LIMNOLOGY and
OCEANOGRAPHY
Limnol. Oceanogr. 00, 2018, 00–00 C 2018 Association for the Sciences of Limnology and Oceanography V
doi: 10.1002/lno.10778
Phosphorus recycling in deeply oxygenated sediments in Lake Superior controlled by organic matter mineralization Jiying Li 1 2
,1,a* Yishu Zhang,1 Sergei Katsev1,2
Large Lakes Observatory, University of Minnesota Duluth, Duluth, Minnesota Department of Physics and Astronomy, University of Minnesota Duluth, Duluth, Minnesota
Abstract Remobilization of phosphorus from aquatic sediments has been extensively investigated in systems prone to anoxia, while studies in well-oxygenated systems have been rare. The recycling efficiency of P in the offshore sediments in the Great Lakes, in particular, is still poorly known. We investigated phosphorus cycling at 13 locations (26–318 m water depth) in oligotrophic Lake Superior where oxygen penetrates into sediments by 2–12 cm. Vertical distributions of iron and phosphorus were measured in porewater and solid fractions, and transformation rates and vertical fluxes were calculated. Whereas a significant fraction of P is bound to ferric Fe in surface sediments, P effluxes into the water column (2.5–7.0 lmol m22 d21) are only weakly affected by iron reduction, because Fe : P ratios in surface sediment are high ( 40–80 mol : mol), and P sorption capacity is far from its limit. In contrast to organic rich systems where P effluxes are sensitive to redox conditions, phosphate effluxes in organic-poor well-oxygenated Lake Superior are controlled by the rates of organic phosphorus mineralization, similar to marine sediments. The efficiency of P recycling in Lake Superior sediments, however, is substantially lower than in marine sediments due to different P biogeochemistry. Only 12% of deposited P is returned to the water column. While burial into sediments is the dominant sink for P in the lake, sediments still contribute up to 40% of total water column P inputs. Similar behavior should be expected in other well-oxygenated freshwater systems, such as other large oligotrophic lakes.
sedimentation fluxes of reactive organic matter and iron, and water column sulfate levels (Katsev et al. 2006a; Hupfer and Lewandowski 2008). On seasonal time scales, the classical model of “oxygen-controlled P mobilization” (Mortimer 1942) links the biogeochemical cycles of Fe and P and predicts P releases under anoxic conditions, whereas under oxic conditions adsorption to sedimentary iron oxyhydroxides (FeOOH) prevents P mobilization. On longer time scales, the fraction of deposited P that becomes returned to the water column has been predicted to be controlled by P immobilization and burial in the anoxic sediment layers (Katsev et al. 2006a; Katsev and Dittrich 2013) While these paradigms focus on processes occurring near the sediment oxic-anoxic boundary and below, P fluxes from deeply oxygenated sediments have received substantially less attention. The efficiency of P immobilization in sediments of oligotrophic iron-rich Great Lakes has not been quantitatively investigated. Lakes Superior, Huron, and Michigan are strongly Plimited, and P fluxes between sediments and water column have been shown to strongly affect the TP dynamics in their water columns (Katsev 2017). While sediments in these “freshwater seas” have been found similar to marine coastal or
Phosphorus (P) is an essential nutrient that limits primary productivity in most freshwater lakes (Hecky and Kilham 1988). While phosphorus enters lakes through inputs from watershed and atmosphere, a significant portion of the annual P budget in water columns of lakes is comprised of P recycled from bottom sediments. This internal loading affects the water column inventory of total phosphorus (TP) and determines the time scales on which TP levels respond to external loading (Katsev 2017). The efficiency of P immobilization in sediments is thus an important parameter that defines the lakes’ ecological trajectories, and changes in sediment geochemical conditions can have profound influences on the lake trophic status and food webs (Sondergaard et al. 2001). The recycling efficiency of P depends on environmental variables such as bottom-water concentrations of oxygen,
*Correspondence:
[email protected] a
Present address: Department of Physical and Environmental Sciences, University of Toronto Scarborough, Toronto, Ontario, Canada Additional Supporting Information may be found in the online version of this article.
1
Li et al.
P cycling in deeply oxygenated sediments
NB
49 BB
48.5
NIP
Sta.2
Latitude (oN)
TB CM
48 IR
EM
47.5 FWM
47 46.5
ED
WM
KW
SW Duluth
46 92
91
90
89 88 87 o Longitude ( W)
86
85
84
Fig. 1. Sampling locations in Lake Superior.
been characterized in detail, therefore studies in similar but unaffected Lake Superior may provide baseline information and insights into the puzzling P dynamics in the lower lakes, where water column trends cannot be explained without accounting for changes in sediment-water P exchanges (Katsev 2017). Here we characterize the diagenetic cycling of P and Fe at multiple locations across the lake and present an updated phosphorus budget that takes into account contributions from sediments.
even hemipelagic marine sediments in terms of their carbon and nitrogen dynamics (Li et al. 2012; Li and Katsev 2014) the dearth of sulfate in fresh water can be expected to strongly affect the efficiency of P recycling. Past studies in other lakes have suggested that P recycling efficiencies in freshwater sediments are generally low compared to marine systems (Caraco et al. 1990). Quantitative comparisons in large systems dominated by autochthonous organic matter, however, have not been performed. This paper investigates the phosphorus fluxes and P-Fe interactions in the well-oxygenated deep sediments in the offshore of Lake Superior, the world’s largest freshwater lake by surface area. Lake Superior is oligotrophic with low concentrations of total phosphorus. TP concentrations in the lake decreased from 0.19 lmol L21 to 0.10 lmol L21 between mid-1970s and 1990s (Sterner et al. 2007) following a reduction in external inputs (Dolan and Chapra 2012), and subsequently stabilized. The present water column TP levels are < 2 lg L21 (Chapra and Dolan 2012), and soluble reactive phosphorus (SRP) is low ( 10 nmol L21; Baehr and McManus 2003). Phosphorus losses via riverine outflow are only 10% of allochthonous inputs (Heinen and McManus 2004), suggesting that sediments must serve as important phosphorus sinks. The strength of the sedimentary P sink, however, is yet poorly constrained, with both P settling rates and benthic recycling efficiency having been estimated only in the Western Arm of Lake Superior (Heinen and McManus 2004). In contrast to Lakes Michigan and Huron, Lake Superior has not been invaded by bottom-dwelling filter-feeding dreissenid mussels, which have greatly reengineered the nutrient flows in the lower lakes and likely caused profound changes to sediment-water fluxes by replacing the natural communities of Diporeia bioturbators and strongly altering the composition of uppermost sediment layers (Higgins and Vander Zanden 2010; Mosley and Bootsma 2015). Preinvasion conditions in the affected lakes have not
Methods Sediment sampling and analyses Sediments and overlying waters were sampled across Lake Superior on multiple cruises aboard the R/V Blue Heron in 2009– 2012 (Fig. 1; sampling dates and locations in Supporting Information Table S1) and processed using the procedures described in Li et al. (2012) and Li and Katsev (2014). Sediment cores were anaerobically sectioned, porewaters were extracted from sections using Rhizon samplers with 0.1 lm membranes, and remaining solid fractions were frozen. Separate cores from the same multicorer set were split lengthwise and scanned for major elements at 200 lm spatial resolution using an ITRAX X-ray fluorescence (XRF) scanning sediment analyzer. Porewater samples for the dissolved Fe(II) and soluble reactive phosphorus (SRP) analyses were acidified with hydrochloric acid (1% of 6 mol L21 HCl) immediately after collection and stored at 48C. Separate wet sediment samples were preserved (frozen) and analyzed for an operationally defined biologically available iron fraction, amorphous Fe(III) oxides, and solid-phase Fe(II) compounds, by a 0.5 N HCl extraction (Supporting Information Table S2; Roden and Wetzel 2002) and phosphate concentrations were determined in the extracts for selected samples. Separate frozen sediment samples were freeze-dried, ground, and homogenized, and the distribution of iron phases in them 2
Li et al.
