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Keywords:

  • subduction;
  • metamorphism;
  • trace elements;
  • nitrogen isotopes;
  • diagenesis

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[1] Toward understanding of the subduction mass balance in the Izu-Bonin-Mariana (IBM) convergent margin, we present an inventory of N and C concentrations and isotopic compositions in sediments obtained on Ocean Drilling Program (ODP) Legs 129 and 185. Samples from Sites 1149, 800, 801, and 802 contain 5 to 661 ppm total N (organic, inorganic combined) with δ15NAir of −0.2 to +8.2‰ (all δ15N values <+2.5‰ from Site 800). At Site 1149, N content is higher in clay-rich layers and lower in chert and carbonate layers, and δ15N shows a distinct down-section decrease from 0 to 120 mbsf (near +8.0 at shallow levels to near +4.0‰). Reduced-C concentration ranges from 0.02 to 0.5 wt.%, with δ13CVPDB of −28.1 to −21.7‰. The down-section decreases in δ15N and N concentration (and variations in concentrations and δ13C of reduced C, and Creduced/N) at Site 1149 could help reconcile differences between δ15N values of modern deep-sea sediments from near the sediment-water interface and values for forearc metasedimentary rocks. At Site 1149, negative shifts in δ15N, from marine organic values (up to ∼+8‰) toward lower values approaching those for the metasedimentary rocks (+1 to +3‰), are most likely caused by complex diagenetic processes, conceivably with minor effects of changes in productivity and differing proportions of marine and terrestrial organic matter. However, the forearc metamorphic suites (e.g., Franciscan Complex) are known to have been deposited nearer continents, and their lower δ15N at least partly reflects larger proportions of lower-δ15N terrestrial organic matter. Subduction at the Izu-Bonin (IB) margin, of a sediment section like that at Site 1149, would deliver an approximate annual subduction flux of 2.5 × 106 g of N and 1.4 × 107 g of reduced C per linear kilometer of trench, with average δ15N of +5.0‰ and δ13C of −24‰. Incorporating the larger C flux of 9.2 × 108 g/yr/linear-km in carbonate-rich layers of 1149B (average δ13C = +2.3‰) provides a total C flux of 9.3 × 108 g/yr/linear-km (δ13C = +1.9‰). Once subducted, sediments are shifted to higher δ15N by N loss during devolatilization, with magnitudes of the shifts depending on the thermal evolution of the margin.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[2] Subduction of sedimentary rocks is an important factor in the global cycles of nitrogen (N) and carbon (C), as could be hydrothermally altered oceanic crust (AOC) [see Javoy et al., 1986; Zhang and Zindler, 1993; Bebout, 1995; Javoy, 1998]. Despite much recent interest in cycling of volatile elements at convergent margins, there has been little previous study of the unmodified inputs of N and C (especially N) in subduction zones. Any previous attempts to determine the fluxes of these elements through convergent margins have required conjecture based on analyses of seafloor lithologies that may not be appropriate or estimates based on studies of low-grade metamorphic rocks. The quantification of true input fluxes of C and N into subduction zones, preferably on the scale of individual convergent margins, becomes all the more urgent in the context of recent attempts to mass balance inputs and outputs of the elements at individual convergent margins. Therefore, to better understand the role of subduction on the global chemical cycles of these elements, we attempt to produce an inventory of N and C contents and isotopic compositions in sediments recovered on Ocean Drilling Program Leg 185 (and Leg 129) outboard of the Izu-Bonin-Mariana (IBM) subduction zone. This inventory, combined with knowledge of devolatilization at greater depths in paleo-subduction zones (based on study of metamorphic suites [see Bebout and Fogel, 1992; Sadofsky and Bebout, 2003]), and outputs of volatiles in arc magmas [e.g., Fischer et al., 2002; Hilton et al., 2002; Snyder et al., 2003], will help to better determine the magnitude and mass balance of N and C in convergent margins and elucidate the global cycling of these elements.

[3] The IBM arc system is formed through subduction of old (>170 Ma) oceanic crust of the Pacific plate beneath the Philippine plate (Figure 1). This section of the oceanic crust, along with the overlying sediment section, has been sampled outboard of the Izu-Bonin subduction zone, by Ocean Drilling Program (ODP) Site 1149, and by ODP Sites 800, 801, and 802 outboard of the Mariana Trench (see locations on the seafloor map in Figure 1). The IBM subduction system provides an ideal location for attempts to mass balance elements fluxes through the “subduction factory” because, unlike many other subduction zones, there appears to be no accumulation of sediments in an accretionary prism [see von Huene and Scholl, 1991]. The sedimentary section at site 1149 is 408m thick, and divided into 4 units by the Leg 185 Scientific party. Unit I (0–118 mbsf) is Pleistocene – late Miocene in age and composed of mixed volcanic ash, siliceous oozes, clay, and minor amounts of silt. Unit II (118–179 mbsf) is less well dated, and dominated by dark brown pelagic clay, with some volcanic ash and more minor radiolarians and diatoms. Unit III (179–282 mbsf) is composed of radiolarian chert and zeolitic clays, recovery in this section was relatively low, and these cherts are likely more minor reservoirs of N and C. Unit IV (282–408 mbsf) is composed of radiolarian chert, marl, and chalk overlying the oceanic crustal basement. Thus the section at site 1149 is a nearly classical sequence of oceanic sediment from the deeper parts of the ocean, with cherts and limestones near the bottom, more clayey layers above that and finally some input of volcanic ash. For our attempt to inventory the flux and isotopic compositions of C and N subducting in the modern Izu margin, we obtained samples from cores of the complete sediment section obtained at Site 1149 during ODP Leg 185 (see description of scientific goals and data obtained on this drilling leg by Plank et al. [2000]), and for comparison, a smaller number of samples of sediment at ODP Sites 800, 801, and 802 collected during Leg 129 and kept frozen since their collection.

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Figure 1. Map of IBM convergent margin and adjacent seafloor, showing bathymetry (see color-coded legend), structures and topography in the seafloor basement, the locations of the volcanic arc, and the locations of Sites 800, 801, 802, and 1149 (image used courtesy of the ODP; figure from Plank et al. [2000], where a detailed description of the data sources is provided).

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[4] Nitrogen is bound into sediments primarily in organic detritus, but other complex ions, such as NO3, can be adsorbed onto sedimentary particles. The primary δ15N of sediments depends on the relative proportions of N from organic sediments, with different values depending on type of organism and food source, and N from other types of particles, such as marine nitrate bound onto clay minerals. Nitrogen fixation by cyanobacteria and atmospheric input of N into the ocean tends to initially create a reservoir with δ15N near atmospheric values, denitrification of a marine reservoir tends to preferentially return 14N to the atmosphere thereby increasing the δ15N of the remaining reservoir, grazing increases δ15N at each trophic level, and diagenetic alteration has been shown to either increase or decrease δ15N depending on redox conditions (see summary by Ettwein et al. [2001, and references therein]). Variations in the proportions of marine and terrestrial organic inputs are also known to affect δ15N in sediments [see Peters et al., 1978; Minoura et al., 1997; Pride et al., 1999].

