Early diagenesis in a reducing fjord, Saanich Inlet, British Columbia—I. chemical and isotopic changes in major components of interstitial water

Early diagenesistin a reducing fjord, f3aankh Inlet, British Columbia-L Chemical and isotopic changes in major components of interstitial water* ARIE NISSENBAUM,? B. J. PRESLEY: and I. R. KAPLAN Department of Geology and Institute of Geophysics and Planetary Physics, University of California, LOSAngeles 90024

(Received9 XeptembeT

I971; acceptedin retied

fm

30 March 1972)

&&&--Water and intemtitialwater from the reducing fjord of S8anich Inlet, British Columbia were analyzed for their major element composition, ammonia, phosphate and silica contents, and for stable isotope composition of sulfur and carbon species. Ca WES the only major element to show 8 sign&ant change with depth (a 75 per cent decreasein some oases). Annnonie and phosphate are highly enrichedin the interstitial water (I.W.), concentration8reaching 250 ppm and 39 ppm, respectively. Total diasolved CO, in I.W. increa8e8strongly with depth (20 to 30 times th8t in overlying see-water) and it becomes enriched in Cl* (dC&n M +17-S%). Both sulfate and dissolved sulfide decrease with depth to a complete disappearance of all sulfur speciesfrom the interstitialwater. The dissolved sulfideis highly enrichedin iF (SS= m + 18y&,). All these uhangeaare attributed to strong biological aotivity in the sediments. INTRODUCTION

A CHEMICAL study on the water, interstitial water and sediment of Saanich Inlet, British Columbia, has been undertaken in our leboratory. The ares, WM chosen as a model environment for geochemicitl studies because it is an isolated basin, and because of its anoxic character. Major and minor element concentrations and organic components were determined in water, ~terstiti~l water and various fractions of the sediment, and in rtddition, such components as dissolved COa, PO,“, SO,“, S” and NH, were also measured. Eh and pH were determined as were carbon and sulfur isotope ratios of isolated compounds. Location and general features of Saalaich Inlet Saanich Inlet is a fjord on the southeastern side of Vancouver Island (Fig. 1). The inlet is about 25 km long and O-4 to 7-6 km wide, and has an average depth of 200 m. It receives little direct runoff from streams, such as the Goldstream River in the southeastern extreme, but receives diluted seawater from the input of the Co~ch~n and Fraser Rivers through the Sattelite ChanneI (H~XU,XXVEA~X, 1962). The exchange of water between the inlet and the Sottelite Channel is restricted by a sill at the north end of the inlet, reaching within 70 m of the surface. The partial isolation of the inlet causes an ~termittent stagnation of the bottom waters for a few months in the year, characterized by oxygen disappearance and production of hydrogen sulfide. Late summer or early fall flushing causea * ~ub~~tion No. 872: Institute of Geophysicsand Planetary Physics, Wnivemity of California at Los Angeles 90024. t Present address: Isotope Department, Weizmann Institute, Rehovot, Israel. $ Preeent enidress: Department of Oceanogmphy, Texas A & M University, College Station, Texas. 1007

1008

*STATION

LOCATION

KILOMETERS I

I

I

40’

iZ3Q30’

20’

1I IO

Fig. 1. Location map and sampling stations in Sartnich Inlet.

Water overlying the sill never becomes re-oxygenation of the water column. anoxic. Saanich Inlet sediments have been described by GUCLUER and Gaoss (1964) and GILOSS(1967). The central part of the basin is covered with a black, varved, clayey silt, rich in sulfides, diatom frustules and organic matter. The sill sediments, on the other hand, are olive-gray silts with lower contents of opal and organic matter. The rate of sedimentation is extremely fast, and was estimated by Gucluer and Gross to be about 4 m per 1000 yr. BUDDEMEIER (1969) using CY techniques, also found comparable rates of sedimentation.

Most of the water and sediments used in this study were collected during a cruise of the University of Washington research vessel, ‘Thomas G. Thompson’ in

E:srly d&genesis in a reducing fjord, S&oh

Inlet, British Columbia-I

1009

December, 1967. The sampling locations are shown in Fig. 1. PVC 6-1 Niskin@ samplers were used for collecting the water, and a 6-in. diameter PVC-lined piston corer was employed for sediment sampling. The PVC core liners containing the sediments were capped immediately after collection and were stored in a horizontal position until the sediment was extruded in the laboratory, three or four days later. Both water and sediment samples were kept at 4°C during transport to the laboratory, and during subsequent storage. Several samples were obtained through the courtesy of Mr. R. Buddemeier of the University of Washington, Chemistry Department. Samples marked 3B were obtained from him and had been collected by Global Marine Company in June, 1968. ANALYTICAL

PROCEDWRES

Sea water and interstitial water Interstitial w&%r was removed from the sediment (which was generally still cold) by a Teflon~-I~~, ~trogen-oper&t~ squeezer at room temperature (PRESLEY, et aI., 1967). Eh, pH, dissolved CO, and H&Z were memed on small sliquots of water immediately. Separate "e were collected for analyses of ephemeral nutrients, (e.g. ammonis, phosph&e) and were frozen in glass ampoules for storage until analysis, whereas aliquots of I.W. intended for major and trace element analysis were stored at 4°C. Sea water from the water column was analyzed in part on board ship (pH, nutrients, dissolved CO,) and in part in our laboratory (major cations and snions). The sliquot of water used for CO, and H&3 determination was obtained by collecting the effluent from the squeezer directly into a hypodermic syringe. This procedure minimizea gas exchange with the atmosphere and allows the wster to be conveniently transferred into a ma&ion flask. Dissolved HsS and CO, were determined on 20 ml samples in a flask containing 6% silver nitrate in a side arm. The H&3 was trapped in the side arm containing AgNO,, and weighed as AgsS. The CO, was purified by passing it through dry ice-seetone traps. Its volume was measured and the srunpleswere collected for C1g/C1z analyses. The technique has been described in greater detail elsewhere (PRESLEY and HAPUN, 1968). Ratios of CH,:CO,:N,O were obtained by gas-chromatography using a 5-ft Poropak column at a room temperature and a Carl& thermistor microdetector. Ammonia, phosphate and silicate were measured spectrophotometrically following the techniques of RICHLLRDSsnd KLETSCH (1964) for ammonia and STRTCKUKD and PARSONS (1960), for total reactive phosphate and silicate. Chloride eoncentmtion was determined by the argentometric Mohr titmtion and sulfate gmvimetricslly as BaSOp. Ca, Mg, K and Ne were analyzed by stomic absorption spectrophotometry. The ratios of S34/S52 and C!13/C”were measured on a Nuclide Corp. 6in. radius, McKinu%yNier double-collector type mass spectrometer. The results are expressed versus PDB standard (for &Y) and C&ion Diablo meteorite standard (for 6S94). RESULTS

