment

Phosphate (PM)

5 10 15

3

25 25

0 1 2 3 4 5

Fe (PM)

30 - 0 10 20 30 40 50

Fe @Ml

0 1 2 3 4 5

Fe UM)

Fig. 9. Porewater phosphate and iron data as a function of depth in the sediments. Open circles and circles with dots represent duplicate cores. Note the different axes for both phosphate and iron.

too deep (5-6 cm) to play a significant role in phosphate transport across the sediment-water interface. Furthermore, it appears from the comparison of the porewater phosphate to that derived from modeling the ammonia profile (Fig. lOc,d) , that porewater phosphate is being removed relative to the amount predicted by organic matter decomposition below approximately 5 cm depth in this basin.

Tanner and Catalina Basins, like San Pedro basin, have elevated iron concentrations fairly close to the sediment- water interface; however, given the uncertainties in our dataset it is not possible to discern what role, if any, iron cycling may have on phosphate cycling in these locations. It should be noted, however, that both Tanner and Catalina basin porewater profiles have coincident Fe and phosphate maximums suggesting a coupling between their cycles within these sedi-

ments. In addition, although these basins clearly exhibit porewater patterns consistent with down-core P removal, other porewater data suggests that the down-core decrease in these constituents could be caused by macrofaunal irrigation.

In summary, porewater phosphate data suggest that P cycling in this region is being influenced by iron cycling. The high porewater iron concentrations in San Pedro and Santa Monica basins make these basins likely locations for iron- derived phosphate. An additional, nonbiogenic, source of porewater phosphate is consistent with the elevated phosphate concentrations relative to the ammonia concentrations observed in San Pedro Basin. Likewise, Catalina and Tanner basins both have phosphate maximums that coincide with porewater iron maximums, suggesting that iron cycling is playing at least some role in P cycling in these basins.


Ammonia (PM) 0 100 200 300 400

Phosphate (PM) ‘“a 100 150 200

0“““““““” cr’ %

0 ??e (b) -j

Ammonia (KM) Phosphate QIM)

25 San Clemente

??o \ ?? 0

??o

??o

??o

0

Fig. 10. Ammonia and phosphate as a function of depth in (a, b) San Pedro and (c, d) San Clemente Basins. Porewater ammonia profiles were modeled using an exponential concentration-depth function after Bemer ( 1980) and Schuffert et al. (1994), where the concentration (C) of a constituent is described as a function of depth (n) using, C = C,, + a,( 1 - e-“‘). Using the results from the ammonia profile, we modeled the phosphate profile assuming that the ammonia profile will produce a down-core distribution of porewater phosphate consistent with the decomposition of organic matter. The resulting theoretical porewater profiles are indicated by the lines in (b) and (d). Because of the relatively small difference between the two ammonia profiles from San Pedro Basin we only present the results from a single fit. This steady-state model accounts for adsorption, diffusion, and sediment burial. We assume first-order reaction kinetics, negligible advective flow, and no irrigation. Further details of the model are described extensively elsewhere (Bemer, 1980; Schuffert et al., 1994). The values used for the diffusion coefficient at infinite dilution (Do) are from Krom and Bemer (1980) where DO( x 10e6) = 19.8 + 0.40( T - 25°C) for ammonia and &( X10m6) = 7.34 + O.l6(T - 25°C) for phosphate. In San Pedro Basin, it is shown that in the upper few centimeters of sediment there is a large excess of phosphate relative to that predicted from the ammonia profile. Further down core, we observe decreases in porewater phosphate (core 2) or increases that are smaller than those expected from the stoichiometric decomposition of organic material (core 1 ), thus indicating a down-core sink for P in this basin. In San Clemente Basin (c, d), P concentrations at depths below -5 cm are less than those expected from the decomposition of organic matter, thus suggesting down-core removal of phosphate at this site as well. Similar plots for other sites would not be meaningful because phosphate profiles at Tanner and Catalina Basins clearly reflect down-core removal; however, in the case of these basins, removal could be caused by either authigenic precipitation of P or because of porewater irrigation. At the Patton Escarpment site there is no significant ammonia gradient.

4.5. Significance of Interactions among Oxygen Penetration, Iron, and Phosphorus

As has been pointed out previously and as is explicit in our work, the phosphorus system is influenced by a variety

of factors. We have suggested that iron cycling may be an important, although not the sole, constraint on the phosphorus system along the eastern Pacific boundary of North America. For example, It has been previously argued that, although Fe may be important it can not be the sole source


or sink of phosphate in excess of organic material in this region (Ingall and Jahnke, 1996). As Ingall and Jahnke (1996) show, P associated with Fe in Santa Monica basin can only account for -30% of the total regenerated P, which, when coupled to the P expected from organic matter degradation, is still insufficient to account for the large observed P fluxes reported here and in their work. We should note, however, that Santa Monica basin is enigmatic in our study as well as in that of Ingall and Jahnke ( 1996). If, for example, we removed all the Santa Monica data from our discussion, some of the highest phosphate fluxes relative to organic matter decay would be from locations where bottom water oxygen concentrations range from 50 to 150 PM, but again these sites have relatively shallow oxygen penetration depths (e.g., Ccl-5 and -6; Tables 1, 2).

The focus on oxygen penetration depth can be somewhat confusing if not misleading because at relatively shallow oxygen penetration depths, we assert that phosphorus burial efficiency can be increased or decreased depending on other, poorly quantified, factors. On average, at oxygen penetration depths <0.3 cm we would expect P fluxes in excess of those anticipated with the degradation of organic material, consistent with the data of Ingall and Jahnke ( 1994, 1996). However, at oxygen penetration depths that encompass this depth range to slightly deeper (0.1-0.6 cm), the opposite trend may result. While the cycling of Fe under these conditions will undoubtedly play some role in phosphate cycling, the redox sensitivity of certain microorganisms will likely influence phosphate cycling as well (Ingall and Jahnke, 1994, 1996). How these two processes vary in importance with oxygen penetration depth is as yet unknown.