[5] Many previous studies have documented or considered sedimentary, including diagenetic, N-isotope behavior [Peters et al., 1978; Macko, 1989; Muzuka et al., 1991; Williams et al., 1995; Luther et al., 1997; Minoura et al., 1997; Wilson and Thomson, 1998; Thamdrup and Dalsgaard, 2000; Freudenthal et al., 2001]. However, most of these studies have not focused on the deep-sea lithologies most likely to be subducted, that is, in the sedimentary environments similar to those outboard of modern subduction zones (dominated by the circum-Pacific). This is the first study of its kind to examine N and δ15N in sedimentary lithologies deposited outboard of a subduction zone and lacking a metamorphic overprint related to complex subduction and exhumation history (the latter experienced by the paleoaccretionary metamorphic suites). It is also the first to investigate in detail the deep-sea sediment representative of an ocean-ocean subduction zone setting; most of the previous work, including the work to date on paleoaccretionary suites (e.g., work on the Catalina Schist, Franciscan Complex, both in California [Bebout et al., 1999a; Sadofsky and Bebout, 2003]), has been conducted on sedimentary lithologies deposited near continents (in general, experiencing higher clastic sedimentation rates and with the greater potential for addition of terrestrial organic matter). The wide range in sedimentary δ15N values from the literature (representative summary in Figure 2) illustrates the potential pitfall in using an averaged modern sedimentary N composition, based on sediment data for a wide variety of sedimentological environments, to estimate the fluxes of N through subduction zones. Also, the sedimentary δ15N values from the literature are obviously different from (mostly higher than) values observed for low-grade metasedimentary rocks from subduction-zone settings for which deep-sea sediments are thought to be the protoliths [Bebout and Fogel, 1992; Sadofsky and Bebout, 2003]. Thus, as an additional goal of this study, we sought an explanation for the disparity in δ15N illustrated in Figure 2, with possibilities including, (1) that the other studies were carried out on sediments not representative of sedimentological processes and organic deposition on the deep seafloor outboard of trenches, and perhaps were conducted on inappropriate sedimentary lithologies, and (2) that diagenetic and/or very low-grade metamorphic processes altered (decreased) the δ15N of the seafloor sediment section during its initial transit into the trench and to depths of 5–40 km represented by the paleoaccretionary suites. In this paper, we present new C and N concentration and stable isotope data for deep-sea sedimentary sections obtained on ODP Legs 129 and 185 and discuss possible causes of variations observed in these sections in the context of the sedimentological and tectonic setting in which the sediments were deposited. We then use the concentrations and isotopic compositions to estimate the C-N fluxes into the Izu-Bonin convergent margin, based on the data for Site 1149 (obtained on Leg 185), and provide a synthesis of recent observations regarding the effects of forearc metamorphic processes on the deeper entrainment of N based on studies of metamorphic rocks exposed in paleoaccretionary complexes.

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Figure 2. Nitrogen isotopic composition of marine sediments, compared with those for low-grade metamorphic rocks from subduction zone settings. Surface sediments off the shore of North America, ovals show mean and standard deviation for several basins off the western shore of North America [Peters et al., 1978]. Oman Margin, mean (line) and standard deviation (field) are shown for each core. Data are from organic-rich sediments at depths of up to 200 m [Muzuka et al., 1991]. Japan Sea, data are from the top 6 m of sediment [Minoura et al., 1997]. S.E. Atlantic data are from cores between the Canary Islands and the Moroccan Coast [Freudenthal et al., 2001]. Baffin Bay and Labrador Sea data are collected from depths of up to 60 m from ODP Leg 105 [Macko, 1989]. Franciscan Complex and Western Baja Terrane [Sadofsky and Bebout, 2003]. Catalina Schist, LA = lawsonite-albite facies, LBS = lawsonite-blueschist facies [Bebout and Fogel, 1992; Bebout, 1997].

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2. Analytical Methods

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[6] Multiple samples of each sedimentary lithology were recovered from newly drilled core from site 1149 and from archived, frozen collections from sites 800, 801, and 802 (the latter obtained from the ODP Gulf Coast core repository, College Station, Texas, and from J. Alt; all samples had been kept frozen since collection). In our laboratory, the frozen sediment samples were thawed and dried at 45°C for approximately 48 hours and then crushed finely enough to homogenize for stable isotope analyses.

[7] Nitrogen isotopic compositions of total N (combined organic and inorganic fractions) were obtained by sealed-tube combustions (910°C for 4 hours [see Bebout and Fogel, 1992; Sadofsky and Bebout, 2000; Bebout and Sadofsky, 2004]). Nitrogen yields were determined by measuring the signal at m/z = 28, for calibrated inlet volumes in the mass spectrometer. Combustions at higher temperatures (up to 1060°C), for one sediment sample, produced δ15N values and N concentrations indistinguishable from those for the lower-temperature experiments; thus all other combustions for N were performed at the lower temperatures (910°C). For reduced-C analyses of δ13C, samples were first reacted overnight in 10 ml of 1N HCl to remove any carbonates. The samples were then centrifuged and rinsed three times to remove all HCl, then dried and prepared for isotopic analysis by sealed-tube combustion (850°C for 1.5 hours). Concentrations of C were determined by Hg manometry after cryogenic purification of the resulting CO2. CO2 from carbonate-bearing samples was prepared by overnight dissolution in 100% phosphoric acid at 25°C, then purified and measured in the same manner as the CO2 from the reduced-C experiments. All gases were analyzed for their isotopic compositions in dual-inlet mode on the Finnigan MAT 252 mass spectrometer at Lehigh University, and isotopic compositions are presented as δ15NAir, δ13CVPDB, and δ18OVSMOW. Uncertainties, determined by replicates of internal and international standards, are ≤0.10‰ (1σ) for all isotope analyses. For the concentration data, detection limits are, for C, <0.1 wt.%, and for N, on the order of several ppm (depending somewhat on sediment type and preparation involved).

3. Results

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[8] Forty bulk-sediment samples from sites 1149, 800, 801, and 802 contain 5 to 661 ppm total N (organic and inorganic N fractions combined) with δ15N ranging from −0.2 to +8.2‰ (Table 1, Figure 3a; the only two samples with δ15N < +2.5‰ are from Site 800). Various ash and clay units recovered from Site 1149A contain abundant N (>200 ppm) and show some variation in concentration as a function of depth within individual lithologies, particularly within ash in the upper 120 m of the section (see Figures 3a and 4a). Clastic (ash and clay) samples from Sites 800, 801, and 802 have highly variable N concentrations N, whereas chert, carbonate, and some other pelagic clay samples are generally lower in N than samples from site 1149 (Table 1; Figures 3a and 4a). At Site 1149A, δ15N varies with depth in Unit I (the upper 120 m of the section), with the sample closest to the sediment-water interface having δ15N of +8.2‰, and showing decrease with depth toward values of +4.7‰ at ∼120 meters (Figures 3a and 4b). At depths greater than ∼120 mbsf, δ15N ranges from +3.7 to +4.9‰ in clastic layers, and δ15N values of chert and carbonate layers at greater depths are lower and more variable (+2.5 to +4.6; Figure 4b).

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Figure 3. Whole-sediment N (a) and C (b) concentrations and corresponding isotopic compositions of all samples analyzed in this study (Sites 800, 801, 802, and mostly 1149), as a function of lithology (for lithologies of Sites 800, 801, and 802 samples, see Table 1).

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Figure 4. Nitrogen concentration and isotope data for the sediment section, and ammonium concentrations from the same depth interval, at Site 1149. (a) Sediment N concentration for Site 1149. Note slight trend toward lower N content as a function of depth. (b) Nitrogen isotopic composition of samples recovered from Site 1149. Note the shift in δ15N in the upper ∼100 mbsf, mostly within the ash- and biogenic-silica-bearing clay part of the section. Below 180 m depths, lithologies are radiolarian cherts and nannofossil chalk/marl, both containing low N concentrations with relatively low δ15N. (c) Ammonium concentrations in interstitial waters in the upper 200 m of the Site 1149 sediment section, demonstrating maximum concentrations near 50 m, resulting from organic breakdown, and decreased concentrations at greater depths related to sequestration by diagenetic clays (data from Plank et al. [2000]).