P cycling in deeply oxygenated sediments
Table 1. Major reactions affecting P cycling in sediments.
was determined using the sequential extraction procedure of Poulton and Canfield (2005; Supporting Information Table S2) to extract exchangeable iron, carbonate iron (siderite and ankerite), and reducible iron oxides (ferrihydrite, lepidicrocite, goethite, hematite, and akaganeite), using magnesium chloride, sodium acetate, and sodium dithionite, respectively. Solid sediments were also analyzed for phosphorus fractions, following the sequential extraction procedure of Ruttenberg (1992; Supporting Information Table S2), which separates the sediment phosphorus into five pools: exchangeable P, ferric Fe-bound P, authigenic carbonate fluorapatite 1 carbonate fluorapatite 1 biogenic apatite 1 calcium carbonate associated P, detrital apatite P of igneous or metamorphic origin, and residual P. Dissolved Fe concentrations in both the extracts and filtered porewater samples were determined by a modified Ferrozine colorimetric method (Viollier et al. 2000) using a Thermo Spectronic GENESYSTM 6 Spectrophotometer: Fe (II) and total Fe were measured before and after reduction with hydroxylamine, and Fe (III) were calculated as the difference between Fe (II) and total Fe. SRP concentrations in porewaters, bottom water samples, and sediment extracts were measured using a modified molybdenum blue method (QuikChemV Method 10–115-01-1Q) on a Lachat Quickchem 8000 flow injection auto-analyzer. Matrix effects were taken into account for calibrations during measurements of phosphate in extracts.
Processes Redox reactions Organic matter
(CH2O)x(NH3)y(H3PO4)z1x H2O ! x CO21
mineralization* Fe reduction†
y NH31x H2O1z H3PO414e–14x H1 Fe(OH)31e– ! Fe211OH–
Fe oxidation‡
Fe2113H2O ! Fe(OH)313H1e–
Mineral precipitation reactions Iron sulfide, pyrite Fe211H2S ! FeS12H1; FeS1H2S!FeS21H2 22 e.g., 10 Ca2112 PO32 4 1CO3
Calcium phosphate Vivianite
§
12F2 ! Ca10(PO4)2(CO3)F2 3 Fe2112H3PO418H2O ! Fe3(PO4)28H2O16H1
Adsorption/desorption Fe21, Fe31 (dissolved) ! Fe(II), Fe(III) (solid) 32 PO32 4 (dissolved) ! PO4 (solid)
Iron adsorption Phosphate adsorption
* The half-reaction may be coupled with reduction of various electron 22 acceptors (O2, NO2 3 , Fe (III), Mn (IV), SO4 , etc.). † The half reaction is typically coupled with oxidation of organic matter. ‡ The half reaction may be coupled with reduction of oxygen and/or nitrate. § Carbonate-fluroapatite is considered the major form of marine sediment authigenic Ca-P precipitation (Ruttenberg 1992), which has a variable of chemical composition, e.g., (Ca, Mg, Na)10(PO4, CO3)6(F, OH)2–3. Other forms of calcium associated phosphorus include minerals fluoroapatite Ca5(PO4)3F, hydroxyapatite Ca5(PO)3OH, and some less stable amorphous phases, e.g., octacalciumphoshate Ca4H(PO4)32.5H2O (Gunnars et al. 2004).
R
Calculations of fluxes and rates Molecular diffusive fluxes (Fi) of dissolved Fe21 and SRP were calculated using the Fick’s law of diffusion. The bulk molecular diffusion coefficients Ds at the in situ temperature 48C and corrected for sediment tortuosity are DFe21 5 123 cm2 yr21, DH2 PO23 5 155 cm2 r21, and DHPO22 5 125 cm2 3 y21 (Boudreau 1997). In determining the fluxes of SRP (mostly HPO22 and H2 PO2 4 4 at the in situ pH), the diffusion coefficients for the two ionic species were averaged, and the calculated fluxes are presented here with uncertainties that account for the possible variations in the composition ratios. Where concentration gradients near the sediment-water interface were poorly resolved, the diffusive fluxes across the interface were calculated from the gradients calculated from the measured porewater concentrations below the interface, the measured bulk bottom water concentrations, and the assumed thickness of boundary layer of 1 cm (based on the 0.05 cm resolution oxygen profiles, reported in Li et al. 2012). The total fluxes of P may be higher than the molecular diffusion fluxes due to contributions from processes such as bioirrigation (Meile et al. 2005; Glud 2008). For oxygen fluxes in these sediments, Li et al. (2012) estimated the difference at 30–50% of the total flux. Bioturbation in Lake Superior is limited to the upper 2 cm of sediment (Li et al. 2012), thus no contributions from benthic fauna are expected below this depth.
The major reactions involved in sediment phosphorus and iron cycling are listed in Table 1. The rates of iron reduction were estimated from the concentration profiles of the dissolved Fe21 and solid-phase Fe(II) using diagenetic equations (see Li 2014). Using Fe (II) profiles benefits from the well-resolved concentration gradients at the Fe reduction/oxidation boundaries (see Results below), while avoiding complications that arise when rate calculation is performed using Fe (III) profiles (see Results). Briefly, below the bioturbation zone (> 2 cm in Lake Superior; Li et al. 2012), under the approximation of nearly steady state conditions in Lake Superior sediments, the diagenesis of iron can be described as dCdiss:FeðIIÞ d 05 uDdiss:FeðIIÞ (1) 1Rdiss:FeðIIÞ dx dx
05
d nUCsolid:FeðIIÞ 1Rsolid:FeðIIÞ dx
(2)
Here Ddiss.Fe(II) is the diffusive coefficient for dissolved Fe(II), Cdiss.Fe(II) and Csolid.Fe(II) are the concentrations of dissolved Fe(II) and solid-phase Fe(II), respectively, and Rdiss.Fe(II) 3
Li et al.
