The oxygen content of ocean bottom waters, the burial efficiency of organic carbon, and the regulation of atmospheric oxygen

Palaeogeography, Palaeoclimatology, Palaeoecology (Global and Planetary Change Section), 97 (1991) 5-18

5

Elsevier Science Publishers B.V., Amsterdam

The oxygen content of ocean bottom waters, the burial efficiency of organic carbon, and the regulation of atmospheric oxygen J.N. Betts 1 and H.D. Holland Department of Earth and Planetary Sciences, Harvard Uni~,ersity, Cambridge, MA 02138, USA (Received January 23, 1991; accepted June 9, 1991)

ABSTRACT Betts, J.N. and Holland, H.D., 1991. The oxygen content of ocean bottom waters, the burial efficiency of organic carbon, and the regulation of atmospheric oxygen. Palaeogeogr., Palaeoclimatol., Palaeoecol. (Global Planet. Change Sect.), 97: 5-18. Data for the burial efficiency of organic carbon with marine sediments have been compiled for 69 locations. The burial efficiency as here defined is the ratio of the quantity of organic carbon which is ultimately buried to that which reaches the sediment-water interface. As noted previously, the sedimentation rate exerts a dominant influence on the burial efficiency. The logarithm of the burial efficiency is linearly related to the logarithm of the sedimentation rate at low sedimentation rates. At high sedimentation rates the burial efficiency can exceed 50% and becomes nearly independent of the sedimentation rate. The residual of the burial efficiency after the effect of the sedimentation rate has been subtracted is a weak function of the 0 2 concentration in bottom waters. The scatter is sufficiently large, so that the effect of the O z concentration in bottom waters on the burial efficiency of organic matter could be either negligible or a minor but significant part of the mechanism that controls the level of 0 2 in the atmosphere.

Introduction Much of the geochemisty of atmospheric O 2 is now well understood (see for instance Holland, 1978 and 1984). Atmospheric 0 2 cycles through the biosphere once in approximately 5000 years. This cycle is almost closed. However, a very small fraction, ca. 0.2%, of the organic matter generated photosynthetically escapes oxidative destruction and is buried, largely as a constituent of marine sediments. The rate at which molecular oxygen is liberated during the production of this quantity of organic matter is sufficient to double the present amount of 0 2 in the atmosphere in ca. 3 m.y. The growth of the O 2 content of the atmosphere is restrained by 0 2 use during oxidative surface weathering and by the burning of

1 Present address: Department of Earth, Atmospheric, and Planetary Sciences, Mail Stop E34-201, Massachusetts Institute of Technology, Cambridge, MA 02139, U.S.A.

reduced volcanic gases. Recent studies of the size of ancient trees and the presence of charcoal in coal (see for instance Robinson, 1989) have shown that the O e content of the atmosphere has probably been greater than ca. 13% and less than ca. 30% during the last 350 m.y. This implies that the rate of 0 2 production has been balanced quite well by the rate of O 2 loss during a period some 100 times longer than the residence time of O 2 in the atmosphere. This is surely not a coincidence. Some control mechanism must have been operating during the last several hundred million years to keep the system in balance, and to maintain the O 2 content of the atmosphere at or close to its present level. Figure 1 shows schematically the outlines of such a control mechanism. The lower curve shows a likely relationship between the O e content of the atmosphere and the rate of 0 2 loss by oxidative weathering and the burning of volcanic gases. At present, the O z output from the atmosphere is a very weak function of Po:, since the oxidation of volcanic gases and of rocks ex-

0921-8181/91/$03.50 @ 1991 - Elsevier Science Publishers B.V. All rights reserved

[~

J.N. BI2TTS AND H.D. HOLLAND

4

&

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sphere. Since then this hypothesis has been expanded (Holland, 1978, ]984), and a great deal of observational data has been obtained that bear on the burial efficiency of organic matter with marine sediments. Some of these data have been summarized recently (see for instance Calvert and Pedersen, 1991), but it seemed useful to assemble the entire database, and to assess the health of the hypothesis advanced in 1973.

-4

The burial rate of organic matter at sea "EIx[O-4

I

x

l

O

~

1.0

Po, (otto.)

Fig. I. Schematic representation of the feedback control mechanism on atmospheric O z.

posed to weathering is nearly complete (see for instance Holland, 1978, ch. 6). Hence, the effect of changes in Po2 on the O 2 output cannot presently serve as an effective control mechanism for Po2. The control must be exercised on the O 2 input side. Specifically, the burial rate of organic matter and of compounds of S 2- and Fe 2+ is probably a strong function o f / o 2 ' The curve showing the net 0 2 change in Fig. 1 is the sum of the 0 2 input and 0 2 output curves. It can be shown readily that the strength of the control on 0 2 is determined by the slope of the net 0 2 change curve where it intersects the x-axis, i.e. where d M o 2 / d t = O, and hence where the system is at steady state. The steeper the slope of the net 0 2 change curve, the more rapid the response of the system to disturbances in Poz and the smaller the change in the steady state Po2 due to shifts in the position of the 0 2 output curve. Some years ago it was suggested that the burial rate of organic matter with marine sediments might be influenced by the concentration of 0 2 in bottom waters (see for instance Demaison and Moore, 1980). Since this, in turn, is a function of the 0 2 content of the atmosphere, a possible connection was established between Po2 and the rate of 0 2 generation by the burial of organic matter at sea (Holland, 1973). The feedback by this mechanism is negative and could, therefore, serve to stabilize the 0 2 content of the atmo-

The burial rate of organic matter with marine sediments depends on the production rate of organic matter at sea, the input of terrestrial organic matter, and the destruction of both within the oceans, at the water-sediment interface, and during burial. Berger et al. (1989) have recently summarized the available data regarding ocean productivity and the loss of organic matter during its passage from near-surface environments, to marine sediments. We have collected data from the literature for 69 sites around the world that bear on the efficiency with which organic matter is sequestered. These data are summarized in Table 1. F c, the flux of carbon to the ocean floor in g-2 a-1 is best estimated from sediment trap data. However, the number of stations where sediment trap data are available are still rather few. The eighteen sites where sediment trap data are available are identified by a letter S. In the 34 sites identified by the letter P the carbon flux at the water-sediment interface was estimated from Berger et al.'s (1987) relationship between the primary productivity, PP~ and the depth, Z: Fc = 9 P P / Z + 0 . 7 P P / Z °5