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Table 1. Carbon and Nitrogen Concentration and Stable Isotope Compositions of Sediments From Sites 1149, 800, 801, and 802
LithologyaSample IdentificationSectionIntervalDepth, mbsfNitrogenReduced CarbonCred/NCarbonate
LegSiteHoleCoreppmδ15Nairδ13CVPDBWt.%Wt.%δ13CVPDBδ18OVSMOW
Leg 185, Site 1149 (Holes A, B, C)
Ash- and Biogenic Silica-Bearing Clay1851149A001H01WR140–15013308.2−23.60.134.0   
Ash- and Biogenic Silica-Bearing Clay1851149A001H03W30–3233716.1−22.60.359.4   
Ash- and Biogenic Silica-Bearing Clay1851149A002H01WR140–15063506.7−22.10.298.3   
Ash- and Biogenic Silica-Bearing Clay1851149A004H02WR140–150264626.1      
Ash- and Biogenic Silica-Bearing Clay1851149A004H03WR114–116274486.2−22.20.357.8   
Ash- and Biogenic Silica-Bearing Clay1851149A004H03WR140–150285086.1−22.20.244.7   
Ash- and Biogenic Silica-Bearing Clay1851149A005H04WR140–150393955.8−22.50.215.4   
Ash- and Biogenic Silica-Bearing Clay1851149A05H05WR140–150403665.7−22.70.205.3   
Ash- and Biogenic Silica-Bearing Clay1851149A006H02WR140–150453836.3−21.70.359.0   
Ash-Bearing Siliceous Clay1851149A007H03WR140–150563555.7−23.20.164.5   
Ash-Bearing Siliceous Clay1851149A007H04WR140–150582035.5      
Biogenic Silica- and Ash-Bearing Clay1851149A008H03W140–150664165.5−23.10.174.0   
Ash-Bearing Silica-Rich Clay1851149A009H03W140–150753835.3−23.60.225.7   
Ash-Bearing Siliceous Clay1851149A010H03W140–150852955.3−22.00.258.4   
Diatomaceous Clay1851149A011H03W140–150933654.8−23.10.102.9   
Ashey Clay1851149A012H03W140–1501044005.0−23.40.143.5   
Silt-Bearing Clay1851149A013H04W18–201132814.7−24.50.113.8   
Ash-Bearing Clay1851149A014H02W140–1501212404.7      
Clay (Bioturbated)1851149A015H03W104–1061352813.9−23.20.041.5   
Clay (Bioturbated)1851149A016H03W140–1501422143.7−22.80.094.0   
Clay (Bioturbated)1851149A017H03W110–1121513294.6−24.40.113.3   
Clay (Bioturbated)1851149A018H03W123–1251603404.7−25.50.082.4   
Clay (Bioturbated)1851149A018H03W140–1501613354.9−25.60.072.0   
Silt-Bearing Clay1851149A019X01W58–601652854.5−25.00.082.7   
Silt-Bearing Clay1851149A020X01W140–1501712594.9−25.00.083.1   
Zeolite-Rich Clay1851149A021X01W40–42180974.6−28.10.1616.7   
Clay1851149B003R04W140–1501762364.6−24.60.094.0   
Radiolarian Chert1851149B011R01W19–22237262.5      
Clay-, Ash, and Radiolarian-Bearing Nanofossil Marl1851149B016R01W93–98283     152.729.7
Calcareous Radiolarian Marlstone1851149B020R01W25–35321     362.429.5
Nannofossil-Bearing Marlstone1851149B022R01W20–25340     122.828.5
Radiolarian Chert1851149B022R01W106–110341183.0      
Clay-Bearing Nannofossil Chalk1851149B027R01W49–55388234.0−26.40.026.8901.829.4
Radiolarian-Bearing nannofossil Marl1851149B028R02W48–56399     742.229.6
Nannofossil Marl1851149B029R01W28–35407     751.729.3
 
Leg 129, Sites 800, 801, and 802
Pelagic Brown Clay129800A05R214–2032  −25.90.04    
Volcaniclastic Sandstone129800A26R2110231  −26.20.54    
Volcaniclastic Siltstone129800A28R180248661−0.9      
Volcaniclastic Siltstone129800A30R111–1627214       
Sandstone Breccia129800A41R11153645       
Clay129800A53R11546537−0.2−22.40.0411.7   
Pelagic clay with Zeolites129801A03R02W120–122232895.0−24.70.062.1   
Pelagic Clay with Nannofossil Ooze129801A05R03W120–12244        
Volcaniclastic Sandstone/Claystone129801A16R01W62–6414632 −26.70.0721.4   
Volcaniclastic Sandstone/Claystone129801B01R01W142–14619513 −26.00.038.4   
Volcaniclastic Claystone/Turbidite129801B05R02W10–12231        
Volcaniclastic Turbidite129801B06R03W136–14024314    9.31.827.8
Volcaniclastic Turbidite129801B08R03W125–12726222       
Red Radiolarite129801B33R01W133–135444783.9 0.034.2   
Pelagic Claystone129802ODP38R3353421274.5      
Silty Claystone129802ODP47R21104221164.2      

[9] Reduced C concentrations range from 0.02 to 0.5 wt.% (for 32 samples), with δ13C ranging from −28.1 to −21.7‰ (Table 1, Figure 3b), and the concentrations show some decrease with depth in the ash and clay layers at Site 1149 (Figures 5a and 5b). The decrease in reduced C concentration with depth in Unit I at Site 1149 (Figure 5a) appears to be correlated with the slight decreases in N concentration (see Figures 3a, 3b, and 4a), and Creduced/N shows subtle decrease over the upper 180 m of section (Figure 5c). Calcite from carbonate-rich layers (12–90 wt.%) of 1149B (all from >280 mbsf) has δ13C of +1.7 to +2.8‰ and δ18O of +28.5 to +29.7‰ (see Table 1).

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Figure 5. Carbon concentration and isotope data for the upper 180 m of the sediment section at Site 1149. (a) Sediment C concentration for Site 1149. Note slight trend toward lower C content as a function of depth. (b) Carbon isotopic composition of samples recovered from Site 1149. Note the shift in δ13C in the upper ∼100 mbsf, mostly within the ash- and biogenic-silica-bearing clay part of the section. Below 180 m depths, lithologies are radiolarian cherts and nannofossil chalk/marl. (c) Creduced/N ratios (by weight) as a function of depth (mbsf) in the upper 180 m of the Site 1149 sediment section. Colored lines indicate two possible weak down-section trends, with either all of the variation occurring in the upper 120 m (the two separate blue lines), or a more continuous trend toward lower Creduced/N over the depth interval shown in this figure (green line).

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4. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

4.1. Downhole Variations in N and C Concentrations and Isotopic Compositions

[10] Sediments are highly variable in N concentration and isotopic composition and these variations have been used to trace a wide range of processes in organic-rich sediments, including diagenesis [e.g., Sweeney et al., 1978; Peters et al., 1978; Rau et al., 1987; Libes and Deuser, 1988; Altabet and Francois, 1994; Minoura et al., 1997; Muller and Voss, 1999; Ettwein et al., 2001]. Many of these studies were carried out in relatively organic-rich systems (compared with the section at Site 1149), very near the sediment-water interface [e.g., Peters et al., 1978; Minoura et al., 1997] or nearer continents, and relatively little attention has been paid to N behavior in deep-sea sediment sections such as that at Site 1149 (notable exceptions in the literature, but lacking N and C isotope data, are the studies by Muller [1977], Waples and Sloan [1980], and Waples [1985]). Here we will attempt to outline some of the possible reasons for the shifts in concentration and isotopic composition of N and C with depth at Site 1149.