P cycling in deeply oxygenated sediments
and Rsolid.Fe(II) are the iron reduction rates (in lmol cm23 d21) that affect the dissolved and solid-phase concentrations of Fe(II); x is the sediment depth; U is the burial velocity and n is equal to (1–u)q, where u is the porosity and q is the density of dry sediment (2.65 g cm23; Johnson et al. 1982). By integrating Eqs. 1, 2 from the upper boundary of iron reduction (x 5 L) to deep sediment where both the dissolved and solid-phase Fe(II) concentrations no longer change with depth (x 5 1), the depth-integrated rates (in mmol m22 d21) can be written as: dCdiss:FeðIIÞ x51 5 Rdiss:FeðIIÞ 5uDdiss:FeðIIÞ dx x5L (3) dCdiss:FeðIIÞ 2uDdiss:FeðIIÞ 5Fdiss:FeðIIÞ;x5L dx x5L x51 5nUCsolid:FeðIIÞ; x51 Rsolid:FeðIIÞ 5nUCsolid:FeðIIÞ (4) x5L 2nUCsolid:FeðIIÞ; x5L
concentrations of SRP in sediment porewaters relative to bottom waters indicate P effluxes from sediments at all sites. Diffusive effluxes calculated from measured SRP gradients range from 2.5 lmol m22 d21 to 7.0 lmol m22 d21 (Table 2). Scanning XRF profiles of total iron and manganese reveal multiple Fe- and Mn-rich layers (Fig. 3 and more in Supporting Information Fig. S1). These layers in Lake Superior form dense crusts that are visible to the naked eye (see images in Li et al. 2012) and often feel noticeably harder to the touch. Distributions of extractable solid phase iron and phosphorus fractions reveal that the iron-rich layers are dominated by 0.5-HCl extractable or dithionite-extractable iron fractions, which operationally correspond to reactive (reduceable) iron. These Fe enrichments coincide with maxima in iron-bound phosphorus (Fig. 4). Reduced solid-phase Fe(II) accumulates in the sediment below the Fe(III) enrichment. The Fe : P molar ratio is generally high near the sediment surface ( 20–60) with an exception at Sta. IR (Fe : P 5–10), decreases with depth within the oxic layer, and reaches minimum at the depth of iron and phosphorus enrichments (Fig. 4). The vertical distributions of P and Fe(II) mobilization rates in the sediments are shown in Fig. 5. P mobilization rates generally decreased with depth within the surface oxic sediments (Fig. 5; see Fig. 2 for OPD and NPD). P mobilization rates decreased and became negative (P immobilization) near the depth of oxygen penetration (eg., at TB) and nitrate penetration (e.g., SW), coincident with Fe immobilization (negative Fe rates and solid Fe enrichment [Figs. 3, 4]), and increased again in deeper anoxic sediments, coincident with increased Fe mobilization rates (Fig. 5). The depth-integrated rates of iron reduction (R*Fe) averaged 0.047 mmol m22 d21 (Table 3). Phosphorus mobilization in the iron reduction zone (Eq. 5) averaged 0.034 mmol m22 d21 (Table 3).
The total rate of iron reduction is then estimated as R*Fe5 R*diss.Fe(II) 1 R*solid.Fe(II). The porosities and burial velocities measured at our sampling locations are described in Li et al. (2012) and Li (2014). The net rate of phosphorus mobilization was estimated below the bioturbation zone (> 2 cm) similarly to the rate of Fe21 production (Eq. 1): d dCi uDi (5) Ri 52 dx dx Here, Ri is the net rate of phosphate production, Di is the diffusion coefficient of phosphate, and Ci is the concentration of phosphate (SRP). The depth-integrated rates of P mobilization were estimated in anoxic sediments similarly to Eq. 3, from the fluxes of SRP at the upper boundary of the iron reduction zone.
Discussion A conceptual diagram of the obtained profiles is shown in Fig. 6. The profiles illustrate the representative patterns in the concentrations and transformation rates of phosphorus and iron. The processes regulating the corresponding distributions and vertical fluxes are discussed below.
Results The sediments at the sampled locations in Lake Superior are characterized by low sedimentation rates and oxygen penetration depths (OPD) of 2 cm or greater, as described in detail in Li et al. (2012). In sediment layers where oxygen and nitrate concentrations were high, porewaters contained no detectable dissolved Fe(II) (< 1 lmol L21) (Fig. 2). In deeper sediments, dissolved Fe (II) concentrations increased below the depths of oxygen and/or nitrate penetrations (nitrate penetration depth 5 NPD; Nitrate profiles are presented in detail in Li and Katsev [2014]). Porewater SRP concentrations were below the detection limit of 0.5 lmol L21 at the sediment-water interface (SWI), increased to low but detectable concentrations (0–5 lmol L21) within the surface oxic sediments where dissolved Fe(II) was not detectable (Fig. 2), and increased significantly into the deep anoxic sediments in parallel with dissolved Fe(II), typically reaching 40–90 lmol L21. The higher
Iron and manganese cycling The geochemical cycling of phosphorus is tightly linked to the redox cycling of iron (e.g., Katsev 2016). P strongly binds to iron oxyhydroxides and is remobilized upon their reductive dissolution in anoxic sediment. We, therefore, first discuss the geochemistry of iron. At nearly all sampled locations in Lake Superior, broad, dense metal-rich layers (Fig. 3) found near the oxic-anoxic boundary indicated extensive redox cycling of Fe and Mn around that boundary. While Fe and Mn enrichments are typical in aquatic sediments, these layers in Lake Superior are exceptionally prominent and visible to the naked eye. Layers similar in appearance were previously reported in Lake Baikal (Granina et al. 2004; Och 4
Li et al.
P cycling in deeply oxygenated sediments
SRP ( μ mol L 1) 0 Depth (cm)
0
30 60 90
0
30 60 90
FWM Jun09
0
30 60 90
Jun10
0
30 60 90
Jul10
0
50
100
0
30 60 90 120
Apr11
Sep10
Aug11
5 10
SRP
15
Fe2+
20 25 0 Depth (cm)
0
30 60 90 IR
0
Jun10
30 60 90 0 July10
30 60 90 0 Sep10
30 60 90
0 30 60 90 0 Aug11
30 60 90 120 July12
5 10 15 20 Apr11
25 0
2
Depth (cm)
0
4
0
EM Jun09
30 60 90 EM
0
30 60 90
Jun10
0
30 60 90
July10
0 30 60 90 120
Sep10
5
5
10
10
15
15
20
20
25 0
2
Depth (cm)
0
4
0
2
WM Jun09
4
0
30 60 90
Jun10
0
2
4
July10
6
0 0
Sep10
5
5
10
10
15
15
20
20
25 0 Depth (cm)
20 SW
2
25 40 0 Aug11
20
6
40
60
July12
20 CM
40
0
20
Jun10
40
0
July10
2
4
0
20
Sep10
40
60
0 0
Apr11
2
4
6
0 0
NIP July12
5
5
5
10
10
10
15
15
15
20
20
20
25 0 0
4
Sta.2 Jun09
25 0
Depth (cm)
0 0
July12
20
40
0
20
40
0
30
60
0
25 30 60 90 120
0
10
20
10 KW
25 30
0
0
0
5
5
5
10
10
10
10
15
15
TB
15
15
20
20
July12
20
20
25
25
25
25
ED
July10
Sep10
Apr11
NB July12
0
20
30
Jun10
20 BB
40
60
July12
5
Fig. 2. Dissolved SRP and Fe21 in sediment porewater. Horizontal dashed and dotted lines represent the depths of oxygen and nitrate penetration, respectively (dashed-dotted line is used when OPD and NPD overlaps).
5
Li et al.
P cycling in deeply oxygenated sediments
Table 2. Diffusive fluxes of soluble reactive phosphorus (FSRP) across the SWI. Fluxes are positive into the sediments (see Li et al. 2012 and Li 2014 for the diffusive oxygen fluxes [Fo2 ] and the measured and calculated [indicated by *] total oxygen uptakes [TOU]). P recycling efficiency (P recycling E) is defined as the ratio of phosphorus effluxes to the total sedimentation flux of organic phosphorus.