(1)

For the remaining sites, identified by the letter R, F c was estimated from the sum of the organic carbon respiration rates and the burial fluxes. With one exception, this method was restricted to cores in extremely low 0 2 setting, where sulfate is the exclusive oxidant (Canfield, 1989). The exception is data point 34, for which oxidation by 0 2, NO 3, MnO2, Fe203 and SO 2 were all taken into account. Fluxes derived in this manner for other sites by the same workers agree with the values derived from sediment trap data (Bender

OXYGEN OF OCEAN B O T f O M WATERS, BURIAl. EFFICIENCY OF ORGANIC CARBON, AND OXYGEN

and Heggie, 1984). Nearly half of the bottomwater oxygen concentrations [02], mostly at open ocean sites, were derived from measurements in nearby areas, and are based on the data compiled by Mantyla and Reid (1983) and Pak et al. (1980). The uncertainties in F~, the flux of carbon to the ocean floor, S, the sedimentation rate, and BF~, the burial flux of organic carbon are summarized in Table 2. Values of the particulate flux of organic carbon to the sea floor based on sediment trap data are thought to be uncertain by a factor of ca. 1.3-5.4. Many measurements were made more than a decade ago, and hence without benefit of recent findings concerning swimmers and the use of poisons, brines, and screens (see Knauer and Asper, 1989). Cryptic swimmers (Michaels et al., 1990) have been found to complicate measurements in high-productivity, shallow or near-shore environments, where up to 96% of the organic carbon flux may be due to this source. Values of Fc based on measurements of surface productivity and the relationship in Eq. 1 between F~ and depth are estimated to have an uncertainty of ca. a factor of 2.8 (Bishop, 1989). In some instances our estimates of the burial efficiency was based on measurements of BFc and the benthic repiration rate, R:

BF~ ) 100 BE= ( BFc + R

(2)

Respirometer measurements appear to be fairly accurate. Most measure only oxygen consumption, and hence do not include the oxidation of organic carbon by sulfate reduction if the sulfur is buried in reduced formf. This is not a serious concern, since sulfide precipitated in sediments in lieu of organic carbon does not change the net oxygen production rate due to the burial of reduced species. The interpretation of respirometer measurements designed to include the loss of organic carbon due to sulfate reduction by measuring the CO 2 flux may be complicated by reactions to form CaCO 3 in sediments (Emerson and Bender, 1981). The precision of most measurements of the 0 2 concentration in bottom waters is better than 5%. However, [0 2] can decrease sharply in the viscous

7

boundary layer a few mm above the sedimentwater interface (Silverberg et al., 1987). Most measurements of bottom water [0 2] reported here are made a good deal farther from the sediment-water contact, and are often considerably greater than [O z] at the contact. The magnitude of the difference depends on the respiration rate and the velocity of bottom currents in the immediate vicinity of the contact. The burial rate of organic carbon, BFc, is the product of the sedimentation rate, S, and the carbon content of the sediments. The carbon content must be taken for sediments below the zone of active diagenesis. This will, of necessity, be a zone of relatively old sediments, which may have been deposited at a time when the flux of organic matter to the ocean floor differed considerably from the present day flux. The percent C of sediment at the sediment-water interface frequently changes significantly over rather short time intervals. Differences in the carbon content of sediments due to changes in the percent C at the sediment-water interface are often large enough to obscure the position of the base of the zone of remineralization in sediment cores. In such cases a reasonable average value of the percent C was used in calculations of the burial flux of organic carbon. The uncertainty in BF~ due to these difficulties is thought to be about a factor of 1.5 (Reimers and Suess, 1983). Whenever possible, the uncertainties assigned to the values of BE are those derived from the original sources of data. Where this was not possible, the uncertainties were based on the estimates in Table 2. These uncertainties represent approximately 1 standard deviation in the log-transformed value. Since BE = (BFJF¢) x 100, and since the major uncertainty in BFc is due to uncertainties in the sedimentation rate, S, log

BE error --- V'(log S error) 2 + (log F c error) 2

~3) when the data are log-normally distributed, and when uncertainties in S are independent of uncertainties in Fc. This is quite reasonable when Fc is derived from sediment trap data. When BE is based on respiration rate measurements, S and

Location

C e n t r a l N. P a c i f i c C e n t r a l N. Pacific C e n t r a l N. P a c i f i c C e n t r a l N. P a c i f i c C e n t r a l N. P a c i f i c C e n t r a l N. P a c i f i c C e n t r a l N. P a c i f i c E. Eq. P a c i f i c E. Eq. Pacific E. Eq. Pacific E. Eq. Pacific E. Eq. P a c i f i c Pac.-Antarct. Ridge Pac.-Antarct. Ridge Pac.-Antarct. Ridge Santa Barbara Basin S a n Nicolas B a s i n San Pedro Basin Santa Monica Basin San C l e m e n t e B a s i n Santa Catalina Basin

Peruvian Margin Peruvian Margin Peruvian Margin Peruvian Margin Peruvian Margin Peruvian Margin Hatteras A.P. Hatteras A.P. Hatteras A.P. Hatteras Cont. Rise Hatteras Cont. Rise Bermuda Rise E. E q u . A t l a n t i c Argentine Basin

No.

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21

22 23 24 25 26 27 28 29 30 31 32 33 34 35

Box C o r e No. 1 4 5A 7706-36 7706 39 7706-49 7706-50 SC-3 TC-3 PC-I 1 2BC-3 6BC 9 I,GC-2 13519 V-15-141

10132-1 10140-1 10141-1 10145-1 10147-1 10175-1 10127-2 MANOP C MANOP S MANOP M MANOP H G.R. 7812-05 7812 0 7 7812-10 239 & G1

Core/Loc.