Some data describing the composition of the water column are graphically presented in Fig. 2. It can clearly be seen that an abrupt change occurs at about 70 m (depth of sill), below which oxygen content almost disappears and nitrate diminishes; phosphate, silica and total CO, increase. These data constitute strong evidence for biologic activity within the water column. Major element composition and ion-chlorinity ratios are shown in Tables 1 and 6

T *C

Fig. 2. Chemical charact~erist~icsof Seanich Inlet water column.

200

160

160

140.

SF

Early

1011

diagenesis in a reducing fjord, Saanich Inlet, British Columbia-I

Table 1. Major component concentration in Saanich Inlet water and interstitial water Cl

Ne

K

Ca

g/l

g/l

S/l

g/l

Mg g/l

SO* g/l

lO*OO

11.06

0.397

0.424

1.330

2.79

8.2

Station 1 (Surfwe) 20m 40m 60m

14.93 16.78 16.85 17.05

7.70 9.00 9.00 9.01

0.36 0.38 0.41 0.42

0.28 0.31 0.29 0.31

0.90 1.09 I.09 1.12

1.85 2.23 2.28 2.33

6.0 7.0 7.0 7.2

Station 4 (Surface) 40 m 100 m 200 m

11.81 17.35 17.65 17.95

6.56 9.38 0.45 9.44

0.30 0.41 0.43 0.43

0.23 0.30 0.32 0.32

0.76 1.11 1.12 1.24

1.61 2.29 2.38 2.39

4.6 7.3 7.3 7.4

17.37 17.67 17.90

8.81 9.01 9.10 9.10 9.56

0.46 0.48 0.46 0.56 0.49

0.31 0.23 0.13 0.09 0.08

1.08 1.08 1.02 1.14 1.24

2.20 1.10 0.03 0.07 0.03

5.6 3.7 4.9 2.8 4.6

17.75 17.87 17.90 17.40

9.54 9.54 9.54 9.54

0.49 0.49 0.51 0.51

0.31 0.31 0.31 0.30

1.06 1.23 1.24 1.26

0.13 0.02 0.05 tr.

6.8 7.1 7.2 7.0

18.49 18.49 18.30 18.41 18.10

9.45 9.45 9.54 9.45 9.28

0.45 0.45 0.45 0.45 0.46

0.31 0.29 0.31 0.30 0.31

1.22 1.22 1.22 1.22 1.22

0.12 0.03 0.03 tr. tr.

6.9 6.8 6.8 6.8 6.6

17.83 17.91 17.28 17.35

9.17 9.33 9.25 9.08

0.49 0.44 0.53 0.52

0.33 0.20 0.03 0.05

1.20 1.23 1.19 1.23

0 0 0.18 tr.

5.5 4.0 2.0 2.6

17.79 17.99 18.12 18.04 18.12 18.11

9.64 9.54 9.72 9.54 9.35 9.19

0.45 0.48 0.48 0.48 0.47 0.46

0.34 0.34 0.31 0.34 0.31 0.30

1.22 1.22 1.23 I.24 1.22 1.22

0.21 0.18 0.12

7.2 7.2 6.9 6.0 7.0 6.8

AVWegEl 888 water*

Sr %/I

Water column

Interstitial water Core 1 Cl5 cm 40-50 cm 85-100 om 135-150 cm 175-186 cm Core 2 O-15 cm 75-85 cm 150-165 cm 225-235 cm Core 3 O-10 cm 50-60 cm 100-110 cm 150-160 cm 190-200 om Core 3B 7.9-8.2 m 17.1-17.4 m 26.2-26.5 m 34.5-34.8 m Core 4 O-16 cm 5&65 om 100-110 cm 150-160 cm 200-210 cm 240-250 cm

* C&mlaulated from CULKIN (1965) Table 1, pqe

2.

0%2 0.02

122.

There is a positive salinity gradient in the inlet at all times (HERLINVEAUX, and this is clearly seen from Table 1. Herlinveaux reports large variations in salinity in the upper few meters of the water column with time, due to a changing influx of fresh water. The water below sill depth, however, remains fairly constant in salinity throughout the year, and the values given in Table 1 are typical. 1962),