Station
Fo2 (mmol m22 d21)
TOU (mmol m22 d21)
FSRP (lmol m22 d21)
P Recycling E
FWM.1
1.6
6.4
22.9 6 0.3
—
FWM.3
4.1; 2.1
—
23.0 6 0.3
—
FWM.4 FWM.5
3.2; 2.5 1.5
— 7.1
22.1 6 0.2 21.6 6 0.2
— —
FWM.6
3.3; 3.1; 4.2; 7.8
—
23.1 6 0.3; 25.9 6 0.7
—
Average EM.1
3.3 2.6
6.8 —
–2.8 6 0.7 23.4 6 0.4
10% —
EM.3
3.1; 2.3
—
22.5 6 0.3
—
EM.4 EM.5
4.6 2.5
— —
25.3 6 0.5 22.3 6 0.2
— —
EM.6
2.1; 2.7
—
24.9 6 0.5
—
2.9 3.4
4.4* 4.4
–3.7 6 0.5 23.7 6 0.2
16% —
Average WM.1 WM.3
1.2
—
21.6 6 0.2
—
WM.4 WM.5
2.2; 2.1 2.8
— —
22.2 6 0.2 22.6 6 0.3
— —
Average
2.3
4.4
–2.5 6 0.3
13%
IR.1 IR.2
4.5 5.9
— —
23.9 6 0.4 25.4 6 0.3
— —
IR.3
3.6
4.9 6 0.8
25.0 6 0.6
—
IR.4 IR.5
5.7; 7.3 3.8
— —
211 6 1; 26.0 6 0.7 24.3 6 0.5; 23.6 6 0.4
— —
Average
5.1
7.7*
–5.4 6 1.0
13%
CM.1 CM.2
2.8 3.7
— —
24.9 6 0.5 23.3 6 0.2
— —
CM.3
1.9
—
23.3 6 0.4
—
CM.4 Average
3.8 3.2
— 4.8*
211 6 1.0 –5.6 6 3.0
— 21%
ED.1
3.0
—
22.5 6 0.3
—
ED.2 ED.3
2.7 6.8
— —
22.7 6 0.3 211 6 1
— —
Average
4.2
6.3*
–5.4 6 2
16%
6.5 4.0; 3.9
— —
23.7 6 0.4 24.3 6 0.5
— —
Average
4.9
7.4*
–4.0 6 0.5
10%
KW.1 NIP.1
2.9
4.4*
–2.1 6 0.2 –6.6 6 0.7; –1.6 6 0.2
9% —
NB.1
6.8; 6.1
9.7*
–6.5 6 0.7
12%
BB.1 TB.1
8.7; 5.4 4.4; 5.8
11* 7.7*
–4.0 6 0.3 –7.0 6 0.7
7% 17%
SW.1 SW.2
further supported by the distributions of dissolved iron (Fig. 2) and solid phase iron speciations (Fig. 3). Increases in dissolved Fe21 and solid-phase Fe(II) below the OPD indicate iron reduction, which accounts for only < 0.2% of total carbon mineralization (R*Fe 5 0.047 mmol m22 d21, Table 3; carbon mineralization is 5.7 mmol m22 d21 [Li et al. 2012]).
et al. 2012). Comparison of the XRF iron and manganese profiles with the distributions of Ti, which is an element typically associated with detrital sediment components, suggests that these Fe- and Mn-rich layers are diagenetic in origin (Fig. 3). Fe and Mn levels within these layers are well in excess of their detrital components. The diagenetic origin is 6
Li et al.
P cycling in deeply oxygenated sediments
XRF Intensity (arbitrary unit) 0
0
5
10
0
5
10
0
5
10
0
5
10
0
5
10
0
5
10
15
5
Depth (cm)
10 15 20
Fe:Ti Mn:Ti
25
FWM.1
IR.1
EM.3
WM.4
CM.2
ED.3
30 Fig. 3. Scanning XRF intensity for total Fe and Mn in sediments of Lake Superior. Fe and Mn are compared to Ti, a detrital sediment element. Horizontal dashed and dotted lines represent the depths of oxygen and nitrate penetration, respectively (dashed-dotted line is used when OPD and NPD overlaps).
Solid Fe (mg/g) 0
0
10
20
Solid Fe (mg/g)
30 0
10 20 30
Solid P (mg/g) 0
1
2
3
4
Solid Fe (mg/g) 5
0
5
FWM.1
Solid P (mg/g)
10 15 20 WM.4
0
0.5
1
1.5
HCl Fe(II) 2
HCl Fe(III) HCl Total Fe NaAc Fe
SM.2
Depth (cm)
5
Dithionite Fe Total Fe Detrital P Authigenic P
10 15
Fe bound P
20
Residual P
25
Solid Fe (mg/g) 0
Total P Fe:P XRF Fe
0 20 40 60 80 100 Fe:P (mol:mol)
0
10
20
Solid P (mg/g)
Solid Fe (mg/g)
30 0 0.5 1 1.5 2 2.5 0
10 20 30
IR.1
Solid P (mg/g) 0
0.5
1
Solid Fe (mg/g) 1.5 0 10 20 30 40
EM.3
Solid P (mg/g) 0
2
4
6
8
ED.3
Depth (cm)
5 10 15 20 25
5
10 15 Fe:P (mol:mol)
0
20 40 60 80 Fe:P (mol:mol)
0 10 20 30 40 50 Fe:P (mol:mol)
Fig. 4. Solid phase Fe and P in sediments. Horizontal dashed and dotted lines represent the depths of oxygen and nitrate penetration, respectively (dashed-dotted line is used when OPD and NPD overlaps).
7
Li et al.
P cycling in deeply oxygenated sediments
Fig. 5. Calculated rates of Fe and P mobilization in sediments. Horizontal dashed and dotted lines represent the depths of oxygen and nitrate penetration, respectively. deeply oxygenated sediments, the depth of oxygen penetration is sensitive to the fluxes of reactive organic matter and can move by several centimeters in response to relatively minor variations in sedimentation flux (Katsev et al. 2006b). In Lake Superior, seasonal migrations in oxygen penetration by up to 2 cm have been observed (Li et al. 2012), and longerterm displacements can occur in response to changes in sedimentation, sediment resuspension, or changes in primary production patterns (Li 2014). Discordant positions of metal layers relative to the present-day OPD may provide insights into the redox boundary excursions. For example, at Sta. FWM, millimeter-wide Mn and Fe-rich layers are found within the oxic sediment layer, about 4 cm above the measured OPD. This violation of the traditional redox sequence suggests that the OPD had been shallower in the past but deepened to its present location. At Sta. FWM, the initial shallowing of the redox boundary was likely caused by the deposition of taconite (depleted iron ore) tailings from the nearby town of Silver Bay in 1950–1980s, which increased the oxygen demand. The redox boundary then migrated downward following the cessation of taconite discharges (Li 2011). At several other locations, additional iron-rich layers are found below the depth of
The rates of iron reduction are reasonably well correlated with the depth of oxygen (and nitrate) penetration, because a deeper oxygen penetration limits the amount of reactive carbon reaching the iron reduction zone (Fig. 7). Yet, over time, reoxidation of the produced upward-diffusing Fe(II) produces the observed prominent solid-phase Fe layers, at the depth where it meets oxygen (or an alternative electron acceptor such as nitrate or Mn oxides). The observation that manganese layers are located immediately above the uppermost iron layers is consistent with the typical diagenetic redox sequence: manganese oxides are more thermodynamically favorable as electron acceptors for organic carbon degradation than iron oxides, thus they are reduced (and reprecipitated) above the zone of iron reduction. The Fe-rich layers in Lake Superior are associated with large quantities of Fe-bound P, therefore redox cycling of Fe has the potential to strongly influence the mobilization/immobilization dynamics of phosphorus, at least in the vicinity of the sediment redox boundary. The presence of multiple metal-rich layers at many locations in Lake Superior (Fig. 3) indicates that the redox boundary within these sediments experienced vertical migrations. In 8
Li et al.