15 ° 0 4 ' S 15 ~ 1 0 ' S 13°37'S 11 ° 1 5 ' S 11 ° 1 5 ' S 11°37'S 32 ° 4 4 ' N 32 ° 5 9 ' N 32 ° 4 6 ' N 36° 10'N 36 ° 2 5 ' N 35 ° 2 0 ' N 5 °40'N 45 ° 4 4 ' S

13 ° 4 2 ' N 9 ° 15'N 9 ° 07'N 4 o 00'N 3 °50'N 9 ° 19'N 13 ° 4 2 ' N l°30'N 11 ° 0 1 ' N 8 °48'N 6 °33'N 0 °36'N 62 ° 5 4 ' S 66 ° 5 0 ' S 62 ° 0 5 ' S 34 ° 1 4 ' N 32 ° 5 5 ' N 33 ° 3 0 ' N 33 ° 4 5 ' N 32 ° 3 5 ' N 33 ° 1 5 ' N

Latitude

75 o 3 0 ' W 75 ° 3 5 ' W 76°51'W 77 ° 5 7 ' W 79 ° 1~VW 79 ° 5 0 ' W 69 ° 5 0 ' W 70 ° 5 0 ' W 71 ° 0 0 ' W 71°24'W 71 ° 5 7 ' W 60 ° 5 0 ' W 19 ° 5 1 ' W 50 0 4 5 ' W

151 ° 3 9 ' W 148 ° 4 5 ' W 148 ° 4 7 ' W 144 ° 4 9 ' W 145 ° 0 2 ' W 146 o 0 1 ' W 151 ° 3 9 ' W 138 ° 5 7 ' W 140 ° 0 5 ' W 103° 30'W 92 ° 4 8 ' W 86°6'W 174°45'E 174 ° 1 4 ' E 168 ° 4 7 ' W 120 ° 0 2 ' W 119° 02'W 118 o 2 5 ' W 118 ° 5 0 ' W 118 ° 1 0 ' W 118 ° 3 6 ' W

Longitude

19 19 24, 27 24, 28, 35 30, 32, 35, 36 30, 32, 35, 36 18, 24 18, 24 18, 24 18, 24 18, 24 18, 24 3, 6, 24, 44 24, 27, 35

24, 26, 27, 35 24, 26, 27, 35 24, 26, 27, 35 24, 26, 27, 35 24, 26, 27 24, 26, 27, 35 24, 26, 27 3, 24, 31, 35 3, 24, 31, 35 3, 13, 24, 35 3, 13, 15, 35 10, 24 33 33 33 12, 37 5, 14, 20 5, 14, 20 5, 14, 20 4, 13, 14, 20 14, 20, 40

References

92 268 370 186 3970 4902 5330 5341 5342 4215 3955 4595 2862 5934

5004 5144 5189 4599 4619 5164 5686 4450 4915 3175 3500 2690 4139 3260 2930 575 1800 900 908 1800 1300 90 9f) 20.036 32.905 1.4714 1.0651 0.06 0.054 0.047 1.39 1.51 0.133 1.5 1.0604

0.8771 0.8632 0.8589 1.2279 1.2248 0.8613 0.5433 2.47 2.16 1.61 1.44 1.8 0.2219 0.4206 0.3201 28 14 24 20 7.9 21

Fc (g m 2 a z)

Depth (m)

S S P P P P R R R R R R R P

P P P P P P P R S S S S P P P S S S S S S

*

400 900 50 160 20 6 3 3 3 10 10 10 2 2.9

0.58 0.41 0.36 0.32 0.43 0.23 0.19 1.8 0.1 1 0.66 5 5.9 1.8 2.6 400 17 54 92 30 29

(cm/ ka)

S

3 3 9 9 161 161 273 273 273 277 277 273 255 232

179 179 179 170 170 179 188 167 167 110 110 125 224 224 224 5 22 7 5 55 15

(/zmol /kg)

[O 2]

40 70 38 48 1.49 0.16 0.004 0.019 0.017 0.23 0.19 0.019 0.05 0.14

0.0068 0.0051 0.0042 0.0036 0.0052 0.0037 0.002 0.024 0.0012 0.02 0.0067 0.22 0.031 0.021 0.029 24 8.4 14 8.8 3 11

(gm 2 a i)

BF c

44 78 190 146 101 15 7 35 36 17 13 14 3.3 13

0.8 0.6 0.5 0.3 0.4 0.4 0.4 1.0 0.06 1.2 0.5 12 14 5 9 86 60 58 44 38 52

(%)

BE

100

330 330 110 90

17 28 20

75 75 75 100 100 75 50

(gC m a 1)

PP

2

S u m m a r y of a v a i l a b l e d a t a f o r c a l c u l a t i n g the b u r i a l e f f i c i e n c y o f o r g a n i c c a r b o n in m a r i n e s e d i m e n t s (/-[. - c a r b o n r a i n to b e n t h i c i n t e r f a c e : S - s e d i m e n t a t i o n r a t e ; [O2], b o t t o m - w a t e r o x y g e n c o n c e n t r a t i o n ; BF~. = b u r i a l flux o f o r g a n i c c a r b o n in s e d i m e n t : B E - b u r i a l efficiency; P P = p r i m a r y p r o d u c t i v i t y . )

TABLE 1

Argentine Basin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Margin N.W. African Rise N.W. African Rise N.W. African Slope W. Baltic: Kiel B i g h t Baltic: G o t l a n d D e e p Baltic: G o t l a n d D e e p Baltic: B o r n h o l m D e e p Baltic: B o r n h o l m D e e p Baltic: B o r n h o l m D e e p W. B l a c k S e a W. B l a c k S e a W. B l a c k S e a W. B l a c k S e a Black Sea Black Sea Cape Lookout Bight Oregon Margin Limfjorden, Denmark Framvaren, Norway Lawrentian Trough Saanich Inlet Long Island Sound

7 Core Mean BH

V-15-142 123210-3 12327-4 12328-4 12329-4 12336-1 12337-4 12344-3 123454 12347-1 12392-1 13209-2 K N 58 DOS 2 D W D 106 17-23 m ~i30 031 Mud Basins B.D. Mean ~ 12 571 580 584 601 1432 1470 A-1 7610-08 1-8 M e a n Dccp Basin