water column Station I (Surface) 20m 40m 60 m

14.93 16.78 16.85 17.05

0.516 0.536 0.534 0.528

O-0241 0~0226 0.0243 0.0246

0.0188 O%I185 0.0172 0.0182

0.0603 0.0650 0.0649 oax7

0.124 Q-133 0.133 0.137

Station 4 (Surface) 40m loom 200 m

11.81 17.35 17.65 17.95

0.555 0.541 O-536 0.5126

0.0254 O-0236 0.0244 O-0240

0.0195 0.0173 0.0181 0.0178

0.0644 0.0640 0.0635 0.0691

0.136 0.132 O-135 0.133

0.390 0.420 0.414 o-412

17.37 17.67 17.90

0.507 0.615 0.534

0.0265 O-0272 0.0273

0.0178 0.0130 0.005

0.0622 0.0611 0.0693

0.127 0.001 0+01

0.322 0.277 0.268

17.75 17.87 17.90 17.40

0.537 0.534 0.533 0.548

0.0276 0.0274 0.0285 0.0293

0.0175 0.0173 0.0173 0.0172

0.0597 0.0688 0.0693 0.0724

0.007 0.001 0.002 0.001

0.383 0.397 0.402 0.402

18.49 18.49 18.30 18.41 18.10

0.51 I 0.511 0.521 0.513 0.513

0.0265 0.0265 0.0246 0.0244 0.0254

0.0168 0.0157 0.0169 0.0163 0.0173

0~0660 0.0660 0.0666 0.0663 0.0674

0.007 0.002 0~002 o-000 0.000

0.373 0.368 0.372 0.369 0.365

17.83 17.91 17.28 17.35

0.514 ov521 0.535 0.523

0.0275 0.0246 0.0307 0.0300

0.0185 0~0110 0.0002 0.0002

0.0673 0.0687 0*0689 0.0709

0.000 o*ooo 0.010 0.000

0.308 0.223 O-116 0.150

17.79 17.99 18.12 18-04 18.12 18.11

0.536 0.530 0.536 0*52Q 0.516 0.507

0.0263 0.0267 0.0265 0.0266 0~0260 0.0254

0.0190 0.0189 0.0171 0*0189 0.0171 0.0166

0.0685 0.0678 0.0679 0.0687 0.0673 0.0674

0.012 0.010 O-007 o*ooo 0.001 0.001

0.405 0.400 0.380 O-382 0.386 0.375

Interstitial

0*402 O-417

0.415 0.422

water

Core 1 O-16 cm 85-100 cm 175-185 cm Core 2 O-15 om 75-M cm 150-165 cm 22Ec235 cm Core 3 Q-10 cm 60-60 em 100-I 10 om 150-160 om 190-200 cm Core 3B 7.9-8.2 m 17.1-17.4 m 26.2-26-5 m 34x%-34.8 m Core 4 o-15 cm 50-65 cm 100-l 10 cm 150-160 om 200-2 10 cm 240-250 cm

* Calcult-hed from COLKIN (1966) Tablo 1, page 122.

Interstitial water Cl is not si~i~ca~tly different from that of the overlying bottom waters in Saanich Inlet. Cores 1 and 4 show a tendency for increase in chlorosity down the core by 3 a/Oin Core 1 and 1.7 o/Oin core 4. The cores nearest the sill (cores 1 and 2) contain the lowest chloride concentration in the interstitial water. The chlorosity of the deep sediment interstitial water (core 3B) is approximately 4% lower than that in the interstitia1 water from the shallower core

Early diagenesis in a reducing fjord, S&oh

Inlet, British Columbia-I

1013

(No. 3). This may constitute evidence for a gradual salinity increase with time in the water of the fjord. Na concentrations (Table 1) are slso lower, and the Na/Cl ratio (Table 2) for interstitial water is 4 to 9 y0 lower than average sea water but not significantly different from bottom waters of the fjord. Potassium and magnesium The Saanich Inlet water below sill depth was found to be 10 to 13 % higher in K than average sea water. The interstitial water seemed to be further enriched at most depths by 10 to 12% over that in the water column, but in the deepest layers of core 3B, the enrichment was more than 20 %, or about 33 y. greater than the concentration of K in average sea water. As Saanich water has a lower chlorinity (the K/Cl ratios showed even more pronounced differences), the K/Cl ratios in deep I.W. are about 50 Oiogreater than in sea water. Similar effects have been noted in California Borderland interstitial water (BROOKSet al., 1968), but there potassium was enriched by only a maximum of 10 y. over the water column value. The sediment used for this study was kept refrigerated until just before squeezing, but may have warmed to room temperature during squeezing, thus, at least some of the K enrichment may be an artifact of the analytical technique, and the values presented represent maxima only (MANGELSDORF et al., 1969; PRESLEY, 1969). If the interstitial water has been enriched in potassium, as it appears to be, the enrichment most likely resulted from the breakdown of feldspars and other detrital silicates. Plagioclase is the most common feldspar in the Saanich sediments, and although it is not a potassium feldspar, more than enough potassium substitutes in the structure of average plagioclases to account for any enrichment of interstitial water. The magnesium-to-chloride ratios for the Saanich Inlet water differed from that of average sea water by only 1 or 2 per cent. The interstitial water ratios similarly were near normal, except for a single sample st the bottom of core 2 where an enrichment in magnesium of about 7.5 per cent occurred. This anomaly could be due to analytical error. SIEVERet al. (1965) reported that interstitial water was commonly about O-1 ppt lower in magnesium than the overlying sea water, and attributed the depletion to a probable uptake by exchange with clay minerals and/or carbonates. In studies on the California Borderland (BROOKSet al., 196S), a magnesium depletion of 5 per cent maximum was found, which may partly have been caused by 8 re-equilibration of the solid with liquid as a result of warming during squeezing. However, because sea water is 10 times richer in Mg than K on an equivalents basis, and because Mg is doubly charged, ion exchange effects on the Mg concentration would be much smaller if they occurred. The depletion of Mg, by exchange with Fe 2+ in clay lattice sites, resulting from metal sulfide deposition as a consequence of sulfate reduction (DREVER, 1971) does not seem to have occurred, probably as a result of abundant iron-rich detrital minerals. It appears that magnesium is relatively inert during early diagenesis of carbonate-poor, detritus-rich marine sediments.

ARIE NISSENBAUM, U. J.

1014

PEESLEY

and I. R.