P cycling in deeply oxygenated sediments
Table 3. Depth-integrated rates of iron reduction (R*diss.Fe(II), R*solid.Fe(II), R*Fe reduction zone (R*P d21.
Station
rel.;
total), rates of phosphorus mobilization in the iron Eq. 5), and percentage of carbon mineralization by iron reduction (R*Fe total : R*C). All rates are in mmol m22
R*diss.Fe(II)
R*solid.Fe(II)
R*Fe
R*P
total
rel.
R*diss.Fe(II) : R*P rel.
R*C
R*Fe
total
: R*C
FWM.1
0.022
—
—
0.0095
—
—
—
FWM.3
0.0047
0.031
—
0.032
—
—
—
FWM.4 FWM.5
0.0077 0.0011
— —
— —
0.017 0.015
— —
— —
— —
FWM.6
0.051
—
—
0.051
—
—
—
FWM.7 Average
0.023 0.018
— —
— 0.049
0.024 0.025
— 0.72
— 5.2
— 0.1%
IR.1
0.015
0.031
—
0.034
—
—
—
IR.2 IR.3
0.044 0.043
— —
— —
0.073 0.028
— —
— —
— —
IR.4
0.046
—
—
0.18
—
—
—
IR.5 IR.6
0.080 0.068
— —
— —
0.18 0.063
— —
— —
— —
Average
0.045
—
0.076
0.093
0.48
5.3
0.2%
EM.3 EM.4
0.0033 0.012
0.01 —
— —
0.014 0.035
— —
— —
— —
EM.5
0.0090
—
—
0.037
—
—
—
EM.6 Average
0.0027 0.0067
— —
— 0.017
0.030 0.029
— 0.23
— 3.8
— 0.1%
WM.4
0.0034
—
—
0.009
—
3.3
—
CM.1 CM.2
0.0014 0.0037
— —
— —
0.013 0.0072
— —
— —
— —
CM.4
0.0049
—
—
0.022
—
—
—
Average ED.2
0.0033 0.0013
— —
— —
0.014 0.0077
0.24
4.1 —
— —
ED.3
0.0020
—
—
0.0083
—
—
Average KW.1
0.0017 0.0015
— —
— —
0.0080 0.0083
0.21 0.18
5.4 3.8
— —
SW.1
0.016
—
—
0.029
—
—
—
SW.2 Average
0.0041 0.0099
— —
— —
0.017 0.023
— 0.43
— 6.4
— —
NB
0.0092
—
—
0.016
0.58
8.4
—
TB BB
0.079 0.007
— —
— —
0.14 0.0055
0.56 1.27
6.6 9.4
— —
Average
0.018
—
0.047
0.034
0.49
5.7
1%
rich layer that is 0.25 cm thick and has solid Fe concentrations on the order of 30 mg/g (Fig. 3) and porosity around 0.85, the upward fluxes of Fe21 on the order of 0.018 mmol m22 d21 (Table 3) correspond to the time of formation of about 75 yr. Formation/dissolution of much thinner Mn-rich layers could correspondingly occur over a time scale of several years to decades. This explains, in particular, the absence of a Mn layer at the upper boundary of the relict Fe-rich layer at Sta. FWM, which is currently located within anoxic sediment.
oxygen penetration (Fig. 3), indicating that the downward migration of OPD is not a lake-wide phenomenon, and longterm displacements in the positions of this important redox boundary are likely governed by local processes such as shifts in sedimentation patterns. Under slowly varying conditions, the rate of precipitation is expected to be similar to the rate of dissolution of the layer as it gets buried below the oxicanoxic boundary. Therefore, dissolution of iron layers upon a vertical migration of the redox boundary could take place over similar time scales as formation of Fe layers. For a Fe9
Li et al.
P cycling in deeply oxygenated sediments
(arbitrary unit) Concentration (arbituary unit) Rate (arbituary unit) (arbitrary unit)
100 200 300 400 O2 Solid P
0.5
0
0.5
1
R * Fe reduction (mmol cm –2 d–1 )
0
0
RP
5
Depth (cm)
Solid Fe
RFe
10 SRP
15 20
Fe2+
R vs. OPD R vs. NPD
10–1
10–2
10–3
25
0
2
4
6
O 2 and
30 O2
RFe
Solid Fe
RP
Fe2+
8 NO –3
10
12
14
16
penetration (cm)
Fig. 7. Rates of Fe reduction vs. depths of oxygen and nitrate penetrations.
the iron reduction zone, SRP concentrations increase concomitantly with the release of dissolved Fe(II) (Fig. 4). The ratio of the depth-integrated rates of Fe and P mobilization (RFeðIIÞ : R*SRP, ranged 0.18–1.27 and averaged 0.49; Table 3) in the iron reduction zone was lower than the corresponding concentration ratio Fe : P in the solid phase (min 3–6; Fig. 4), indicating that more P is released into the porewater relative to Fe. This likely reflects a proportionally greater immobilization of reduced Fe into solid Fe(II) phases (Fig. 4) such as Fe sulfides, which do not adsorb phosphate, as well as continued release of phosphate from organic matter. The oxidized surface layer of sediments generally serves as an efficient trap for phosphate. Despite the availability of fresh organic matter, the net phosphate production rates there are small (Fig. 5) and SRP levels within this layer increase downward from the SWI at a slower rate compared to the rates in anoxic layers. The actual rates and SRP concentration gradients within the bioirrigation zone (< 2 cm) may be higher (by 30–50%), when taking bioirrigation into account (Li et al. 2012). The low concentrations of SRP and low rates of P production in the surface sediments likely reflect the high rate of P immobilization with Fe(III) phases. In contrast to the Fe rich layer at the Fe reduction-oxidation boundary, surface sediments are typically characterized by high Fe(III) : P ratios (40–80) (with the exception of Sta. ED where the Fe(III) : P ratio is lower near the surface), indicating strong capacity for binding phosphate, provided that the sorption capacity of solid phases in the surface oxic layer is similar (at least within an order of magnitude) to the sorption capacity of the ironrich layers at the oxic-anoxic boundary. Phosphate regenerated from organic matter in the oxidized surface sediments thus likely becomes adsorbed to oxidized sediment solids.
Solid P PO SRP 4 Fig. 6. Schematic profiles of dissolved O2, Fe21, SRP, solid Fe, solid P, and rates of Fe mobilization (RFe) and P mobilization (RP).
Phosphorus cycling The cycling of P in Lake Superior sediments is shaped by the balance between P mobilization and immobilization within the sediment. Dissolved phosphate is produced by mineralization of organic phosphorus throughout the sediment column, while Fe-bound P is mobilized within the zone of iron reduction. Immobilization of P occurs during iron reoxidation through P adsorption and co-precipitation (Table 1; Carignan and Flett 1981; Roden and Edmonds 1997), and in the deeper anoxic sediment through the formation of reduced iron phosphates such as vivianite (Table 1; Rothe et al. 2014). In Lake Superior, strong interaction with iron oxides is evident by the dominance of the Febound solid P pool, whose distribution matches the Fe-rich layers (Fig. 3). The iron oxidation zone near the Fe(III)-rich layers coincides with the consumption of dissolved phosphorus (Fig. 4). Within the layers, the Fe : P ratios are significantly lower than in surrounding sediment (Fig. 3), revealing their role as disproportionately strong P sinks. The low Fe : P molar ratio of 4–7 within the Fe-rich layers at most locations (Fig. 3) corresponds to the typical values at which the surfaces of iron oxyhydroxides become saturated with phosphate (Jensen et al. 1992; Caraco et al. 1993; Roden and Edmonds 1997), depending on mineralogy. Below the Fe-rich layers, in 10
Li et al.