44 ° 5 4 ' S 23 ° 3 0 ' N 23 ° 0 8 ' N 21 ° 0 9 ' N 19 ° 2 2 ' N 16 o 1 4 ' N 15 ° 5 8 ' N 15 ° 2 6 ' N 15 ° 2 9 ' N 15°50'N 25 ° 1 0 ' N 12 o 2 9 ' N 38 ° 2 8 ' N 38 ° 1 9 ' N 38 ° 5 0 ' N 54 ° 3 2 ' N 57 o 0 0 ' N 57 ° 2 0 ' N 55 o 1 8 ' N 55 o 1 3 ' N 55 ° 2 1 ' N 43 ° 2 0 ' N 43 ° 1 0 ' N 42 ° 5 0 ' N 42 ° 3 0 ' N 43 o 0 1 ' N 42 o 0 3 ' N 34°37'N 44 ° 3 6 ' N 56 ° 4 7 ' N 58 * 1 0 ' N 49 ° 4 0 ' N 48°36'N 41 o 1 6 ' N

51°32'W 18 ° 4 3 ' W 17 ° 4 4 ' W 18 ° 3 4 ' W 19°56'W 20 ° 2 6 ' W 18 ° 0 7 ' W 17 ° 2 1 ' W 17 ° 2 2 ' W 17°51'W 16 ° 5 1 ' W 20 o 0 3 ' W 72 ° 0 2 ' W 69 ~ 3 7 ' W 72 ° 3 1 ' W 10 ° 0 3 ' E 19 o 2 0 ' E 20 ° 0 3 ' E 15 o 0 4 ' E 15 o 0 5 ' E 15 o 1 7 ' E 29 ° 5 8 ' E 28 ° 4 8 ' E 29 ° 3 5 ' E 30 ° 3 5 ' E 34 ° 0 5 ' E 41 o 1 8 ' E 76 ° 3 3 ' W 126 ° 2 0 ' W 9 ° 15'E 6 °45'E 68 ° 3 0 ' W 123°30'W 72 o 5 4 ' W

24, 17, 27, 17, 17, 17, 17, 17, 17, 17, 17, 24, 23, 23, 23, 1 8, 8, 8, 8, 8, 9, 9, 9, 9, 16 16 2, 24, 21 29, 38 11, 7, 35 35 35 35 35 35 35 35, 44 35 41 41 41

35 35

28 43

39

25 27

22, 23 22, 23 22 22 22 42 42 42 42

27, 27, 35 27, 27, 27, 27, 27, 27, 27, 27, 27, 34, 34, 34,

5885 3076 2037 2798 3315 3645 3085 711 966 2710 2575 4713 2816 3577 2192 18 165 230 70 70 90 1420 320 2080 1000 2248 906 10 2060 10 160 350 225 3

1.0654 1.3992 2.5906 3.7835 1.3385 1.0548 2.4832 8.1712 6.6862 2.9343 1.2967 1.029 2.3 4.2 6.3 40 6.5424 5.1172 48.03 48.03 84 3.3 13.2 8.7 2.9 1.6891 8.9612 1800 1.9792 170 22 72 121 670

P P P P P P P P P P P P S S S S P P P P R R R R R P P R P R S S R R

3.3 3.8 8.3 15 2.6 2.2 4.9 16.5 16.5 12.5 5.1 3.2 4.5 8 16 100 17 23 100 40 275 1.7 27 5 5 22 40 10000 10 200 200 610 1050 5000

232 230 160 220 230 230 230 100 120 220 220 259 277 277 277 317 0 0 25 25 25 0 0 0 0 0 0 200 125 200 0 320 0 0

0.2 0.084 0.45 1.4 0.097 0.066 0.41 2.4 3.4 1.66 I).17 0.093 0.45 0.8 1.6 7 2.6 3.2 15 6 58 0.42 6.8 2.7 1.4 1.3 2.4 1400 0.44 20 20 25 48 360

19 6 17 37 7 6 17 29 51 57 13 9 20 19 25 18 40 63 31 12 69 13 52 31 48 77 27 78 22 12 91 35 40 54 100

90 270

60 60 90 90 90

100 90 130 230 90 75 160 210 210 175 75 85

1. B a l z c r c t a l . , 1986. 2. B a r t l c t , K., 1990 (pers. c o m m ) 3. B e n d e r a n d t l e g g i e , 1984. 4. B e n d e r et al., 1989. 5. B e r e l s o n et al., 1987. 6. B e r g e r et al., 1989. 7. B e r n e r a n d W e s t r i c h , 1985. 8. B o e s e n a n d P o s t m a , 1988. 9. C a n f i e l d , 1989. 10. C o b l e r a n d D y m o n d , 1980. 11. D e v o l el al., 1984. 12. D y m o n d et al., 1981. 13. E m e r s o n , 1985. 14. E m e r y a n d Bray, 1962. 15. F i n n e y et al., 1988. 16. G l e n n a n d A r t h u r , 1985. 17. H a r t m a n et al., 1976. 18. H e g g i e et al., 1987. 19. H e n r i c h s a n d F a r r i n g t o n , 1984. 20. J a h n k c , 1990. 21. J o r g e n s e n , 1 9 7 7 . 2 2 . K 6 g l e r a n d L a r s e n , 1978. 23. M a n h e i m , 1961. 24. M a n t y l a a n d R e i d , 1 9 8 3 . 2 5 . M a r t e n s a n d K l u m p , 1984. 26. Miiller a n d M a n g i n i , 1980. 27. M f i l l e r a n d Suess, 1979. 28. M u r r a y et al., 1978. 29. N2es et al., 1988. 30. P a k et a L 1 9 8 0 . 3 1 . R e i m e r s et al.. 1984. 32. R e i m e r s a n d Suess, 1981. 33. R e i m e r s a n d Suess, 1983. 34. R o w e a n d G a r d n e r , 1 9 7 9 . 3 5 . S a r n t h e i n et al., 1987. 36. S c h e i d e g g c r a n d K r i s s c k , 1 9 8 1 . 3 7 . S h o l k o v i t z , 1973. 38. S i l v e r b e r g et al., 1987. 39. Skei et al., 1 9 8 8 . 4 0 . S m i t h et al., 1987. 41. T u r e k i a n , 1965. 42. V a i n s b t e i n et al., 1985. 43. W e s t r i c h , 1983. 44. Z a h n et al., 1986. * P: F c values d e r i v e d f r o m p r i m a r y p r o d u c t i o n a n d d e p t h . S: F c values e s t i m a t e d f r o m s e d i m e n t t r a p d a t a . R: F~ values c a l c u l a t e d as t h e s u m o f r e s p i r a t i o n a n d b u r i a l flux.