4.1.1. Oceanographic Considerations

[11] The decrease in δ15N with depth at Site 1149 could in part reflect change in δ15N of the primary sediments being deposited at this site through time. An increase of δ15N with time requires an increase in productivity or a decrease in nutrient supply that makes the preferential uptake of 14N by organic matter less efficient [see Ettwein et al., 2001]. Samples from 90 mbsf have δ15N values similar to those at greater depths at Site 1149, and samples at around 50 mbsf are clearly higher in δ15N than those below 90 mbsf. Therefore the shift in δ15N (and less clear trends in N and C content and δ13C) would have begun within that period of deposition (Table 1, Figure 4), corresponding to ∼2.5 to 3.5 Ma [Lozar and Mussa, 2003]. The most likely change in oceanographic factors around this time is the increase in the Kuroshio current (due to the closure of the seaway between North and South America) which is documented in the biostratigraphic record at Site 1149 by three events between 75.11 and 68.46 mbsf [Lozar and Mussa, 2003]. The general motion of the site toward the higher productivity waters of this current could help account for the gradual increases in δ15N, δ13C-org, and N and C concentrations. However, the region known to have been strongly affected by the Kuroshio current is extremely high in productivity, with most samples containing 0.5 to 1.5 wt% C as documented during ODP Leg 186 [see Mora, 2002]. Furthermore, although the waters affected by the Kuroshio current are higher in productivity, which could cause an increase in δ15N, they are also higher in nutrient supply, possibly mitigating the effects of productivity on marine δ15N (see discussions by Farrell et al. [1995], Milder et al. [1999], Pride et al. [1999], and Ettwein et al. [2001]). Finally, it is unclear whether a change in productivity of this magnitude would produce a trend in Creduced/N as seen in the data for Site 1149 (see discussions by Minoura et al. [1997], Milder et al. [1999], and Pride et al. [1999]).

[12] Variations in δ15N similar to those at Site 1149 have been attributed in other studies to varying mixtures of terrestrial and marine organic matter [e.g., Peters et al., 1978; Minoura et al., 1997], with the marine component having δ15N near +8‰, with δ13C near −20.5‰, and the terrestrial component having δ15N near +1.8‰, with δ13C near −26.5‰. The sediments analyzed in these two studies were deposited in settings more proximal to continental sources (for Peters et al. [1978], on the NE Pacific continental shelf; for Minoura et al. [1997], near Japan in the Japan Sea), and we consider the significant additions of terrestrial organic matter to the Site 1149 sediments (and certainly the sediments obtained on Leg 129) less likely. It is uncertain whether the volcanic ash-rich sediments deposited at Site 1149 contained appreciable terrestrial organic matter during their deposition (see discussion of sediment sources at Site 1149 by Urbat and Pletsch [2003]). However, this possible mixing warrants further study, and more detailed biogeochemical work could help elucidate the sources of the organic matter [e.g., Madureira et al., 1997; Meyers and Doose, 1999; Shipboard Scientific Party, 2000]. In Figure 6, we compare the δ15N-δ13C data for the upper part of the Site 1149 section with the “Marine” and “Terrestrial” organic end-members proposed by Minoura et al. [1997] (very similar to the end-members proposed by Peters et al. [1978]). From this comparison, in particular based on the relative scatter of the Site 1149 data about the mixing line for these two end-members, it is apparent that the Site 1149 sediments do not show a simple marine-terrestrial mixing behavior. The Site 1149 data do show some correlated variation, however (see Figure 6; noted above), and when three outliers are removed from consideration (indicated in boxes on Figure 6), the data show a linear relationship (r2 = 0.60) with a slope quite different from that of the mixing line of Minoura et al. [1997]. The data in Figure 6 for Franciscan metasedimentary rocks are discussed in section 4.2. Finally, vascular land plants typically have C/N higher than that of algae [see Meyers and Doose, 1999, and references therein], and a down-section increase in the terrestrial organic component would thus have likely resulted in an increase in C/N rather than the observed decrease in C/N at Site 1149 (see Figure 5c). Furthermore, the Creduced/N values for the Site 1149 sediments (2–9 on atomic basis) fall in the range for algae (4–10 on atomic basis) and not in the range of significantly higher C/N for vascular land plants (≥20 on atomic basis [Meyers, 1994; Meyers and Doose, 1999]).

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Figure 6. Plot of δ15N vs. δ13C of the upper part of the sediment section at Site 1149, compared with the mixing line (thicker, red line) between “Terrestrial” and “Marine” organic components proposed by Minoura et al. [1997; cf. Peters et al., 1978]. When the three outliers indicated in the small boxes (the sediment sample from 1 mbsf, and two bioturbated clays from near 135–142 mbsf) are removed from consideration, an r2 of ∼0.60 is obtained for the Site 1149 data (see thinner, orange line), but the slope of this line differs significantly from that of the mixing line of Minoura et al. [1997]. The data for low-grade Franciscan Complex metasedimentary rocks plot near the “Terrestrial” organic component of Minoura et al. [1997], consistent with the known deposition of the Franciscan sediments near the continental margin in W. North America.

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4.1.2. Diagenetic Processes

[13] A third possibility is that the shift in δ15N values in the sediments recovered from the upper part of the Site 1149 section (Site 1149A), with δ15N decreasing steadily within the uppermost 120 m (Unit I) from +8.2‰ at 1.4 mbsf to +4.7 at 113 mbsf (Figure 4b), is due to complex diagenetic processes, conceivably with superimposed more minor effects of the two oceanographic alternatives presented above (changes in productivity related to ocean currents and/or deposition of differing proportions of marine and terrestrial organic fractions). The down-section change in δ15N at Site 1149 is accompanied by a decrease in total N concentration (Figure 4a), which could reflect the loss of N with δ15N significantly higher than +8‰, perhaps a nitrate component. Reduced-C concentrations and δ13C also decrease slightly with increasing depth in the upper 150 m of site 1149, contributing to a subtle down-section trend in Creduced/N, and indicating some parallel behavior throughout the organic reservoir. Together, the shifts in reduced C, total N, and Creduced/N with increasing depth at Site 1149 are compatible with the shifts in the same parameters over similar depth horizons reported for other Pacific Ocean deep-sea sediment cores and attributed to diagenesis [Muller, 1977; Waples and Sloan, 1980; Waples, 1985] (unfortunately, N and C isotope data were not obtained in these previous studies). A full explanation of these organic geochemical phenomena, and a more comprehensive reconstruction of diagenesis in this section, would require analyses of separated organic and inorganic N fractions (see studies by Muller [1977], Waples and Sloan [1980], and Williams et al. [1995]) and a more comprehensive consideration of the chemical compositions of the interstitial waters from the sediment cores. However, some reasonable speculation would be that the shift in δ15N, and the associated other variations in the C-N signature, are the result of diagenesis, perhaps a microbially mediated reprocessing of organic matter such as that suggested to occur in algal mats [Lehmann et al., 2002].