P cycling in deeply oxygenated sediments
FC 5 6/7 Fo2 (Li 2014). Given that Lake Superior sediments mineralize 88% of deposited organic carbon (Li et al. 2012), this translates into the ratio of P efflux to C sedimentation of (1 : 909) 3 (6/7)/0.88 5 1 P : 885 C. Given the approximately 1 P : 106 C stoichiometry of organic matter mineralization, this suggests that a fixed fraction ( 12%) of phosphorus deposited to the sediment with organic matter is released back into the water column as SRP. The remaining 88% of P must be immobilized in the sediment. Using the fluxes of phosphorus and oxygen at individual sites, P recycling efficiency was calculated to be in the range of 7–21%, averaging 13% (Table 2). The P recycling efficiency (average 12–13%) in Lake Superior is consistent with the typical P recycling efficiency 5–20% in other freshwater lakes (Caraco et al. 1990). Higher values are found in anoxic or more productive and seasonally anoxic lakes (e.g., Moore et al. 1998; Katsev and Dittrich 2013; Matisoff et al. 2016). The sediments in Lake Superior have rates of sedimentation rates and carbon mineralization that are similar to marine sediments in 200–2000 m water depths (Li et al. 2012). Yet marine sediments have much higher P regeneration efficiencies (> 50%; Caraco et al. 1990; Sundby et al. 1992; McManus et al. 1997), and their Cox : P regeneration ratios are closer to the Redfield ratio of 106 : 1. It is commonly believed that high rates of sulfate reduction and the associated high levels of H2S decrease the abundance of iron oxides and sediment P retention, therefore P regeneration is largely controlled by organic matter mineralization (Capone and Kiene 1988; Caraco et al. 1990, 1993). Even in ocean sediments where oxygen penetration and the zone of sulfate reduction are deep (e.g., deep-sea sediments; Canfield 1991) so that H2S would not affect Fe and P cycling in surface sediments, P effluxes could be expected to be controlled by organic matter mineralization near the surface. P regeneration efficiency in these deep-sea sediments is consistent with the Redfield decomposition of organic matter (McManus et al. 1997), and is substantially higher than in sediments of Lake Superior. The difference in P retention/regeneration efficiency between the deep-sea sediments and sediments in Lake Superior suggests the difference in mechanisms of P retention. In marine sediments, the major P sink is formation and burial of carbonate fluorapatite in deep sediments (Ruttenberg 1992). The composition of surface sediments (i.e., the mineralogy of deposited material and possibly the amounts of organic substances that may serve as P sorption sinks), on the other hand, may be important for controlling the P regeneration near the sediment surface (Hupfer and Lewandowski 2008). The acetate- and dithionite-extractable iron (carbonate iron and reducible iron oxides) are the major components that adsorb phosphorus (Ruttenberg 1992). In Lake Superior, they account for 1.5–3.5% of the solid sediment weight. In marine sediments, in contrast, dithionite extractable iron typically accounts for only 0.2–1.5% (Canfield 1989). The P retention capacity in Lake Superior thus
12 FWM EM WM CM ED SW KW IR BB TB NB
F SRP ( mol m –2 d–1 )
10 8 6 4
F SRP =1.1x10 –3 F O 2 r 2 = 0.62
2 0 0
2
4
6
8
F O 2 (mmol cm –2 d–1 ) Fig. 8. Flux of SRP out of sediment vs. total oxygen uptake. Diffusive oxygen fluxes (Fo2 ) are from Li et al. (2012) and Li (2014).
Controls on phosphorus fluxes across the sediment-water interface Whereas a significant amount of P in Lake Superior sediments is strongly bound to iron, the processes that take place near the deep oxic-anoxic boundary have relatively little effect on phosphate concentrations in porewater near the sediment-water interface. The great depth at which iron reduction occurs in these well-oxygenated sediments ensures that the effluxes of phosphorus from sediments are only weakly affected by iron reduction. The thick oxidized layer provides a strong P-adsorption capacity that mitigates any changes in P mobilization below. In many shallower and more productive lakes, seasonal onset of reducing conditions cause the release of sediment P into the water column (Mortimer 1942). In contrast, in Lake Superior, the flux of phosphate across the SWI is largely controlled by processes that take place entirely within the oxic layer. P effluxes are determined by the balance between the phosphate production from organic phosphorus and its adsorption to iron oxides and other sediment solids within the oxic zone, often as much as several cm above the sediment redox boundary. The surfaces of iron phases within the oxic zone are likely far from being saturated with phosphate, therefore the sorption isotherm becomes an approximately linear function of porewater P concentration (Jensen et al. 1992). The effluxes of SRP across the SWI then vary approximately linearly with organic carbon loading, because within the oxic zone dissolved P comes predominantly from organic matter. Indeed, the SRP fluxes exhibit a ratio of 1 P : 909 O2 with oxygen uptake rates (Fig. 8; FSPR 5 1.1 3 1023 Fo2 ). Carbon mineralization in Lake Superior sediments can be estimated as 11
Li et al.
P cycling in deeply oxygenated sediments
fluxes are likely to be on the order of 6 lmol m22 d21. These are broadly consistent with fluxes previously estimated in sediment core incubations in the Western Arm of the lake (5 lmol m22 d21; Heinen and McManus 2004). A similar figure can be obtained using the estimated P recycling efficiency (12–13%) and the organic P sedimentation rate estimated from carbon sedimentation (assuming 1P : 106C stoichiometry); P efflux is estimated using this method at 7.4 6 2.3 lmol m22 d21 (organic carbon settling rate 6.5 6 2.0 mmol m22 d21; Li 2014). Based on previously estimated P inputs from watershed (Fig. 9), inputs of P from sediment (release of dissolved P from sediment to water) account for 40% of the total P inputs into the water column, a significant contribution to supplying the productivity-limiting nutrient. The low efficiency of P recycling makes sediment an important sink for P. A large percentage of the deposited phosphorus (88%) is permanently buried in the sediments. Based on organic P sedimentation numbers, detrital and non-reactive organic P is buried at the rate of 38 6 13 lmol m22 d21 (Fig. 9). The burial flux of P can be also estimated from solid-phase phosphorus concentrations (CL; Fig. 4) and known burial velocities (UL) and porosity (uL) in the deep sediment as Fburr. 5 CLUL(1-uL)q. The obtained value of 29 lmol m22 d21 is largely consistent with the estimate above (38 6 13 lmol m22 d21). The current P budget for the lake appears imbalanced, with P sinks significantly exceeding sources (Fig. 9). The 2306 13 lmol m22 d21 unbalanced in P fluxes would correspond to a > 1 lmol L21 decrease in the water column TP levels since the 1970s (calculated using surface area of 82,100 km2 and volume of 11,920 km3). Historical data for the water column TP concentrations, however, show much smaller decreases in Lake Superior (less than 0.1 lmol L21 over 40 yr), indicating the need to reconcile P loading estimates from different sources. Potential reasons for the discrepancy may include, for example, an underestimation of the external P sources to Lake Superior in previous studies. This may include phosphorus inputs from shoreline erosion that were probably underestimated in previous studies, because data were largely unavailable (Chapra and Sonzogni 1979). There may have been a significant and underappreciated role of sediment resuspension that enhances P remobilization from sediments beyond diffusive and bioirrigation fluxes, or a heterogeneity of the lake floor that extends beyond the range covered by sampling locations in this study. These remaining significant uncertainties in several major P sources and sinks (Fig. 9) indicate the need for further studies of the P fluxes and dynamics in this important lake.