We suggest that at high bottom water oxygen concentrations along the continental shelf, in places where oxygen penetration is shallow and ample reactive iron is present, iron cycling will dominate sediment-water phosphorus exchange. We should note however, that the role of Fe may simply be transient, and that the predominant end product is more likely to be authigenic apatite or some other phase that will ultimately remain within the sediment (e.g., Ingall and Jahnke, 1996; Reimers et al., 1996). This scenario has previously been reported in an environment having high bottom water oxygen concentrations, but shallow oxygen penetration depths (Heggie et al., 1990). At this location, the biogeochemical cycling of Fe results in a near-surface sedimentary trap for Fe. Phosphate and fluoride are scavenged by the iron oxides near the sediment-water interface. As this phase is subsequently buried and undergoes reductive dissolution, phosphate and fluoride are released for incorporation into apatite. The dataset presented here are inadequate to assess the potential importance of this process along the California margin. However, the data do clearly indicate that such a process is possible along this margin, and it is possible that such a process is important elsewhere along the eastern boundary of the Pacific. Thus, the interplay between Fe and phosphate may provide a mechanism for the ultimate burial of phosphate in authigenic apatite in this region.

Under high oxygen conditions where oxygen penetration is deep, increased retention of P is probably more dependent upon the storage of P by microorganisms as the end result of oxidative phosphorylation (e.g., Ingall and Jahnke, 1994).

One example of a location where P remobilization is much less than that expected from the degradation of organic matter is in the equatorial Pacific (e.g., Hammond et al., 1996). Here, the C,,:P regeneration ratio was found to vary between two separate cruises (360 t 170 vs. 113 ? 30) suggesting that, whatever process is responsible for this variation, it is not necessarily a steady-state phenomenon.

As with high oxygen environments, low oxygen environments also seem to produce two possible results with respect to phosphorus cycling. As noted earlier there is a substantial body of literature demonstrating the contemporary accumulation of authigenic apatite in low oxygen environments, yet there is also a growing body of literature suggesting that benthic regeneration of reactive P is more efficient when bottom water oxygen concentrations are low (e.g., Ingall and Jahnke, 1994, 1996). In their study, Reimers et al. ( 1996) note that a critical factor in mineral paragenesis is that iron oxyhydroxides are reduced, which raises pH and releases sorbed phosphate, fluoride, and Fe*+. Thus, as has been shown before, a variety of conditions are required for phosphate uptake by sediments during diagenesis; these include: the presence of reactive iron, perhaps high concentrations of organic matter, or the presence of a microbial mat (e.g., Froelich et al., 1988).

Finally, the question of steady-state may also be important to our study. We might expect that variations in oxygen penetration depth, due to seasonal variations in primary pro- ductivity will lead to variations in the sign and magnitude of P regeneration in our study area. We note that in a recent particle flux study in the Monterey Bay region, Pilskaln et al. ( 1996) report organic carbon export fluxes that vary by more than an order of magnitude. Such large variations in organic carbon delivery to the seafloor could produce variations in the oxygen penetration depth that could change the sign of the phosphate flux in a given region. Furthermore, in Santa Barbara basin, Reimers et al. ( 1996) have interpre- ted their data as an indication of the nonsteady-state nature of sediment surface interactions that may lead to apatite formation. One way to assess whether or not the fluxes reported here are at steady-state is to measure sediment phosphate contents. For sites having P-fluxes into the sediments, we would expect significant sediment P contents (perhaps as high as 2%). While P contents of this magnitude might imply a significant calcium phosphate (apatite) sedimentary component, we have no way of confirming this possibility. Unfortunately, we do not have sediment samples from the locations where we observe very small or negative P fluxes. Thus, we have no way of knowing if P uptake is related to phosphorus-iron interactions or phosphorus uptake via apatite precipitation or both processes.

5. CONCLUSIONS

Data presented here suggest that suboxic sediments have the potential to either increase or decrease the burial efficiency of phosphorus. It was found that under conditions of low bottom water oxygen concentration and when the rate of organic carbon oxidation is - 1 mm01 m-* day -’ or less, P may be preferentially buried relative to organic carbon. At stations where bottom water oxygen is low and the degra-


dation rate of organic material is greater than 1 mmol mm2 day -’ , P may be released from sediments at a rate exceeding the production expected from the oxidation of organic matter. At stations having high bottom water oxygen concentrations, rates of organic matter decomposition <7 mmol m-* day -‘, the C,,:P regeneration ratio is consistent with the Redfield ratio. At high bottom water oxygen concentrations and high rates of organic matter decomposition, which yield shallow oxygen penetration depths, phosphate appears to be buried preferentially relative to organic carbon. While it is clear from this and other studies that oxygen penetration depth and iron cycling are important to consider in the early diagenetic behavior of phosphate, the longer-term effects on phosphate burial are as yet uncertain.

Acknowledgments-Support for this research was provided by N.S.F. grants to J.M., W.B., K.J., and K.C. We’d like to thank D. Burdige, D. Colbert, G. Elrod, J. Nowicki, and S. Tanner for analytical and field assistance and the captains and the crews of the R.V. Point Sur and R.V. New Horizon for their assistance. The thoughtful criticism of P. Froelich, P. van Cappellen, and two anonymous reviewers are greatly appreciated.

P. N. Froelich

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