45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69

44

36 37 38 39 40 41 42 43

o

Z

©

Z

8 z

>

o © >

Z (3 ,.<.

t~

C

),

o

Z

o ©

o

X .<

10

J N . B E T T S A N D H.D. H O L L A N D

TABLE 2 y = 1.5084 - 2.5971e-3x

R~2 = 0.136

Uncertainties (1 o-) in the parameters used in estimating BE, the organic carbon burial efficiency

R: Benthic respiration rate Respirometer: factor of 1.5 (Reimers and Suess, 1983) Pore-water model: factor of 2 Solid-phase model: factor of 4 F~.: Particulate organic carbon flux to the sea floor Sediment traps: factor of 1.5-4 or more, depending on water depth (E. Druffel, pers. comm, and Michaels et al., 1990) Calc. from PP, Z: factor of 2.8 (Bishop, 1989)

--0

BFc: Net burial flux of organic carbon in sediments S and %C: factor of 1.5 (Reimers and Suess, 1983)

_1 -2

Fc are not entirely independent, and Eqn. 3 may overestimate the error in BE. The calculated values of BE at three sites are greater than 100%. This might be due in part to the movement of organic matter along the ocean floor or to underestimates of F c. However, the differences between the calculated values of BE and 100% lie within 0.05 and 0.8 standard deviations. Since 21 data sets yielded BE values within one standard deviation of 100%, it is not surprising that three fall above 100%. The three anomalous data points were not eliminated from the entire data set to avoid adding bias to the data.

t , ;0

2;°

400

Fig. 2. Plot of the burial efficiency, BE, vs. the oxygen concentration in bottom waters.

and it is most likely that the sedimentation rate is the most important parameter controlling the burial efficiency of organic matter with marine sediments. This is not a new discovery. Toth and Lerman (1977), Heath et al. (1977) and Mfiller and Suess (1979) observed this relationship more than a decade ago. Since then their observations

Parameters that influence the burial efficiency of organic matter at sea

Parameters that could influence the burial efficiency of organic matter with marine sediments include the 02 concentration of bottom water, the flux of organic matter to the sea floor, and the sedimentation rate. Figure 2 is a plot of the values of BE in Table 1 plotted against the 0 2 concentration in bottom water. The scatter is very large, and there is no obvious correlation between BE and [O2]. Figure 3 is a plot of the burial efficiency, BE, versus the benthic carbon rain rate, F c. There is, again, a large degree of scatter and no indication of a close connection between the two parameters. Figure 4 is a plot of the burial efficiency, BE, versus the sedimentation rate, S. The correlation is very satisfactory,

300

[o~1 gM

1

It t

ttttI

t

0

-2 Iog(Fcl

gC/m2/ll

Fig. 3. Plot of the burial efficiency, BE, vs. the benthic organic carbon rain rate, F~.

OXYGEN OF OCEAN BOTTOM WATERS, BURIAL EFFICIENCY OF ORGANIC CARBON, AND OXYGEN

11

data in Fig. 4 can be fitted adequately by the equation A log(S) log(BE) 1

m v

0

--~

f(S)

-1. o

•

-2 -2

,

.

-1

=

.

0

, 1 log(S)

•

, 2

•

log(BE)

. 3

,

-

4

¢mlka

Fig. 4. Plot of the burial efficiency, BE, vs. sedimentation rate, S.

have been confirmed by Sarnthein et al. (1987), Henrichs and Reeburgh (1987), and Canfield (1989). Roughly a third of the data used in our study were taken, at least in part, from the paper by Sarnthein et al. (1987). In their data set the burial flux, BFc, was related to the rain rate F c and the sedimentation rate Sb_ ~ by the relationship

BF~. = 0 .0168FcSb~ l¢,S~

(4)

where the sedimentation rate is in c m / k a and excludes the contribution from organic carbon. In this nomenclature, and adding the contribution of organic carbon to the sedimentation rate, the burial efficiency BE becomes

BE = 1.68S I°4

(5)

A line generated by Eqn. 5 fits the data in Fig. 4 reasonably well except at sedimentation rates in excess of ca. 60 cm/ka, where it predicts burial efficiencies in excess of !00%. Sarnthein et al.'s (1987) data did not extend into the high-sedimentation range region, where Eqns. 4 and 5 must break down. Using non-linear statistics on the log-transformed data, we have shown that the

log( S + B)

+ C =f(S)

(6)

where: A = 1.39; B = 7.90 cm/ka; C = 0.34. At low sedimentation rates this reduces to a linear relationship between log BE and log S. At high sedimentation rates BE approaches 10 173 , i.e. 54%. The residuals [log(BE)-f(S)] represent the variability in BE which is not explained by f(S), and it is of interest to see whether these are a function of the 0 2 concentration in bottom waters. Figure 5a is a plot of all the residuals versus [O2]. The scatter is large, and there is no obvious correlation between the residuals and [O2]. As expected, the limits on the upper and lower slope of the regression line at the 95% confidence limits are small: - 6 x 10 s /zM-~ arid - 1.2 x 10 -3 /zM -I. These limits derive solely from the uncertainty due to the scatter in the data. They do not incorporate the uncertainty estimates for the measurements themselves. It is worth pointing out, however, that the deviations from the regression line are roughly comparable to the uncertainty estimates. The inclusion of all the residuals in Fig. 5a could be misleading if the effect of the 0 2 concentration in bottom waters on the preservation of organic matter is a function of the sedimentation rate. To test this possibility we have repeated the analysis after deleting the data for stations where the sedimentation rate was less than 10 cm/ka. The residuals from a linear regression, g(S), on data for the high-S stations are plotted in Fig. 5b. The limits on the upper and lower slope at the 95% confidence limits are - 5 x 10 4 ~ M -1 and - 2 . 0 x 1 0 -s /xM -1. This range of slopes is larger than that for the entire data set, but not more so than might be expected solely from a reduction in the number of data point from 69 to 35. We see no evidence, therefore, for a significant effect of the sedimentation rate on the effect of [0 2] on BE. Sarnthein et al. (1987) attempted to demonstrate that the effect of [0 2] on Fc is negligible by plotting Fc versus [0 2] and superimposing iso-