KAPLAN

Calcium and &or&urn Samples of Saanieh Inlet water proved to be as much as 19%) lower in calcium to-chloride ratio than average sea water. However, the ratio at most sampled depths does not decrease below 15% of the sea water value. This ratio is consistently lower in the interstitial water and in two cores (Nos. 1 and 3B) it decreases to
core 1 o-15 40-50 85-100 138-150 178&185

7-6 8.0 7.9 8.0 8.0

+ 330 --120 -140 -120

2.7

tr.

13.6 23.7 29.7 36.2

3.8 5.0 0.6

7.8 7.8 7.7 7,7

-60 j-260 j-250 +260

40.6 49.2 .-. 61.2

8.0 7.6 7.6 7.5 7.7

-100 -140 -130

1.2

-

22.9 11.5 0.3 0.7 0.3

0.15 0.53 2.08 3.50 4.26

35 155 73 122

1.3

4.05 8.58 10.0 14.0

120 115 91 136

-

3.71 5-24 6.41 6.71

195 280 325 420 375

-

400 240 160 200

383 358 333 320 363 291

42

Core 2 u-15 75-85 150-165 225-235

tr. 0

tr. tr. tr.

- 100

30.2 40.6 36.2 39.1

5,3 4.4 3-1 3.1 2.x

1.4 0.3 o-3 tr. tr.

7.5 7.7 1.8 7.9

$40 + 340 +3&o + 370

49.5 66.0 52.7 44.6

1.3 0 0 0

2.0 tr.

5.0 0 0 0

7.7 7.1 7.8 7.7 7.9 7.7

-120 -110 -120 -135 -110 -110

32.4 34.6 36.8 39.7 42.4 40-l

4.1 4.1 3.8 3.8 3.4 3.1

2.2 1.9 1.2 tr. 0.2 0.2

4.26 7.92 10.4 10.6 9.8 9‘1

0 0

Core 3 O-10 5&60 10%110 15Q-160 190-200 Core 3B 790-820 1710-1740 2620-2650 3445-3475

0 0

Core 4 o-15 50-65 100-110 150-160 200-210 240-250

2.72 2.50 2.04 2.04

1015

Early diagenesisin a reducingfjord, Saanich Inlet, British Columbia-I

not surprising that the calcium concentration is so low in cores 1 and 313; rather, it is surprising that it remains so high in all the other cores. The samples showing greatest depletion in Ca have two properties in common : relatively lower phosphate content (Table 3) and higher concentration of dissolved organic matter (NISSE~BAUM et al., 1972).

Aliquots of interstitial water, left standing at 4°C in sealed glass containers, formed a precipitate after a few days. On drying and aging (for several days) this precipitate, which originally gave a poorly-defined X-ray diffractogram of calcium carbonate, was converted into calcite. It thus appears that the water in the sediment may be in equ~b~um with some metastable phase of calcium carbonate and not with calcite. This argument is further supported by the fact that the C13fC12ratio (BROWN et al., 1972) of the dissolved CO, (Table 4) is generally quite different (either Table 4. gC13and 135~ of dissolved carbonate, sulfate and sulfide in interstitial water

core#

Depth(em)

1

6-M 40-60 85-100 13&150 175-185 O-16 76-86 156-166 226-235 O-10 SO-60 10~110 160-160 UN-200 790-820 1710-1740 2626-2660 3445-3478 O-15 60-65 100-110 16~160 200-210 240-250

3B

4

6C” fCO,)%G -11.2 -37.1 -11.3 +3*3 +9*4 + 12.0 +1&o $16.7 -6+3 -0.1 +6*6 + 12.7 -j- 10.3 -i_ 17.8 $16*1 +13*3 -0.7 -+-l-6 +4-o 14.7 +5*3 -f-6,6

&P

&P’

Wd%,

G&WI&

$21.9 $39.1 -

+4*3 + 18.3 +11*6 + 14.6 -

$19.4 tii.2 -t-22.7 -

t_ 16.2 + 16.6 j-14.2 j-13.7 +12*7 +;.2 +9-1 + 13.9 $-13-l +s*a + 12.0

heavier or lighter) from the bulk sediment carbonate (which is & ly&, PDB), indicating strong ~seq~b~urn with the dominant biogenic calcite and aragonite in the sediment. Aging (and dehy~ation) within the sediment column may ‘drive’ its conversion to the more stable calcite, resulting in loss of Ca2+ from solution due to supersaturation and precipitation. As has been observed previously in interstitial water measurements (BROOKS et cd., 1968; PRESLEY and KAPLAN, 1968) strontium content closely parallels that of calcium. P~cipitation of CaCO, under these non-equi~brium con~tions scavenges strontium. It has also been found in the deep sea cores from the JOIDES project, that when the calcium content in interstitial water increases (probably by solution of carbonate) there is usually a concomitant increase in Sr.

The dissolved sulfide in Sua,nich Inlet sediment is the result of sulfate reduction b>anaerobic bacteria (see R~ICNARUS,1065 and references given therein) which fumtion after dissolved oxygen has been depleted. Table 2 shows that the water in Saanich Inlet was only slightly depleted in sulfate relative to sea water when these samples were taken, and that the SO&l ratio in the bottom water was only 5% less than that in average sea water. However, the interstitial waters were in all cases very strongly depleted, with only the surface sample of the sill core containing more than a few per cent of the original sea water Most interstitial water samples from cores 1, 3 and 4 contained concentration. measurable dissolved sulfide, as did sample 7.9-8.2 m, core 3B. In general, the pre-, sence of dissolved sulfide correlated with low Eh measured in the water.