Atmospheric deposition Outflow 0.43 Watershed (tributary) 2.4 0.89 0.29 Point source 4.1 Shoreline erosion Sedimentation 43 ± 13 Flux at SWI 5.2 ± 1.6 Water column: Total input = 13 ± 2 Total output = 43 ± 13
38 ± 13 Burial
Entire lake: Total input = 7.7 Total output = 38 ± 13
Fig. 9. Sources and sinks of phosphorus in Lake Superior (lmol m22 d21). Organic P sedimentation flux is calculated from the organic carbon settling rate of 6.5 6 2.0 mmol m22 d21 (Li 2014), with a 106C : 1P stoichiometry (Heinen and McManus 2004), assuming 30% of the lake floor is nondepositional (actual number not known; Kemp et al. 1978). The flux of P across the SWI is estimated using the organic P sedimentation and P recycling efficiency of 12%. Data for atmospheric deposition, watershed, and point source input are from Dolan and Chapra (2012); data for shoreline erosion are from Chapra and Sonzogni (1979); outflow is from Weiler (1978).
appears to be enhanced not only by the deep penetration of oxygen but also by the high sedimentation fluxes of iron. Phosphorus budget in Lake Superior The sediment P fluxes estimated above allow better estimates for the sediment contribution to the total phosphorus loading into the Lake Superior water column. Although compiling the P budget for such a large lake necessarily involves significant uncertainties associated with a finite number of sampling locations, compiling such a budget is instructive for understanding the P dynamics in the lake and identifying existing knowledge gaps. The large number of locations sampled in this study across the entire lake offers a substantial improvement over the previous estimates, which were significantly more limited in scope. Figure 9 presents the phosphorus budget in Lake Superior with updated contributions from sediments. A previous budget (Heinen and McManus 2004) was calculated based on sedimentation fluxes measured in sediment traps in the Western Arm of the lake. The sediment traps, however, are known to underestimate P sedimentation, because organic material may undergo substantial degradation while in the traps. Figure 9 also accounts for the more recent P input data (atmospheric deposition, watershed loading, point-source pollution, etc.) that were not available previously (Dolan and Chapra 2012). Organic P sedimentation (–37 lmol m22 d21) appears to be a major sink of P for the lake’s water column. The calculated diffusive effluxes of phosphorus from sediments are 2.5–7.0 lmol m22 d21, average 4.4 lmol m22 d21. Total fluxes may exceed the diffusive fluxes by 30–50% due to processes such as bioirrigation (Li et al. 2012), therefore, the average total
References Baehr, M. M., and J. McManus. 2003. The measurement of phosphorus and its spatial and temporal variability in the western arm of Lake Superior. J. Great Lakes Res. 29: 479– 487. doi:10.1016/S0380-1330(03)70452-3 12
Li et al.
P cycling in deeply oxygenated sediments
paradigm in limnology. Int. Rev. Hydrobiol. 93: 415–432. doi:10.1002/iroh.200711054 Jensen, H. S., P. Kristensen, E. Jeppesen, and A. Skytthe. 1992. Iron-phosphorus ratio in surface sediment as an indicator of phosphate release from aerobic sediments in shallow lakes. Hydrobiologia 235–236: 731–743. doi: 10.1007/BF00026261 Johnson, T. C., J. E. Evans, and S. J. Eisenreich. 1982. Total organic-carbon in Lake-Superior sediments - comparisons with hemipelagic and pelagic marine environments. Limnol. Oceanogr. 27: 481–491. doi:10.4319/lo.1982.27.3.0481 Katsev, S. 2016. Phosphorus effluxes from lake sediments, p. 115–131. In R. Lal and B. A. Stewart [eds.], Soil phosphorus. CRC Press. Katsev, S. 2017. When large lakes respond fast: A parsimonious model for phosphorus dynamics. J. Great Lakes Res. 43: 199–204. doi:10.1016/j.jglr.2016.10.012 Katsev, S., I. Tsandev, I. L’Heureux, and D. G. Rancourt. 2006a. Factors controlling long-term phosphorus efflux from lake sediments: Exploratory reactive-transport modeling. Chem. Geol. 234: 127–147. doi:10.1016/ j.chemgeo.2006.05.001 Katsev, S., B. Sundby, and A. Mucci. 2006b. Modeling vertical excursions of the redox boundary in sediments: Application to deep basins of the Arctic Ocean. Limnol. Oceanogr. 51: 1581–1593. doi:10.4319/lo.2006.51.4.1581 Katsev, S., and M. Dittrich. 2013. Modeling of decadal scale phosphorus retention in lake sediment under varying redox conditions. Ecol. Model. 251: 246–259. doi: 10.1016/j.ecolmodel.2012.12.008 Kemp, A., C. Dell, and N. Harper. 1978. Sedimentation rates and a sediment budget for Lake Superior. J. Great Lakes Res. 4: 276–287. doi:10.1016/S0380-1330(78)72198-2 Li, J. 2011. Diagenesis and sediment-water exchanges in organic-poor sediments of Lake Superior. M.S. thesis. Univ. of Minnesota. Li, J. 2014. Sediment diagenesis in large Lakes superior and Malawi, geochemical cycles and budgets and comparisons to marine sediments. Ph.D. dissertation. Univ. of Minnesota. Li, J., S. A. Crowe, D. Miklesh, M. Kistner, D. E. Canfield, and S. Katsev. 2012. Carbon mineralization and oxygen dynamics in sediments with deep oxygen penetration, Lake Superior. Limnol. Oceanogr. 57: 1634–1650. doi: 10.4319/lo.2012.57.6.1634 Li, J., and S. Katsev. 2014. Nitrogen cycling in deeply oxygenated sediments: Results in Lake Superior and implications for marine sediments. Limnol. Oceanogr. 59: 465– 481. doi:10.4319/lo.2014.59.2.0465 Matisoff, G., and others. 2016. Internal loading of phosphorus in western Lake Erie. J. Great Lakes Res. 42: 775–788. doi:10.1016/j.jglr.2016.04.004 McManus, J., W. M. Berelson, K. H. Coale, K. S. Johnson, and T. E. Kilgore. 1997. Phosphorus regeneration in