[14] The results of Lehmann et al. [2002] deserve particular notice in this context for the similarity of results despite experimental conditions that differ significantly from the natural conditions in Site 1149. This study consisted of experimental simulations of early diagenesis of algae under varying redox conditions. These authors found minor shifts in concentration of C and N, during early diagenesis, and fairly significant changes in isotopic composition, including a shift in δ15N of ∼3‰ under anoxic conditions after 60 days. The similarity of the results of our study, showing shifts in δ15N of ∼3‰ in a 120 m sediment column and the Lehmann et al. [2002] study conducted at different pressures and temperatures, and at very different timescales, suggests that a shift in δ15N could be a common feature of early diagenesis in off-trench, deep-sea sediment sections, and that the magnitude of the negative shift (depending on a variety of conditions, notably the redox state) could conceivably be on the order of ∼1–3‰. At Site 1149, the release of organic N, presumably related to the N and δ15N shifts reported here, is recorded by the slight enrichments (up to ∼200 μM) in dissolved NH4+ in the interstitial waters, the overall chemistry of which has led to the conclusion that the Site 1149 diagenetic environment is slightly suboxic [Plank et al., 2000]. Dissolved NH4+ in the waters (see Figure 4c) reaches a maximum through the upper 50 mbsf, reflecting N release during organic matter degradation, and the decreased dissolved NH4+ below this interval reflects enhanced uptake during clay diagenesis [Plank et al., 2000].

4.2. Comparison of N Signature in Deep-Sea Sediments With That in Paleoaccretionary Rocks

[15] Although δ15N is relatively high and highly variable in surface/near-surface oceanic sediments (with some values >+10‰; see Figure 2), the low grade, forearc metasedimentary rocks studied to date appear to have a much narrower range of δ15N values mostly lower and between +1 and +3‰ [Bebout and Fogel, 1992; Bebout, 1997; Bebout et al., 1999a; Sadofsky and Bebout, 2003] (see comparisons in Figure 2). This “disconnect” between the δ15N values of many oceanic sediments (see Figures 234), mostly from relatively near the sediment-water interface, and the values for low-grade metamorphosed deep-sea sediments in forearc metamorphic suites (representative data in Figure 2) could result from several factors. First, many of the deep-sea sediment sections analyzed previously and in the present study, are not representative of the near-continent, trench settings with high sedimentation rates resulting in the voluminous trench sediment and accretionary prisms sampled by the forearc metamorphic rocks (e.g., Catalina Schist, Franciscan Complex, Western Baja Terrane [see Sadofsky and Bebout, 2003]). Enhanced contributions of terrestrial organic matter (with lower δ15N and δ13C than marine organic matter) must be considered a factor when comparing the seafloor sediment data with the N and C isotope data for the circum-Pacific forearc metamorphic rocks as the metamorphosed voluminous clastic sediments in these metamorphic suites were deposited near a continental margin. Correspondingly, the δ15N values or these low-grade forearc metamorphic rocks generally fall in the range of +1 to +3‰, and with low δ13C, quite near the “Terrestrial” organic component proposed by Peters et al. [1978] and Minoura et al. [1997] (see Figure 6). As another alternative, diagenesis (more pronounced with increasing depth), and conceivably also, incipient low-grade metamorphism during very early subduction history (perhaps in part accompanying mechanical fluid expulsion at shallow levels of accretionary prisms), shifts the δ15N of the initially high-δ15N seafloor sediments from their near-surface values (δ15N near +8 per mil [see Libes and Deuser, 1988; this study] (Figure 4) toward values similar to those observed for the low-grade metamorphic equivalents, which represent subduction to depths of ∼5–40 km. We propose (see section 4.1.2) that the variation in δ15N and δ13C (Figure 6) and C and N concentrations and Creduced/N) in the Site 1149 sediment section largely reflects diagenesis, possibly with some lesser superimposed variation due to changes in primary productivity and marine-terrestrial organic sources. However, it appears extremely likely that the lower whole rock δ15N (and the δ13C of the reduced C) of the forearc metasedimentary rocks, relative to that for many marine sediments and most of the sediment section at Site 1149 (see comparison in Figure 2), is largely attributable to greater proportions of terrestrial organic matter in the forearc metamorphic suites deposited nearer a continental sediment source (perhaps with superimposed minor effects of diagenesis).

4.3. Mass Balance Calculations and Implications for Fluxes Into Subduction Zones

[16] Because there are significant variations in the concentrations and isotopic compositions of N and C throughout these sedimentary sections, the calculation of the fluxes of these elements into the IBM subduction zone is not straightforward. The approach that we favor is to use shipboard estimates of the lithologic units, including the thickness of these units (and average densities of the materials), and average all measured values for each lithological layer (see Table 2). We present only data from Site 1149 in these calculations. We will assume that a section like that at Site 1149 is subducted and not take the apparent shift in shift in δ15N as a function of depth into account at this time, because of the difficulty of deciding the cut-off for which samples to include and which to leave out. We suspect that the actual δ15N and N concentration could be slightly lower than the values used in the calculations we present below, but that the magnitudes of the differences produced by these small changes are at this point within error in any likely application of these data (<±0.5‰ for δ15N, <5% relative for the mass of N subducted).

Table 2. Estimated Annual Sedimentary N-C Subduction Flux at the Izu Margin
 Thickness, mN, ppmδ15NAirCreduced, wt.%δ13CVPDBAverage Densitya
  • a

    The average density, used in these calculations, is equal to bulk density minus the pore water content of a given depth interval. This takes into account the actual density of grains and the porosity. High porosity in the upper part of the section leads to low density of the actual sediment (solids).

  • b

    Hilton et al. [2002] use values of 0‰ and −20‰ for “carbonate” and “sedimentary/organic”, respectively (also see discussion by Van Soest et al. [1998]).

Average
Ash and diatom/rad clay (1149 Unit I)1203715.80.22−22.80.641
Dark brown pelagic clay (1149 Unit II)602614.50.09−250.695
Chert and marl layers (1149 Units III, IV, and V)230223.20.02−24.82.04
  N, gδ15NAirC, gδ13CVPDB 
Fluxes (per Linear km)
Fluxes (for C only reduced C) 2.5E + 065.01.4E + 07−24.0 
Calcite C flux, marl layers   9.2E + 082.3 
Total C flux (full section)   9.3E + 081.9 
 
Total Fluxes (Over 1050 km Linear Trench Length)
This study 2.6E + 095.09.8E + 111.9 
Hilton et al. [2002] 5.8E + 097.0 subarc6.2E + 110, −20b 

[17] To calculate the fluxes of C and N entering the subduction zone at the Izu trench, we begin by separating the materials into three general categories chosen to reflect the resolution of the sampling in this particular study. These categories are: Ash and diatomaceous clay (∼120 m thick, Unit I at site 1149); Dark brown pelagic clay (∼60 m, Unit II, 1149); Radiolarian chert, zeolite-rich clay and marl layers (∼230 m). Multiplying the thicknesses of these units by the orthogonal convergence rate of 5 cm/yr (see compilation of convergence rates by Plank and Langmuir [1998]), and the average dry densities for each unit (values from Plank et al. [2000]) allows a simple calculation of the mass of each lithology entering the subduction zone. Note that the volumes are multiplied by dry density to remove porosity (quite high in the upper 180 m) and pore water from the discussion, we assume for simplicity that all pore water is lost during the early stages of subduction. Average concentrations of C and N can then be combined with our knowledge of the mass of that lithology being subducted to determine the masses of C and N entering the subduction zone in solids. Simple averages of the δ13C and δ15N values in each lithology are used to calculate an average isotopic composition for each unit,. The mass of carbonate was calculated by averaging the shipboard analyses of carbonate content for the mixed carbonate/siliceous layers, combined with the isotopic compositions presented in Table 1.