Dissolved CO, and its U3/C12 ratios The distribution of total CO, in the water column is given in Fig. 2. The data on concentration and SC13of carbon species in the interstitial water are given in Tables 3 and 4. The CO, concentration in the water column increases from 1.5 mM/l at the surface to 2.8 m&l/l at 120 m depth. Below this, the value remains fairly constant. Total dissolved CO, in interstitial water is greater than previously reported from studies on continental shelf sediment (PRESLEY and KAPLAN, 1968) but is equivalent to values recently reported by BERNER et al. (1970). All cores (except 3B) show gradual increase of CO, with depth ; to a maximum of 66 rn~~~l at 17 m depth. The two bottom samples from core 3B show slight decrease of dissolved CO, (Table 3) presumably due to CaCO, precipitation in these samples. The 6C13 of the CO, in the water column ranges from +1x, at the surface to -l*l%,, at 100 m depth and -3+5x0 at 225 m depth, just above the sediment-water interface. The depletion in 8% is accompanied by increase in dissolved CO, and clearly points to production of CO, from biological sources in the deeper anoxic part of the basin. Values of 6C1s for the dissolved CO, in interstitial water are surprising. Instead of the CO, increase being accompanied by depletion of CY3, as found previously for interstitial water from continental shelf sediment (PRESLEY and KAPLAN, 1968), we found a reverse situation (Table 4, Fig. 3). Some of the values obtained (i.e. &Y = +17*8%,) are among the most enriched in Cl3 ever measured for dissolved CO,. As content of dissolved CO, increased, this isotope effect became more pronounced.

Several samples were measured for the CO,, methane and N,O content by gas chromatography. N,O was detected in cores 1 and 2, but was not found in core 3B. From the known concentration of CO,, and peak height comparisons, we estimate the quantities of N,O to be between 0.2 and 7 ml/l. In the deep samples from core 3B we found methane in coneentratio~s of up to 10 ml/l. This, however, probably represents minimum concentrations only, as methane would probably diffuse away from the samples during sampling and transportation. J. Cline (personal communication) described the popping of core barrel

Early diegenesis in a reducing fjord, Saanich Inlet, British ColumbiaI

1017

+40 +30 1 +20‘ii

& _CIA

NO-

se0

z m 3 a

o-

water

-_

--

;QO**

cl

d

__

*

0

0 -to-

-30 -20

0

0 0 core I A core 2 0 core 3 --core 38 * core 4

i 0

-40

0

IO

20 CO2

30

concentration

40

50

60

70

(mM/L)

Fig. 3. Conoentmtion and &Ts of dissolved CCO,.

by excess gas within short time after collection, presumably due to methane. A single specimen of methane (core 2, 225-235 cm) was measured for its KY3 and a value of -556%, was obtained. The concentration data for N,O and especially CH, are only approximate and should only be viewed as qualitative. They are given here as an indication of their presence and for the sake of interpretation of data which follows in discussion. caps

i$~lf~r species and their Ss4/S32 ratios No hydrogen sulfide was present in the water column at time of collection. Sulfur isotope measurements were made on sulfate from two water samples at station 4, the bottom sample yielded a measured value for 8S34of +20*9%,, and the surface samples, + 1 9*4x0. These data indicate that the fresh water runoff into the fjord probably contains little sulfate (as river sulfate is isotopically lighter than marine sulfate) and that the bottom water has undergone little isotopic fractionation either through reduction of sulfate or re-oxidation of sulfide. In the interstitial water, sulfate concentration decreases rapidly with depth. In all cores, excepting core 1, the top 10 cm pore water contains ten to twenty times less sulfur than the overlying water. All the cores show a decrease of sulfate with depth until almost a complete disappearance. In all cores, the dissolved H,S concentration decreases from the surface to the bottom reaching a maximum of 170 ppm at the surface of core 3. The only exception is core 1, where the zone of highest biological activity (and highest H:,S concentration) is 40-100 cm below the surface. Core 2 and the three bottom samples of core 3B did not contain any measurable sulfide, except for traces of H,S at the surface of core 2, and also show nearly complete disappearance of sulfate. Dissolved (free or unbound) H,S in these two cores has been totally removed by fixation into the sediment as iron sulfide. It is interesting to note that the sediment in core 2, where the most rapid rate of sulfate reduction occurred (as noted from disappearance of sulfate at the surface), gave a positive Eh reading. This same apparently ‘anomalous’ behavior was also found in a core from San Pedro Basin, southern California borderland (PRESLEY and KAPLAN, 1968). The value of 0.18 g/l SO, in sample 2620-2650 cm, core 3B is probably due to oxidation during squeezing.

The sulfur isotope data show that, as expected, depletion of SO, is accompanied by S31 enrichment in the residual suIfate (up to -+39.1%,). Contrary to that, found in most other marine sediments (Kfiar~a~ et el., 1963) the dissolved sulfide values arc far heavier (+4$& to -+ISy&,) than those previously found in marine sediments. The data indicate that sulfide is formed rapidly in core 4, yielding a low fractionation factor (a = SO,S == l.Oll), but much more slowly at the surface of core 1, where M = 1.035. Calculations of rates of reduction using fractionation factors have been previously made by KAPLAN et al. (1963). XiEicon Silicon in the water column shows bimodal distribution. The upper 80 m zone has a constant value of around 62-64 pgA/l (1.75 ppm Si). Below 100 m the silicon content increases to = 3.2 ppm. The distribution of silicon is remarkably close to that of phosphate (r = O-98) and is inverse to the oxygen pattern. In the pore water, silicon was measured in 4 samples only, those from core 3B (Table 3). The values of 2.1-2.7 mM/l (58-76 ppm) are similar to those reported by SIEVER et al. (1965) for interstitial water from the diatomaceous sediments from the Gulf of California, and indicate that the interstitial water is probably saturated with respect to the opal present (&EVER, 1962). Ammonia and phosphate