Boudreau, B. P. 1997. Diagenetic models and their implementation. Springer Berlin. Canfield, D. E. 1989. Reactive iron in marine-sediments. Geochim. Cosmochim. Acta 53: 619–632. doi:10.1016/ 0016-7037(89)90005-7 Canfield, D. E. 1991. Sulfate reduction in deep-sea sediments. Am. J. Sci. 291: 177–188. doi:10.2475/ajs.291.2.177 Capone, D. G., and R. P. Kiene. 1988. Comparison of microbial dynamics in marine and fresh-water sediments - contrasts in anaerobic carbon catabolism. Limnol. Oceanogr. 33: 725–749. doi:10.4319/lo.1988.33.4_part_2.0725 Caraco, N., J. Cole, and G. E. Likens. 1990. A comparison of phosphorus immobilization in sediments of fresh-water and coastal marine systems. Biogeochemistry 9: 277–290. doi:10.1007/BF00000602 Caraco, N. F., J. J. Cole, and G. E. Likens. 1993. Sulfate control of phosphorus availability in lakes - a test and reevaluation of Hasler and Einsele model. Hydrobiologia 253: 275–280. doi:10.1007/BF00050748 Carignan, R., and R. J. Flett. 1981. Post-depositional mobility of phosphorus in lake-sediments. Limnol. Oceanogr. 26: 361–366. doi:10.4319/lo.1981.26.2.0361 Chapra, S. C., and W. C. Sonzogni. 1979. Great-Lakes total phosphorus budget for the mid 1970s. J. Water Pollut. Control. Fed. 51: 2524–2533. http://www.jstor.org/stable/25040455 Chapra, S. C., and D. M. Dolan. 2012. Great Lakes total phosphorus revisited: 2. Mass balance modeling. J. Great Lakes Res. 38: 741–754. doi:10.1016/j.jglr.2012.10.002 Dolan, D. M., and S. C. Chapra. 2012. Great Lakes total phosphorus revisited: 1. Loading analysis and update (1994–2008). J. Great Lakes Res. 38: 730–740. doi:10.1016/j.jglr.2012.10.001 Glud, R. N. 2008. Oxygen dynamics of marine sediments. Mar. Biol. Res. 4: 243–289. doi:10.1080/17451000801888726 Granina, L., B. Muller, and B. Wehrli. 2004. Origin and dynamics of Fe and Mn sedimentary layers in Lake Baikal. Chem. Geol. 205: 55–72. doi:10.1016/j.chemgeo.2003.12.018 Gunnars, A., S. Blomqvist, and C. Martinsson. 2004. Inorganic formation of apatite in brackish seawater from the Baltic Sea: An experimental approach. Mar. Chem. 91: 15–26. doi:10.1016/j.marchem.2004.01.008 Hecky, R. E., and P. Kilham. 1988. Nutrient limitation of phytoplankton in fresh-water and marine environments - a review of recent-evidence on the effects of enrichment. Limnol. Oceanogr. 33: 796–822. doi:10.4319/lo.1988.33.4part2.0796 Heinen, E. A., and J. McManus. 2004. Carbon and nutrient cycling at the sediment-water boundary in western Lake Superior. J. Great Lakes Res. 30: 113–132. doi:10.1016/ S0380-1330(04)70381-0 Higgins, S. N., and M. J. Vander Zanden. 2010. What a difference a species makes: A meta-analysis of dreissenid mussel impacts on freshwater ecosystems. Ecol. Monogr. 80: 179–196. doi:10.1890/09-1249.1 Hupfer, M., and J. Lewandowski. 2008. Oxygen controls the phosphorus release from lake sediments - a long-lasting 13
Li et al.
P cycling in deeply oxygenated sediments
sediments. Limnol. Oceanogr. 37: 1460–1482. doi: 10.4319/lo.1992.37.7.1460 Sondergaard, M., P. J. Jensen, and E. Jeppesen. 2001. Retention and internal loading of phosphorus in shallow, eutrophic lakes. Sci. World J. 1: 427–442. doi:10.1100/tsw.2001.72 Sterner, R. W., and others. 2007. Increasing stoichiometric imbalance in North America’s largest lake: Nitrification in Lake Superior. Geophys. Res. Lett. 34: L10406. doi: 10.1029/2006GL028861 Sundby, B., C. Gobeil, N. Silverberg, and A. Mucci. 1992. The phosphorus cycle in coastal marine-sediments. Limnol. Oceanogr. 37: 1129–1145. doi:10.4319/lo.1992.37.6.1129 Viollier, E., P. W. Inglett, K. Hunter, A. N. Roychoudhury, and P. Van Cappellen. 2000. The ferrozine method revisited: Fe(II)/Fe(III) determination in natural waters. Appl. Geochem. 15: 785–790. doi:10.1016/S0883-2927(99)00097-9 Weiler, R. 1978. Chemistry of Lake Superior. J. Great Lakes Res. 4: 370–385. doi:10.1016/S0380-1330(78)72207-0
continental margin sediments. Geochim. Cosmochim. Acta 61: 2891–2907. doi:10.1016/S0016-7037(97)00138-5 Meile, C., P. Berg, P. Van Cappellen, and K. Tuncay. 2005. Solute-specific pore water irrigation: Implications for chemical cycling in early diagenesis. J. Mar. Res. 63: 601– 621. doi:10.1357/0022240054307885 Moore, P. A., K. R. Reddy, and M. M. Fisher. 1998. Phosphorus flux between sediment and overlying water in Lake Okeechobee, Florida: Spatial and temporal variations. J. Environ. Qual. 27: 1428–1439. doi:10.2134/jeq1998.00472425002700060020x Mortimer, C. H. 1942. The exchange of dissolved substances between mud and water in lakes. J. Ecol. 30: 147–201. doi:10.2307/2256691 Mosley, C., and H. Bootsma. 2015. Phosphorus recycling by profunda quagga mussels (Dreissena rostriformis bugensis) in Lake Michigan. J. Great Lakes Res. 41: 38–48. doi: 10.1016/j.jglr.2015.07.007 Och, L. M., and others. 2012. New insights into the formation and burial of Fe/Mn accumulations in Lake Baikal sediments. Chem. Geol. 330–331: 244–259. doi:10.1016/ j.chemgeo.2012.09.011 Poulton, S. W., and D. E. Canfield. 2005. Development of a sequential extraction procedure for iron: Implications for iron partitioning in continentally derived particulates. Chem. Geol. 214: 209–221. doi:10.1016/j.chemgeo.2004.09.003 Roden, E. E., and J. W. Edmonds. 1997. Phosphate mobilization in iron-rich anaerobic sediments: Microbial Fe(III) oxide reduction versus iron-sulfide formation. Arch. Hydrobiol. 139: 347–378. Roden, E. E., and R. G. Wetzel. 2002. Kinetics of microbial Fe(III) oxide reduction in freshwater wetland sediments. Limnol. Oceanogr. 47: 198–211. doi:10.4319/lo.2002.47.1.0198 Rothe, M., T. Frederichs, M. Eder, A. Kleeberg, and M. Hupfer. 2014. Evidence for vivianite formation and its contribution to long-term phosphorus retention in a recent lake sediment: A novel analytical approach. Biogeosciences 11: 5169–5180. doi:10.5194/bg-11-5169-2014 Ruttenberg, K. C. 1992. Development of a sequential extraction method for different forms of phosphorus in marine-
Acknowledgments We thank Sean Crowe, Matthew Kistner, and David Miklesh for help with sample acquisition and processing, Jay Austin for sharing his National Science Foundation (NSF)-funded cruise opportunities in June and October 2009, Elizabeth Minor, Stephanie Guildford, and Josef Werne for sharing their laboratory facilities. We gratefully acknowledge the help of Captain Mike King and the crew of the R/V Blue Heron, marine technician Jason Agnich, and laboratory technician Sarah Grosshuesch. The work has been supported by the NSF Ocean Sciences (OCE) grant 0961720, the University of Minnesota Duluth start-up funds to SK, the Water Resources Science Block Grant, and University of Minnesota Duluth Physics Department summer fellowships to JL, and Undergraduate Research Opportunities Program to YZ.
Conflict of Interest None declared. Submitted 16 June 2017 Revised 15 October 2017 Accepted 07 December 2017
Associate editor: Caroline Slomp
14