12

J.N. B E T T S A N D H.D. H O L L A N D

2' 95% upper slope:

(a)

y=-O.OOOO6x+O.OO9

A t~ m

=

-1'

95% lower slope: y=-O.OO13x+O.lg .2 ¸ 1 O0

200 [021

300

~M

(b) 95% upper slope: y='O'OOO47x+O'O5g

i I,M

0= g5% lower slope:

Limits on the role of atmospheric oxygen on the burial efficiency of organic carbon

The previous section showed that the concentration of 0 2 in bottom waters exercises no more than a minor effect on the burial efficiency of organic carbon with marine sediments. It remains to be seen whether this small effect, might still play a significant role in maintaining the redox balance of the atmosphere-biosphere-oceancrust system. To test the potential importance of the effect, we propose to consider the response of the 0 2 concentration of bottom waters and the 0 2 pressure in the atmosphere to changes in the rate of 0 2 use in the exogenic cycle. The 0 2 concentration, ~moz, of surface sea water is close to equilibrium with the partial pressure of 0 2 in the atmosphere. It is, therefore, proportional to Po, and to the total number of moles of 0 2 in the atmosphere-ocean system. Since the number of moles of 0 2 in the oceans is less than 1% of the number of moles of 0 2 in the atmosphere, ~mo' = Bo2Po2 - 6.8 x 10-ISMo,tzM/kg

(7)

y=-O,OO20x+O.25

-2 100

200

300

[o21 ~M Fig. 5. Plot of the residuals vs. oxygen content of bottom water for (a) all sites and (b) sites with S > 10 c m / k a . The residuals are from a non-linear fit in (a) and a linear fit in (b).

pleths of BFc. Their plot contains too few data points to constrain the slope of the isopleths in a conclusive fashion. Henrichs and Reeburg (1987) plotted BE versus S, and found that the low-[O 2] sites did not clump at the high-BE edge of their curve. Their approach led to the same conclusion as ours and Berner and Canfield's (1989): the effect of the O~ concentration in bottom waters therefore exerts at most a minor effect on the burial efficiency of organic matter in marine sediments. This conclusion is at odds with Emerson's (1985); his model was, however, based only on data from a small number of locations.

where: smo2 = concentration of 0 2 in average, present day, surface sea water, B()~= Henry's Law constant for 0 2 in average sea water; and Mob=number of moles of 0 2 in the atmosphere-ocean system; its present value is 3.7 × 10 lo Except in near-shore settings, the concentration of dissolved oxygen at the sediment-water interface is smaller than ~mo2. The value for average bottom water appears to be about 100 ~ M / k g lower than that of average surface water (Broecker and Peng, 1982); thus [02] ~ 6.8 X 1 0 - ' a M o ~ - 100 # M / k g

(8)

This expression is, of course, invalid when the concentration of 0 2 is surface waters is less that 100 ~ M / k g . From the data in Fig. 5a it follows that BE BE o

- 10 -"[°21+b

(9)

13

()XY(;EN ()F"OCEAN BO'ITOM WATERS, BURIAl. EFFICIENCY OF ORGANIC CARBON. AND OX~((;EN

9.0

where BE o is the burial efficiency of organic carbon when [02] = b/a = 150 txM/kg

(10)

7.0

"6

BE

-- 10

1.3)~1(1 3 [ 0 2 ] + 0 . 1 9 = e

3.llXlO 3102]+1).44

BE,, (11) The exogenic redox cycle is very nearly balanced today (see for instance Holland, 1978, ch. 6). fir the 0 2 use rate due to the oxidation of rocks exposed to weathering and the oxidation of volcanic gases, is ca. 1.0 × 1013 tool/yr. The generation rate of 0 2 by the burial of reduced substances, largely organic carbon, with marine sediments, is also close to 1.0 × 1013 mol/yr. Hence, today dM%

( F~)( BE)A

dt

12

=0

(12)

Where: ( F c ) ( B E ) = rain rate of organic matter to the ocean floor in g / m -2 a -t times the efficiency of organic matter burial with sediments, averaged over the ocean floor; and A = area of the oceans in m 2 (13) If we reduce the value of ~ , the 0 2 balance will be upset, and Mo2 will rise. If we assume that the redox balance is reestablished solely by the effect of the increase of [0 2] on BE, we can calculate the necessary change in Mo2 and the time required to reach the new steady state. From Eqns. 8, 9, and 11 it follows that dt

- 1.0 × 1013 = 2.10 X 1013 e

- ~ 2.04×10 mMo2 - -

(14) I//"

(15)

When steady-state has been reestablished, Mo_' = - 4 . 9 0 × 1019 In

x 10 13 2.10

(16)

/

s.0

o,o

minimum slo~

~4t~ Of BE vs" [021

5.0

where a and b are the slope and intercept, respectively, of the regression lines in Fig. 5a. Since we wish first to establish the maximum likely influence of [0 2] on BE and hence on the exogenic redox cycle, we will use the most negative likely value of a. Thus

0.50

'\

8.0

4.0

0.30

maximum rS~Opea , a ~ BE VS. [0 t] ~

0.20 g

of

5.0 2.0

~

0.10

1.0

O25

0.50

0.75 1.00 1.25 ~0 X 1013mollyear

1.50

1.75

Fig. 6. Relationship between Mo,, Po, and at stead3, state fur values of between 0.40× 10 ~3 and 1.5(f× 1013 tool/yr. For the "maximum slope" curve a = - 1.3× l0 3: for the "minimum slope': curve a = + 6 × 111 5.