The phosphate concentration in the water column ranges between 1.7 and 3-O ,ugA/l (0.16 to O-28 ppm) in the upper 70 m, and between 4.4 to 4.7 ,ugA/l (0.42 to 0.45 ppm) in the 100 to 200 m zone. Ammo~a was not detected in the water column, and the major nitrogen reservoir (excepting molecular N,) was nitrate (Fig. 2). A pronounced decrease in nitrate content was found below sill depth. In the interstitial water, both phosphate and ammonia are markedly enriched (Table 3). Ammonia generally increases with depth, while phosphate has a much more erratic distribution. The concentration of dissolved ammonia in the interstitial water of Saanich Inlet is comparable with results obtained by RITTENBERO et ak. (1955) for the reducing Santa Barbara Basin, whereas the phosphate content is on the average greater than that found for Santa Barbara water. Ammonia was not found in the 3 deep samples and phosphate was depleted. This is probably due to removal of both components, the former possibly by oxidation under the high redox conditions present and the latter by precipitation as a calcium salt or co-precipitation as carbonate. pH and the buffering system In Table 3, it can be seen that the measured pH of the interstitial water has a narrow range of 7.5 to 8.0, in spite of the fact that the water has been tremendously For example, phosphate has been enriched in the products of biological activity. enriched by up to 200 times, total dissolved carbonate by up to 20 times, and ammonia by some 10,000 times over normal sea water, and hydrogen sulfide reaches concentrations of over 150 ppm. The most important process which would affect the ionic balance and hence pH,

Early diagenesisin

5 reducing

fjord, Saanich Inlet, British Columbia---I

1019

is biological sulfate reduction, which was found to occur in all cores retrieved. According to FLEMING (1940) and others, the average atomic ratio of carbon-tonitrogen-to-phosphorous in marine organic matter is 106 : 16 : 1. If organic matter can be generalized as (CH,O),,(NH,),,H,PO~, then we can represent sulfate reduction in the manner given by RICHBDS et aE.(1965) : 1/53(CHZO),,,(NH,),,H,P04

+ SOd2- = ZCO, + 2H,O + 16/53NH, + S2+ 1/53H,PO,

(1)

When the normal sea water concentration of 27 m-mole/l sulfate is completely reduced by reaction (l), 54 m-mole/l of molecular CO,, 27 m-mole/l S2-, ~8 mmole/l NH,, and ~0.5 m-mole/l H,PO, will be produced. Assuming that no CO, or NH, escapes from the sediment, and ignoring the small amount of H,PO, produced and the small amount of HCO,- originally present in the water (they approximately cancel in any case) the following set of equations (2)-( 10) can be written for mass balance and dissociation constants : CO, + HCO,- + COa2- = 0.054 M

(2)

H,S + HS- + S2- = O-027M

(3)

NH,+ + NH, = 0.008 M

(4) (5)

(H+)Wb2-)= (HCO,-)

K

t

”

(H+)(HS-) = R , PM)

&

(6) (7)

PO) This gives nine equations, for the ten unknowns (3 carbon species + 3 sulfur species + 2 nitrogen species +OH- + H+). The tenth relations~p can be contested from either a charge balance equation or a proton balance equation, showing that at equilibrium protons lost must equal protons gamed. Starting with the products given in equation (l), the proton balance is shown in equation (11). (NH4+) + 2(H,S) + (HS-) + (H+) = (OH-) + (HCO,-) + 2(COa2-)

(11)

The hydrogen ion activity can be calculated from the ten equations using a computer program to assign a number of diEerent concentrations to the starting

materials, and for different values of the dissociation constants. i\ pH of 7.0 is obtained with concentrations of 0*064 &I CO,, 0.027 M @- and O-008 M NH,, when K,’ and K,’ for H,CO, are set izt 7.4 x IF7 a*nd 4.4 x lo-11, respectively (Buch. quoted in SKIRROW, 1!)65; p. Si), K,' and K,’ for W,S are set at 14 x 10-7 anti 3.5 x lo-l4 (ELLIS and &.KI)ISG, 1959, values corrected using activity coefficients from HELCJESON, 1964). K,’ for N.H, is set at 2.5 x 1fF (HODGXAN, 1958, value corrected using activity coefficient from HARKLD and OWEN, 1967), and lo-l4 Sor K,. Actually, the calculated pH is independent of the concentration of the species as long as they are kept in the same ratio, on the condition the concentrations are high enough to ignore the initial concentration of bicarbonate, etc. prior to biogenic degradation. Although the measured pH values were between 7.6 and 8.0, it is probable that the real in sitzc values were slightly lower, as loss of CO, during sampling and squeezing was unavoidable. It thus appears that only limited reactions between the interstitial water and solid phases are necessary for pH buffering. Three solidsolution reactions may be considered : (i) removal of hydrogen sulfide as iron sulfide, largely as pyrite (ii) removal of dissolved carbonate and calcium as calcium carbonate and (iii) equilibration between silicate minerals and the dissolved cations, including H+. Although each one of these systems can result in a pH shift, their precise role cannot be readily established. For example, in the case of calcium carbonate, the interstitial water is highly supersaturated with respect to calcite, but may not be so with respect to a hydrated metastable form of calcium carbonate of unknown thermodynamic properties. PRESLEY et al. (1972) have, furthermore, shown that dissolved iron and other heavy meta cations are out of equilibrium with known metal sulfides, and here again the system could be in equilibrium with metastable and undefined iron sulfides. A similar situation exists for the case of sea water ionic exchange with silicate minerals. HELGESON et aE. (1969) give a stability field diagram showing that, with the proper choiceofformulasforthecommonsedimentarysilicates, an equilibrium pH of 75 to 8.0 should be reached by sea water. Thus, the a priori deviations suggest that the metabolic activities of microorganisms degrading the organic matter and reducing sulfate should, under ‘closed system’ conditions, lower pH of interstitial water to 7.0. Exchange reactions involving three major solid phases, will result in a buffering to maintain the pH in the range of 7.5 to 8.0. Removal of ions by diffusion or by biological activity (CO, + CH,) resulting in an ‘open system’ are discussed below. Redox p’otential Sulfide seems to be the constituent that controls the Eh of the interstitial water. Platinum electrode measurements in the interstitial water always gave negative values in the presence of dissolved sulfide, but positive values in its absence. BERNER (1963) found a direct relationship between Eh and the concentration of dissolved sulfide, but the range of sulfide concentrations in Saanich Inlet interstitial water was The measured Eh does not, however, fit too small to show any such relationship. Berner’s empirical equation of: Eh = -0.485