[18] On the basis of these mass balance calculations, carried out using the orthogonal convergence rate of 5 cm/yr [Plank et al., 2000] (Table 2), we suggest that a sedimentary section like that at Site 1149 delivers an annual flux into the subduction zone of 2.5 × 106 g of N and 1.4 × 107 g of reduced C per linear km of trench, with average δ15N of +5.0‰ and δ13C of −24.0‰, respectively. In addition, 9.2 × 108 g/yr of C is subducted per linear km in carbonate-rich layers of 1149B with average δ13C of +2.3‰. Because of the relatively low C content of the clastic units and the fairly thick carbonate section, the carbonate (oxidized C) budget overwhelms the reduced C budget at this site and 9.3 × 108 g/yr per linear km are input into the subduction zone with an average δ13C of +1.9‰ near that of the carbonate-rich horizons. The overall Ctotal/N of the subducting sediment is near 370:1 (including both reduced and oxidized C reservoirs), the Coxidized/Creduced is ∼65:1, and the Creduced/N (organic component) is ∼5.6:1. The total N flux into the Izu-Bonin convergent margin, for a trench length of 1050 km [from Plank and Langmuir, 1998], is calculated at 2.62 × 109 g/yr, considerably smaller than the flux of 5.8 × 109 g/yr calculated by Hilton et al. [2002], who estimated concentrations for subducting sediments (using 100 ppm for the entire sediment section of 400 m) and used a somewhat larger sediment flux into this margin (4.18 × 1013 g/yr compared with our use of 3.09 × 1013 g/yr).

[19] A more complete inventory of the incoming N and C budgets in the IBM convergent margin will require analyses of AOC from Sites 801 and 1149. δ15N for MORB glass from the East Pacific Rise ranges from −1.0‰ to −6.5‰ (mean of ∼−3.0‰, n = 7 [Marty and Humbert, 1997]), with N2 concentrations of 3–65 × 10−10 moles/gram (<0.5 ppm) and concentrations of NH4+ in spilites (perhaps analogues to some seafloor basalts) from SW England are as high as 182 ppm (range of 1–182 ppm, with mean N concentration of 53 ppm [Hall, 1989, 1990]). These concentrations may represent the extremes for fresh seafloor basalt (with the baseline of 0.01–1.0 ppm N) and the most highly altered oceanic basalt (with up to ∼200 ppm N), and the alteration in AOC at Sites 801c and 1149 is likely to be intermediate in extent. The high K2O values reported for the crustal sections from ODP Site 801c (which experienced an increase of K2O by 17% due to alteration) and 1149 (which is more heavily altered than 801c) provide an indication that AOC may contain appreciable N that could figure significantly in the subduction flux models. This may be especially true when one considers the large volume/mass of oceanic crust being subducted (∼6 km section); any elevation in N concentration (due to hydrothermal alteration) above the maximum ∼0.5 ppm levels observed in fresh basalt glass [Marty and Humbert, 1997] could result in the subduction of N in AOC rivaling that subducted in thin, shale-poor (e.g., carbonate or chert rich) sediment sections (see Bebout [1995, Table 2] and discussions below). However, these elevated N concentrations are likely to occur only within the uppermost, more hydrothermally altered 1 km of the subducting oceanic crust (see discussion below).

[20] Our study demonstrates the importance of evaluating subduction inputs through analyses of the materials outboard of individual trenches and thought to be subducting in any attempts to chemically mass balance individual convergent margins. As in the case of the considerations of pre-subduction sediment δ15N, which can vary with a number of sedimentological and diagenetic factors (discussed above), the use of a single sediment N concentration as applicable in considerations of inputs and outputs at multiple convergent margins from quite different tectonic and sedimentological environments could easily yield flawed results. Our C and N input flux estimates, based on analyses of the Site 1149 section, differ significantly from recent estimates of C-N inputs based on generalized assumptions regarding C and N concentration made without the benefit of data for the sediment at this margin (latter estimates by Hilton et al. [2002]; see comparisons in Table 2). Attempts to deduce overall efficiencies of N and C return in arc volcanic gases are highly susceptible to uncertainties in the input (and output) estimates, as can be demonstrated for the Central American arc-trench system for which Fischer et al. [2002] recently claimed extremely efficient volcanic arc return of sedimentary N. Using N concentration data for only the upper 160 m of the section at Site 1039, Li et al. [2003] obtained a sedimentary N subduction rate of 9.9 × 109 g/yr (for the 1100 km of trench length). Use of even a quite low concentration for the lower, carbonate-rich section would significantly increase this flux estimate (work on the lower section is in progress). In comparison, Fischer et al. [2002] employed a smaller input rate of 2.4 × 109 g/yr, using an averaged N concentration of 100 ppm for the upper 175 m of the section (and assuming no N in the lower section) in their comparison with arc N outputs in Central America. Fischer et al. [2002] claimed similarity of their input of 2.4 × 109 g/yr with their estimate of arc output of 4.1 × 109 g/yr, and speculated that this similarity reflects extremely efficient return of subducted sedimentary N.

[21] It is worth noting that the δ15N of +7‰ used by Fischer et al. [2002] and Snyder et al. [2003] for sediment contributing to arc magmatism is higher than those of both the aggregate sediment sections at Sites 1149 (IB) and 1039 (Central America; bulk δ15N of +5.0‰ and +5.6‰, respectively for the two margins [see Li et al., 2003]) and the shallowly subducted paleoaccretionary rocks (+1 to +3‰; see Figure 2), a difference that could be related to increase in δ15N during metamorphic N losses (see examples of this relationship by Haendel et al. [1986], Bebout and Fogel [1992], Bebout et al. [1999b], and Mingram and Brauer [2001]). These metamorphic N losses would affect the efficiency with which the seafloor sediment N inventory is conveyed to depths beneath arcs, and thus any mass balance employing seafloor sediment as input and arc gases as output. Regarding the input fluxes in subduction zones, it worth noting that all of the estimates of sediment geochemical inputs assume uniformity in the incoming sediment sections along-strike in active trenches, known not to be the case.

[22] We suggest that the apparent similarity in sediment N input with arc volcanic outputs reported by Fischer et al. [2002] for Central America could be partly coincidental, reflecting significant contributions from both sediments and devolatilizating AOC or significant errors in input or output estimates. Given the uncertainty of this N flux in AOC (see consideration by Bebout [1995]), the overall Central American mass balance of N inputs and arc outputs cannot be addressed in entirety; we estimate a N subduction rate of 1.7 × 1010 g/yr in AOC (four times the sedimentary flux for the same margin), using an estimated N concentration of 10 ppm for the crustal lithology. Use of the total sediment + AOC input (2.7 × 1010 g/yr), based on this 1.7 × 1010 g/yr flux in AOC and the sediment N input flux of 9.9 × 109 g/yr (see above), a ∼15% arc return of subducted N would be indicated (incorporating the arc output rate of 4.1 × 109 g/yr from Fischer et al. [2002]). If the AOC N concentration is reduced to 5 ppm, a ∼22% return in arcs would be indicated, and an ∼36% N return in arcs is indicated if a concentration of 1 ppm N is assumed for the AOC. This crude set of comparisons highlights the dire need for any constraints on the total flux and isotopic composition (for both N and the noble gases) of the AOC volatiles component. Direct use of our information regarding N inputs into the Izu-Bonin convergent margin to mass balance with N outputs in the corresponding arc awaits investigation of arc volcanic gases in this margin.