Figure 6 shows the relationship between the steady state values of Mo~ and ~ for values of between 0.40 × 1013 and 1.50 × 1013 tool/yr. An increase by 50% in the rate of 0 2 consumption decreases the steady state value of Po~ by a factor of 2. A decrease by a factor of 2 in the rate of 0 2 consumption increases the steady state value of Po2 by a factor of 2. If we choose a and b for the line of minimum slope relating BE to [0 2] in Fig. 5a, the required change in Mo~ and Po,_ to compensate for even very small changes in is extremely large. The approach to a new steady state after a sudden change in ~V from 1.00× 1013 mol/yr at t = 0 can be calculated readily by integrating Eqn. 13 to yield: ( 3.7 × 1019 - Mo: )

t=

7

4.90 × 10'~'

,

2.10 × 1013 e - 2ll4X''1 -""%e- q,

Xln

1.00

X](] ~ ~-

)

(17)

where ~ is the rate of 0 2 use after time t = 0. Figure 7 shows the change in Mo, and f))~ after a decrease of ~ from 1.0 X 1013 -mol/yr to 0.50 X 1013 mol/yr and from 1.0x 1013 mol/yr to 0.80 x 1013 mol/yr. The e-folding time in the first case is 8.2 m.y.; in the second case the e-folding time is 5.7 m.y. These intervals are ca. 3 times and twice the residence time of 0 2 in the atmosphere-ocean system. Similar curves can be constructed for the response of Mo2 and Po: if the effect of [0 2] on BE is described by the curve of

14

J.N. BE'ITS AND H.D. HOLLAND 8.0

0.45 I

J

~ = 0.50 x 1013tool/year

! 0,40

7° I

,oi ~:~x

j

0.35

///

5.0

0.30

~/ = 0 . 8 0 x 1013 rnol/year

o.. i

4.0/~

0.25

0.20 30 0

1 5

• 10

~ 15

I 20

25

30

35

t X 106 Years

Fig. 7. Increase of Mo, and Po, with time after reduction of from 1.0(l× 10 I3 mol/yr to 0.5()× 101) mol/yr (upper curve) and (/.80× 10I3 mol/yr (lower curve).

minimum slope in Fig. 5. E-folding times of the response to changes in the value of qt are then on the order of 200 m.y. It is very likely that the rate of 0 2 consumption in the exogenic cycle has varied by more than a factor two during the past 350 m.y. (Kump and Garrels, 1986; Berner and Canfield, 1989). The above calculations show that Mo2 and Po, should then have varied by more than a factor of 2 during this period, even if the control which the 0 2 content of bottom waters exerts on the burial efficiency of organic matter with marine sediments is assumed to be the maximum allowed by the data in Fig. 5a. This range probably exceeds the limits of Po, allowed by the presence of charcoal and the remains of large trees in the coal deposited during the past 350 m.y. (Robinson, 1989). If so, we have shown that the control that [0 2] exercises on BE and hence on Mo, and Po, may not be negligible, but that it is unlikely to have been a major part of the mechanism that has controlled atmospheric 0 2 during the past 350 m.y. Components for a model of the 0 2 control system If the relationship between BE and [0 2] is unlikely to have been the major control mechanism of atmospheric 0 2 during the past 350 m.y., we must look to the connection between Fc and atmospheric oxygen. There are some hints regarding this connection, but the matter is compli-

cated by several other parameters that do or can influence F~ and hence dMojdt. F c depends in large part on the biologic productivity of the oceans, the depth of seawater below the areas of greatest productivity, and the contribution of terrestrial organic matter to Fc. The productivity of the oceans depends in large part on the availability of nutrients. Among these, PO~- and NO 3 are the most important, although iron and perhaps other metals are apparently in sufficiently short supply to limit productivity in some parts of the oceans. PO2 is supplied almost entirely from the continents, and it is likely that during periods of low erosion and weathering rates the rate of 0 2 use, q~, the sedimentation rate, S, and the biological productivity of the oceans are all lower than during periods of rapid erosion and weathering. If so, Fc may well be positively correlated with ~, and the changes in Mo, required to maintain the 0 2 balance of the atmosphere may well be much smaller than the computed changes in Fig. 6. As shown above, the burial efficiency of organic matter decreases with decreasing sedimentation rate. At a constant flux of organic matter the lower value of BE during periods of low erosion rates would therefore tend to offset the lower value of qr during these periods. However, the average rate of sedimentation of the bulk of marine sediments is so high, that the effect of a decrease in the average value of S on the average value of BE is probably too small to balance more than a small fraction of the decrease in the 0 2 use rate, ~. Changes in the rate of chemical weathering can be occasioned by changes in tectonic activity, in the average composition of surface rocks, and in the quantity and composition of volcanic gases. As shown in Fig. 8, the H 2 / C O 2 mol ratio in volcanic gases is quite variable. This variability is due in part to differences in the oxidation state of magmas, in part to differences in their (H 2 + H 2 0 / C O + CO 2) ratio. During periods of volcanism when the flux of volcanic gases is normal but the mn2/rnco, ratio is abnormally high, q~ will be abnormally high, whereas the nutrient flux supplied by chemical weathering will be normal. This may explain (Holland, 1989) the unusual

OXYGEN OF OCEAN BOI~FOM WATERS, BURIAL EFFICIENC'¢ OF ORGANIC CARBON, AND OXYGEN Probable Mean Range Tolbochik (Kamchatka), (1020=)

5.7

While Island (,580-650") e Mr. St. Helens ()400 °)

Nyiro~ngo (970-1020 °) Mt Etno (1075=) Ardoukoba (1070") Erta Ale (1075-1210°) SurtNy (1125=] Pu'u 0'o ()1000~

, Kilauaa (1100-1175=) I 0

e -

0

l}

S~wa

Shinzon

-

Iwo Jima (659 =) Usu (600-7000) Merapi (720-819 °)

0.2

0.4

0.6

0.8

1.0

1.2

1.4

1.6

1.8

2.0

22

mH~ [mol/mol)

moo2 Fig. 8. The ratio of the concentration of H 2 - C O 2 in volcanic gases.