V -f- 0.0295 pSz-

(12)

Early diagenesisin a reducingfjord, Saanich Inlet, British Columbia-I

1021

Total dissolved sulfide in interstitial water of Saanich Inlet ranged from 06 to 5.3 x lo-* M, which is equivalent to S2- activity in the range of 5 x 1O-1o to 5 x 1O-s M and the Eh, according to the above relation, should be approximately -0*200 to -0.230 V, instead of the measured -0-100 to -0-140 V. The Eh reported by Berner was obtained by inserting the electrodes directly into the sediment, but as noted previously (BROOKSet al., 1968), this results in a lower Eh than is obtained from measurements made in the squeezed water. Dissolved sulfide completely disappeared at depth in core 2 and in core 3B, presumably by reactions that produced pyrite and other solid sulfides. Sulfide must have been present in these interstitial waters at an earlier time, because they have been strongly depleted in dissolved sulfate. The Eh was found to be highly positive in these samples, in core 3B the values were only slightly lower than those typical of oxygenated sea water. A possible explanation is that in the absence of continual generation of hydrogen sulfide, the stable metal phases of sulfide (e.g. pyrite) form, which are highly - - insoluble. Redox is then controlled by the solubility of metal oxides oxidized metal ions or carbonates, largely those of Fe and Mn, and the ratio log reduced metal ions controls the measured potential. GENERALCONSIDERATIONS Enrichment of KY3 in dissolved CO,

Although the average value for BC$_,,, of marine carbonate is about zero, isotopically light values (HATHAWAYand DEOENS,1969) and isotopically heavy values (MURATAet al., 1967) have been recorded. The former is generally thought to be formed either from bicarbonate released by metabolism of organic matter, or in isolated cases, where the carbonate is exceptionally depleted in CY3,from carbonate formed by the oxidation of methane. The explanation for isotopically heavy carbonate is not so evident. Isotopically light and heavy CO, has been found in natural gas wells (NAKAI, 1960; WASSERBUR~ et al., 1963). Dissolved CO, in the open ocean is generally in equilibrium with carbonate, hence under marine conditions 683 of the total dissolved CO, is N -lx,. During diagenesis, biogenic CO, is released in about the same isotopic ratio of the original organic matter, and as its concentration increases, the value of KY3 of the dissolved carbonate decreases until it reaches about -2O%, (PRESLEYand KAPLAN, 1968). The isotopically heavy values obtained for dissolved CO, in the present study, appear to be contrary to the observations for other marine environments. The presence of methane indicates that a relationship may exist between this gas and the dissolved CO,, causing isotopic exchange. Three possibilities are extant: (i) an exchange caused by equilibration of the two gases which could lead to fractionation factors of 1.06 to 1.07 (CRAIG, 1953; BOTTINQA,1969), (ii) formation of CH, and CO, by fermentation of organic acids by equation (13) CH,COOH --+ CH, + CO,

(13)

or, (iii) reduction of preformed carbon dioxide by methane-forming bacteria using molecular or organically available hydrogen.

We prefer the third option, There is no available evidence to demonstrate that CO, and CH, could effectively exchange in the short time available for formation oi the sediment at temperatures
(11)

Fractionation will, of course, depend on kinetic considerations. However, the data of ROSENFELD and SILVERMAN (1959) and NAKAI (1960) suggest that a kinetic isotope effect with an average value of u = 1.05 may be valid for enrichment in Cl2 in the methane relative to carbon dioxide. As stated above, carbon dioxide evolved from complex organic matter has the CY3/CY2 value of the average carbon in the organic matter. In Saanich Inlet, this value can actually be derived to be SCY3= -2O%,, from the dissolved CO, content in the upper water column (2.2 mM/l, &.Y3 = + lx,,) and in the bottom water at station 3 (2.8 mM/l, 6U3 = -3*5%,). On this basis it is possible to plot a graph representing change in isotopic ratio of CO, reduced (abscissa) according & ( 1 This has been done in Fig. 4, using a Rayleigh plot

CO, (ordinate) (14).

against percent

In R13/R,12Co, = (1 -

to equation

a) In

and assuming 01= 1.05, and the starting value for CO, is 6C13 = -20x,,. It can be seen from the data from cores 2, 3, 3B and 4, that from 25 to 50% of the original CO, formed has been reduced to methane. An anomalously light value for CO, in 40-50 cm, core 1 (%I3 = -37*1%,)

1023

Early diagenesis in a reducing fjord, Saanich Inlet, British Columbia--I

+60 +80 +100+I20 -

40

60

80

100

Percent reduction of CO2 Fig. 4. Change in 6Cls of dissolved CO, relative to the percent CO, + CH, calculated from a Reyleigh equation.

reduction

of

be due to oxidation of upward diffusing methane at the boundary of the oxidizing and reducing zones of the sediment. These data strengthen the similarity between this environment and that under which the Miocene Monterey shale is formed. Both are highly reducing, rich in diatomite, generally low carbonate and suitable for the generation of methane, resulting in formation of CY3-enrichedcarbonates (MDRATAet al., 1967). would