4.4. Fate of Subducted Sedimentary Nitrogen in the Izu-Bonin Margin (and Other Convergent Margins)

[23] Studies of low-grade metasedimentary suites in paleoaccretionary complexes in Western North America shed some light on the degrees of deep subduction of N in sediment and the stable isotope compositions of this N. Recent study of devolatilization in sedimentary lithologies subducted to 5–40 km depths in the Catalina Schist, California, the Franciscan Complex, California (Coast Ranges), and the Western Baja Terrane, Mexico [Sadofsky and Bebout, 2003] (see estimated peak P-T ranges in Figure 7), affords an assessment of the entrainment of N into forearc regions of a relatively “cool” subduction zone. Interestingly, samples from the Coastal Belt, the lowest-grade unit in the Coast Ranges subducted to only ∼5 km depths, show correlated δ15N and N concentration (see Figure 8a), conceivably reflecting differential loss of “heavy” N (as NO3?) during diagenesis and extremely low-grade metamorphism accompanying subduction to these extremely shallow levels [see Sadofsky and Bebout, 2003]. Moving upgrade in the Coast Ranges, the more uniform δ15N of the somewhat higher-grade (more deeply subducted; see Figure 7) Central and Eastern Belt metasedimentary rocks (near +1.5‰) could reflect the more complete loss of this “heavy” component (see data for the three Coast Ranges units in Figure 8a).

image

Figure 7. Pressure-temperature diagram showing estimates of peak metamorphism for low-grade paleoaccretionary prism rocks from California, USA, and Mexico (see Grove and Bebout [1995] for details regarding the generalized phase equilibria and stability fields compiled on this diagram). Arrows are schematic prograde P-T paths rocks might take in subduction zones, reflecting a wide range in thermal structure; for the Catalina Schist, California, these paths are inferred to reflect rapid cooling in an incipiently formed subduction zone, with warmer conditions first producing the highest-temperature epidote-amphibolite-facies unit (labeled as “EA”) and latest-stage, cool subduction producing the lawsonite-blueschist and lawsonite-albite facies units (labeled in this figure as “LBS” and “LA”, respectively). Various units of the Coast Ranges Franciscan Complex are indicated (“Coastal Belt”, “Central Belt”, and “Eastern Belt”, in order of increasing peak metamorphic pressures thus depths of underthrusting [see Sadofsky and Bebout, 2003; Blake et al., 1987]). Also indicated is the peak P-T for the Franciscan Complex metagreywackes at Pacheco Pass, California (estimates from Ernst [1993]). Fields labeled “ST1”, “ST2”, and “ST3” indicate peak conditions for tectonometamorphic units of the Western Baja Terrane [from Sedlock, 1988], with “ST3” representing the lower-P conditions in this suite.

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image

Figure 8. Nitrogen concentration and isotope compositions and Creduced/N of subduction-zone metasedimentary rocks, and comparison with data for the Site 1149 sediments. (a) δ15N versus whole-sediment N concentration, illustrating differences among the tectonometamorphic units of the Franciscan Complex and Western Baja Terrane, Mexico [from Sadofsky and Bebout, 2003]. (b) Creduced/N versus N concentration for metasedimentary rocks from the same units, demonstrating the extremely uniform Creduced/N for the lowest grade rocks in the Coastal Belt, for which peak metamorphic temperatures are <200°C at approximately 5 km maximum depths of underthrusting. δ15N values are similar to those expected in sediments with organic matter derived primarily from photosynthesizing organisms [see Rau et al., 1987]. Metamorphic devolatilization of N would be expected to produce a trend of increasing δ15N with decreasing N content as N bearing fluids (likely with N2 as the dominant N fluid species) preferentially fractionate the lighter isotope [see Bebout et al., 1999a, 1999b]. Also indicated in this figure is the range of Creduced/N and N concentrations in the Site 1149 sediment section (data for the upper 180 m only).

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[24] Carbonreduced/nitrogen ratios are another useful indicator of the devolatilization of these elements, as both of these elements are provided to the sedimentary rocks largely by organic processes. Most deep-ocean sediments should have Creduced/N ratios <20:1 [Muller, 1977; Sweeney et al., 1978; Waples and Sloan, 1980], and the vast majority of the high-P/T metasedimentary suites fall into that range (see Figure 8b). The Coastal Belt samples have Creduced/N more uniform than that of the higher-grade units (see the line for Coastal Belt samples on Figure 8b), fully within the range for modern seafloor sediments, and perhaps indicating that the loss of isotopically “heavy” N producing the decrease in whole rock δ15N did not result in significant shifts in the Creduced/N of the same rocks (i.e., the δ15N of the N that was lost was significantly higher than that left behind in the sediments). The higher-grade units in the Coast Ranges show scatter in Creduced/N and shifts mostly to higher Creduced/N relative to Creduced/N of the lowest-grade Central Belt, perhaps reflecting the effects of deeper metamorphic devolatilization (i.e., at depths of 10–40 km).

[25] Yet unknown is the extent of N loss, and accompanying shift in δ15N, that occur during subduction of sedimentary lithologies to depths greater than those represented by the circum-Pacific paleoaccretionary suites we've examined (i.e., at depths >40 km). Busigny et al. [2003] reported whole rock and mica separate δ15N values of +2.6 to +4.8‰ for metasedimentary samples of the Schistes Lustres (and metasedimentary rocks at Lago di Cignana, both exposed in NW Italy) subducted to depths corresponding to pressures of 1.5 to 2.5 GPa (approximately 60–90 km). Although the interpretation of the data for the Schistes Lustres and at Lago di Cignana is complicated by the variable, in some cases extreme, overprinting that occurred during exhumation of these rocks [Reinecke, 1998; Agard et al., 2002; Bebout et al., 2003], these somewhat higher values (relative to those for the lower-P suites studied by Bebout and Fogel [1992] and Sadofsky and Bebout [2003]) could conceivably reflect some increase due to small amounts of N loss during subduction to these greater depths. However, Busigny et al. [2003] also analyzed presumed non-metamorphic equivalents, and because the δ15N values of these sediments are similar to those of the metamorphosed rocks, argued for no change in δ15N in sediments subducted to depths approaching 90 km. Confirmation of whether these data for the Schistes Lustres do reflect deep subduction, without appreciable geochemical effects of exhumation-related overprinting, awaits a more detailed investigation of mineral chemistry (and single-grain δ15N and trace element compositions) in these rocks. As briefly discussed above, recent studies of arc volcanic gases have employed δ15N values of near +7‰ for the deeply subducted sedimentary N component [see Fischer et al., 2002; Snyder et al., 2003], and calculations of N isotope shifts due to metamorphic devolatilization at temperatures of less than 300°C (by Rayleigh or batch loss) can easily produce shifts of 3–4‰ (i.e., shifts in δ15N from +3 to +7‰, perhaps with loss of <25% of the initially subducted N (see calculations by Bebout and Fogel [1992] and Bebout et al. [1999a, 1999b]; at T < 300°C, 103lnαNH4+−N2 is >4‰, based on the calculations of Hanschmann [1981]). These relatively small amounts of loss could be difficult to identify in metamorphosed sediments, given the large degree of variability thought to represent variation in the isotopic composition of the seafloor sediment protoliths.