isotopic composition of carbon in sedimentary, rocks formed between 850 and 600 m.y.B.P. (Knoll et al., 1986). The relationship of Fc and ~ is clearly complex and not very well understood. This is also true for the relationship of Fc and Mo2. The O 2 content of seawater does not seem to exercise a large effect on the rate of oxidation of organic matter in seawater (Devol, 1978). However, the 02 content of seawater exerts a major influence on the NO 3 budget of the oceans. As Codispoti (1989) has pointed out, the major loss of NO B from the oceans occurs in O2-minimum regions, where the concentration of 0 2 is reduced to values close to zero by the oxidation of organic matter. NO~ is used as a source of oxygen in this region, and the rate of NO 3 loss is very significant for the oceans as a whole. An increase in Mo2 and hence in the concentration of 0 2 in surface ocean water serves to decrease the rate of NO~ loss and tends to increase the productivity of the oceans. This, in turn, increases the 0 2 demand in O2-minimum regions, but it seems likely that the overall effect of the coupling of NO~ and atmospheric 0 2 produces a net positive feedback on F c. The interaction of O 2 with the PO 3- cycle of the oceans almost certainly produces a negative feedback on F c. The major sinks of PO43 in the modern ocean are organic matter, ferric hydrox-

15

ide-ferric phosphate solid solutions, and apatite (see for instance Ruttenberg, 1990). An increase in the 0 2 content of bottom waters probably increases the quantity of iron sedimented as Fe(OH) 3 and not reduced by organic matter, and hence increase the rate of removal of PO~ from the oceans. This, in turn, decreases the availability of PO~ for removal with organic matter and hence decreases the average value of Fc. The connection between [O2], the rate of phosphate removal with Fe(OH) 3, and Fc has not been quantified. It may turn out to be the most important missing link in the chain that connects the 0 2 content of the atmosphere to the redox balance of the exogenic cycle. Another possible role of phosphorus as a link in the 0 2 feedback system involves changes in the release rate of phosphorus to the oceans from land as a result of forest fires whose frequency is affected by the 0 2 content of the atmosphere (Kump, 1988).

Conclusions (1) The new compilation of data for the burial efficiency, BE, of organic matter with marine sediments has confirmed that the sedimentation rate, S, is the dominant influence on BE. (2) The residuals after the subtraction of the influence of S on BE show a slight negative correlation between BE and the 0 2 concentration in bottom waters. (3) The scatter in these residuals is so large, that the effect of [0 2] on BE could be either negligible or a fairly significant part of the mechanism that controls the level of atmospheric 0 2. (4) The remainder of the O 2 control mechanism almost certainly involves correlated changes in Fc and the rate of 0 2 loss, q'. The removal of PO~- with Fe(OH) 3 in marine sediments is a function of the O 2 content of seawater and may turn out to be an important part of the mechanism that controls the O2 content of the atmosphere.

Acknowledgments The authors wish to thank R.A. Jahnke, T.F. Pederson, K. Winn, F. Sayles, S.M. Henrichs, E.

16

Druffel, and A.W. Mantyla for helpful discussions and data, R.A. Berner and M.A. Arthur for most helpful reviews, and NASA for financial support under grant NAGW-599 to Harvard University. References Balzer, F,, Pollehne, F. and Erlenkeuser, H., 1986, Cycling of organic carbon in a coastal marine system. In: P.G. Sly (Editor), Sediments and Water Interactions. Springer, New York, pp. 325-33/). Bender, M.L. and Heggie, D.T., 1984. Fate of organic carbon reaching the deep sea floor: a status report. Geochim. Cosmochim Acta, 48: 977-986. Bender, M., Janke, R. and Weiss, R., 1989. Organic carbon oxidation and benthic nitrogen and silica dynamics in San Clemente Basin, a continental borderland site. Geochim. Cosmochim. Acta, 53: 685-697. Berelson, W.M., Hammond, D.E. and Johnson, K.S., 1987. Benthic fluxes and the cycling of biogenic silica and carbon in two southern California borderland basins. Geochim. Cosmochim. Acta, 51: 1345-1363. Berger, W.H., Fisher, K., Lai. C. and Wu, G,, 1987. Ocean carbon flux: global maps of primary production and export production. In: C, Agegian (Editor), Biogeochemical Cycling and Fluxes between the Deep Euphotic Zone and Other Oceanic Realms. NOAA Syrup. Ser. for Undersea Research, NOAA Undersea Research Program, vol. 3(2). Berger, W.H,. Smetacek, V.S. and Wefer, G., 1989. Productivity of the Oceans: Present and Past. Wiley Interscience, New York, 471 pp. Berner. R.A. and Canfield. D.E., 1989. A new model for atmospheric oxygen over Phanerozoic time. Am. J. Sci., 289: 333-361. Berner, R.A. and Westrich. J.T., 1985. Bioturbation and the early diagenesis of carbon and sulfur. Am. L Sci. 285: 193-206. Bishop, J.K.B., 1989. Regional extremes in particulate matter composition and flux: effects on the chemistry of the ocean interior. In: W.H. Berger, V.S. Smetacek and G. Wefer (Editors), Productivity of the Ocean: Present and Pasl. Wiley lnterscience, New York, pp. 117-137. Boesen, C. and Postma, D., 1988. Pyrite formation in anoxic environments of the Baltic. Am. J. Sci., 288: 575-603. Broecker, W.S. and Peng, T.-H., 1982. Tracers in the Sea. Eldigio, New York. Calvert, S.E. and Pederson, T.F., 1991. Organic carbon accumulation and preservation in marine sediments: how important is anoxia? In: J. Whalen and J. Farrington (Editors), Production, Accumulation and Preservation of Organic Matter in Recent and Ancient Sediments. Canfield. D.E., 1989. Sulfate reduction and oxic respiration in marine sediments: implications for organic carbon preservation in euxinic environments. Deep-sea Res., 36(11: 121138.

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OXYGEN OF OCEAN BO]q~OM WATERS. BURIAL EFFICIENCY OF ORGANIC CARBON, AND OXYGEN

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17

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