Nutrient regeneration during diagenesis

As stated earlier, several authors (see review by RICHARDS,1965) have stressed that the relationship of biogenically-active elements dissolved in sea water is in the approximate ratio of their abundance in marine plankton (C : N : P = 106 : 16 : 1). During degradation of organic matter under anaerobic conditions, one may expect overall reactions shown in equation (1) to follow. RICHARDSet al. (1965) have shown that in the anoxic water column of Lake Nitinat, the ratio is approximately that suggested by the theoretical relationship of (l), and there, C : S : N: P = 193 : 68 : 22 : 1. The relationship is much more complex in the interstitial water, as shown in Table 5. Furthermore, each core displays a set of characteristic values for the dissolved elements which is generally consistent for that whole core, but different from the other cores. Thus, for example, in core 1 sulfate reduction is great in comparison to CO, released and in general, the phosphate and ammonia contents are lower than expected. The phosphate content is also low in the oxidizing sediment of cores 2 and 3B. The theoretical relationship and that found in Lake Nitinat are approximated most closely by the data in reducing cores 3 and 4 and sample 7-9-8-2 m

1024

ARIE NISSENBAUM,B. J. PRESLEY and 1. I<.

KAPJAY

Table 5. Ratios of dissolved carbon dioxide. slllfide’“, ammonia and phoqh;tt,o in interstkiak T\aber 1tIcHaRos

et al. (1966)t

Theoretical Found Core #

Depth (cm)

1

o-15 40-50 86-100 135-150 175-185 O-15 75-85 150-165 225-235 O-10 50-60 100-110 150-160 19&200 790-820 1710-1740 2620-2650 3448-3475 O-15 50-68 100-I 10 150-160 200-210 240-250

2

3

3B

4

(.‘O2:s;’ (

W,:NH,

2.0

G.6

2.s

8.8

1.9 1.1 1.0 1.3 1.5 1.R 2.0

18.0 25.7 11.6 8.5 8.5

2.1 1.3 1.7 1.5 1.6 2.0 2.7 2.2 1.8 1.5 1.5 1.6 1.7 1.8 1.7

3.7 8.2 5.6 6.8 9.9 7.6 4.4 3.5 3.8 4.3 4.4

10.0 5.7

K”-:Kil, 3.3

3.1

9.3 24.2 11.7 6.8 5.1 B.7 2.8 2.4 1.7 6.2 4.6 3.8 3.6 4.8 5.2 2.8 2.2 2.3 2.4 2.5

,V,“-.I’(, L 53

GA 45 350 160 333 18T 181 206 262 170 114 82 73 58 63 60 100 150 120 58 62 70 74 67 81

NHp:l’O, 1Ii 2’

4 15 13 48 35 34 74 110 103 19 16 15 18 13 11 22 31 33 27 31

* Equal to SO,s- reduced. j’ ‘Theoretio&l’, values derived from equation (2); ‘Found’, values from mezxmmments on components dissolved in Lake Nitinst.

of core 3B. Poor agreement was also found by RITTENBERG et al. ( 1965) for NH, : PO, ratios in 3 basins from Southern California. Average ratios of 10, 20 and 200 were measured in the interstitial water from Santa Catalina, Santa Barbara and Santa Monica Basins, the latter being the most oxidizing. The relatively low CO, : S2- ratios found in all the reducing sediment supports the previous suggestion that CO, is lost from the system, probably as CH, and as CaCO,. Otherwise, one may expect to obtain C/S ratios >2, as seen in Lake Nitinat, in view of the fact that carbon dioxide can be liberated by a variety of metabolic pathways other than sulfate reduction. Ammonia appears to be enriched with respect to both dissolved CO, and PO,, except for the surface layer of the sediment cores. This again suggests loss of CO, from the system. Ammonia may be continually generated at depth, both from the breakdown of organic matter sedimenting on the floor of the fjord and from microorganisms proliferating within the sediment. It is of interest to note that whereas in cores 2, 3 and 4 the C:N ratio in interstitial water generally drops to (6, in the This may indicate that ammonia is being presediment the ratio remains at ~10. ferentially removed from the sediment to the solution. However, it should be noted that the correlation is poor between the dissolved ammonia and the total nitrogen in the sediment, at any one depth, although sediment with high organic nitrogen content, contain interstitial water with a highest dissolved NH,. We cannot explain why we

Early diagenesisin a reducingfjord, Saanich Inlet, British Cohnnbia-I

1026

were unable to detect dissolved NH, in the interstitial water from the deep sediment of core 3B. The phosphate content of the sediment does not correlate with the organic carbon content, indicating it is dominantly present as an inorganic component. Phosphate appears to be preferentially removed from the interstitial water of the oxidizing sediment and is evidently precipitating out of solution. Although we have no direct evidence, the high redox potentials in cores 2 and 3B (Table 4) may allow an iron hydroxide phase to be present in the sediment and would then permit formation of insoluble ferric phosphate or alternatively, as insoluble ammo~um phosphate salts, not Eh dependent. On the basis of a nitrogen balance, RITTENBERU et al. (1955) suggested an upward flux of ammo~a into the water column from the sediment of Santa Barbara Basin amounting to O.4o/o of the total biological nitrogen consumption in the water. Our data do not permit us to demonstrate any diffusion effects, along the sediment colmnn, In many cases, reactions have taken place so close to the sediment water interface (as in core 2,0-15 cm) that significant difFusionwould have occurred through only a very shallow path at, or near, the sediment surface. ~ckn~~~~-We wish to thank CHARI PETROWSKI,ED RUTE, YEFXOSZIUA KOLODNYand JOELCLINEfor assistance on various phases of the study. Samples were obtained through the kindness of the University of Washington, both in making available a research vessel and in assistingin the sample collection. The deep sampleswereprovided by Dr. ROBERTB~DDEXXIER. This project was supportedby AEC contract AT(~-3)-34 P.A. 134. Support for A. ~ISSEXBAUX WELEobtained from the Oceanographicand Limnologioal Company of Israel.

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