[26] It is appropriate to briefly discuss the extent to which our data, and recent work on subduction-zone metamorphic suites, can elucidate N cycling at the Izu-Bonin margin. For this margin (at 32°N), Peacock [2003] calculated temperatures at the slab-mantle interface of ∼245°C at 50 km depths and ∼540°C at sub-arc depths (∼120 km). Thus, for reference, only the lowest-T paths shown in Figure 7 (resulting in temperatures of just over 200°C at depths approaching 40 km) mimic the prograde paths thought to be experienced in this relatively “cool” convergent margin. Even the P-T path experienced by the somewhat higher-grade lawsonite-blueschist-facies unit of the Catalina Schist (patterned field labeled as “LBS” on Figure 7) is somewhat higher-T than would be expected in this modern subduction zone. On the basis of the calculated P-T paths of Peacock [2003], nearly all of the Cottian Alps Schistes Lustres (peak conditions, 300–625°C, 50–70 km [Agard et al., 2002]) and the rocks at Lago di Cignana (peak conditions, ∼625°C, 90 km [see Bebout and Nakamura, 2003]), appear to have experienced peak metamorphic temperatures somewhat higher than those indicated for their respective maximum depths in the modern Izu-Bonin margin. However, the tectonometamorphic units in the Cottian Alps Schistes Lustres traverse studied by Agard et al. [2002], Busigny et al. [2003], and Bebout et al. [2003], combined with the Lago di Cignana rocks [studied by Bebout and Nakamura [2003] and Busigny et al. [2003]), certainly do cover a range of peak P-T broadly compatible with the P-T calculated for the slab-mantle interface in modern subduction zones (see the P-T field for only the lowest-grade unit of the Cottian Alps Schistes Lustres on Figure 7).

[27] In the Catalina Schist, even in the lawsonite-blueschist-facies metasedimentary unit (labeled “LBS” in Figures 2 and 7; peak conditions of 350°C at pressures corresponding to ∼40 km), δ15N values appear slightly higher than those in lawsonite-albite-facies equivalents experiencing “cooler” prograde P-T paths and lower peak metamorphic temperatures of <250°C (see schematic prograde P-T paths in Figure 7). The Schistes Lustres rocks, metamorphosed at temperatures >300°C but at higher pressures, are similarly somewhat higher in δ15N than the units of the Franciscan Complex and Western Baja Terrane, the latter for which peak temperatures are roughly consistent with peak recrystallization at shallower levels of 5–40 km in the modern Izu-Bonin subduction zone (again, 245°C at 50 km calculated by Peacock [2003]). However, Busigny et al. [2003] suggest that the protoliths for the Schistes Lustres had δ15N higher than that of the protoliths for the metasedimentary rocks of the Catalina Schist, Franciscan Complex, and Western Baja Terrane. Together, the work on the circum-Pacific suites [Bebout and Fogel, 1992; Sadofsky and Bebout, 2003] and the work on the Schistes Lustres [Busigny et al., 2003] appear to demonstrate the impressive compatibility of ammonium ions in micas (see discussion by Boyd [2001]), suggesting that efficient deep entrainment of N into convergent margins is facilitated by the extremely broad stability range of particularly the white micas (see experimental study of the stability relations of phengitic muscovite by Domanik and Holloway [2000]).

[28] Overall, given the likelihood that even small amounts of N loss can potentially result in significant shifts in sediment δ15N (discussion above), and taking into account the systematics in circum-Pacific paleoaccretionary suites (Catalina Schist, Franciscan Complex, Western Baja Terrane; +1 to +3‰ in sediments subducted to 5–40 km) and the results presented by Busigny et al. [2003] (somewhat higher values of +2.6 to +4.8‰ in sediments subducted to 50–90 km), it appears that the use of a mean δ15N value of +7‰ (perhaps ±2‰) is reasonable in the studies of volcanic gases (see recent studies by Fischer et al. [2002], Hilton et al. [2002], and Snyder et al. [2003]). However, as highlighted above, knowledge of the N concentrations and isotopic compositions of subducting AOC is critical in any further comparisons of the seafloor, metamorphic, and volcanic gas records of convergent margin N cycling; also critical is the efficiency with which N and noble gases released from devolatilizing AOC (and sediment, for that matter) can be mobilized into the subarc mantle wedge. Integration of the N results with other data (trace element, isotopic) for individual arcs indicating relative slab contributions from AOC and sediment could potentially allow further delineation of AOC and sediment N (and C) sources.

5. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[29] In this paper, we provide an inventory of the C and N concentrations and stable isotope compositions in the sediment sections obtained on ODP Legs 125 and 185, focusing primarily on Site 1149 drilled on Leg 185. The upper 120 m of the Site 1149 section shows correlated variation in the C and N concentrations and isotope compositions that we suggest are primarily related to diagenesis, but with possible superimposed lesser effects of changing productivity and relative contributions of terrigeneous and marine sediment sources over the period of its deposition. More detailed biogeochemical study of the organic matter in this section would provide a more thorough assessment of organic matter sources in the section. With the data we obtained for the Site 1149 section, we estimate the C and N input fluxes (via subduction) into the Izu-Bonin convergent margin to be used in assessments of the input/output mass balance of these elements at this margin. We demonstrate the importance of using geochemical inputs constrained by drilling and geochemical analyses of the sedimentary section outboard of the margin being considered, rather than using estimates based on analyses of sediments from other tectonic and sedimentological settings, or “global” estimates of the concentrations and isotope compositions of subducting sedimentary lithologies. Further work on sediment sections outboard of the other major circum-Pacific (and other) modern subduction zones is warranted. It is worth noting that the majority of modern Earth subduction is circum-Pacific [see Plank and Langmuir, 1998; Hilton et al., 2002] and the C and N budgets in other ocean basins do not contribute as significantly to the modern convergent margin flux of these elements. On the basis of the existing observations regarding the effects of mechanical compaction, deformation, and diagenetic and metamorphic devolatilization in forearcs (from studies of forearc metamorphic complexes), we suggest that a significant fraction of the initially subducted N and C (reduced and oxidized) could be retained to great depths in subducting sediment sections to either return to the surface in arcs or enter the deeper mantle beyond sub-arc regions.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information

[30] We dedicate this paper to Stuart Boyd, who recently passed away quite prematurely. We acknowledge Stuart's extremely creative contributions, including his ideas directed toward the understanding of (bio)geochemical N cycling that helped pave the way for more extensive use of the N isotope system in the growing field of geobiology. Funding for our study was provided by the Joint Oceanographic Institutions, United States Science Support Program (JOI-USSSP), and for the work on metamorphic suites, from the National Science Foundation (EAR-9805050). SJS acknowledges the DFG (SFB 574) for support during the completion of this study. We thank Leg 185 co-chief scientists Terry Plank and John Ludden for their support of our shorebased research. We would like to thank Jeff Alt for providing archived samples from Ocean Drilling Program Leg 129, David Velinsky for helpful discussions regarding sample preparation and sediment diagenesis, and Philippe Agard for discussions of the tectonometamorphic and geochemical evolution of the Schistes Lustres exposed in the Cottian Alps, France and Italy. Finally, thanks to Terry Plank and Bill White for their editorial handling, and to Dave Hilton and an anonymous reviewer for their helpful reviews of our manuscript.

References

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  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information
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Supporting Information

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Analytical Methods
  5. 3. Results
  6. 4. Discussion
  7. 5. Conclusions
  8. Acknowledgments
  9. References
  10. Supporting Information
FilenameFormatSizeDescription
ggge403-sup-0001-tab01.txtplain text document4KTab-delimited Table 1
ggge403-sup-0002-tab02.txtplain text document1KTab-delimited Table 2

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