36. minor element and stable isotopic ......minor element and stable isotopic composition of the...

Duncan, R. A., Backman, J., Peterson, L. C , et al., 1990Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 115

36. MINOR ELEMENT AND STABLE ISOTOPIC COMPOSITIONOF THE CARBONATE FINE FRACTION: SITE 709, INDIAN OCEAN

Paul A. Baker,2 Mitchell J. Malone,2 Stephen J. Burns,3 and Peter K. Swart3

ABSTRACT

Stable isotopic and minor element compositions were measured on the fine fraction of pelagic carbonate sedimentsfrom Ocean Drilling Program Site 709 in the central Indian Ocean. This section ranges in age from 47 Ma to thepresent. The observed compositional variations are the result of either paleoceanographic changes (past oceanic chemi-cal or temperature variations) or diagenetic changes.

The CaCO3 record is little affected by diagenesis. From previous work, carbonate content is known to be determinedby the interplay of biological productivity, water column dissolution, and dilution. The carbon isotopic record is gener-ally similar to previously published curves. A good correlation was observed between sea-level high stands and high13C/12C ratios. This supports Shackleton's hypothesis that as the proportion of organic carbon buried in marine sedi-ments becomes larger, oceanic-dissolved inorganic carbon becomes isotopically heavier. This proportion appears to behigher when sea level is higher and organic carbon is buried in more extensive shallow-shelf sediments.

The strontium content and oxygen isotopic composition of carbonate sediments are much more affected by burialdiagenesis. Low strontium concentrations are invariably associated with high values of S18O, probably indicating zonesof greater carbonate recrystallization. Nevertheless, there is an inverse correlation between strontium concentration andsea level that is thought to be a result of high-strontium aragonitic sedimentation on shallow banks and shelves duringhigh stands.

Iron and manganese concentrations and, to a lesser extent, magnesium and strontium concentrations and carbonisotopic ratios are affected by early diagenetic reactions. These reactions are best observed in a slumped interval of sedi-ments that occurs between 13.0 and 17.5 Ma. As a result of microbial reduction of manganese and iron oxides and dis-solved sulfate, it is hypothesized that small amounts of mixed-metal carbonate cements are precipitated. These have lowcarbon isotopic ratios and high concentrations of metals.

INTRODUCTION

A great deal of work in recent years has been carried out onthe reconstruction of Cenozoic ocean temperature and composi-tion from the isotopic and chemical compositions of marinecarbonate sediments. Most often, these studies have been basedupon the measurement of stable oxygen and carbon isotopic ra-tios in foraminifers or total carbonate compositions of bulk sed-iments. Other measurements have included various elementalconcentrations in the carbonates (e.g., cadmium, lithium, bar-ium, strontium, magnesium, iron, manganese, sodium, and rareearth elements) and various isotopic ratios (e.g., neodymiumand strontium). For some of these parameters, the sources ofsecular variation are reasonably well understood (e.g., total car-bonate, and strontium-, carbon-, and oxygen-isotopic ratios)and useful interpretations can be made from measured data.For many of the measurements, however, especially for many ofthe minor elements, virtually ad hoc decisions have been made,partly by necessity, as to what processes control their measuredvariations.

In this study we present minor element and stable isotopicmeasurements made on a large suite of carefully cleaned sam-ples of the fine fraction of carbonate sediments from Site 709.We propose that the major processes controlling the concentra-tions of all of the reported elements and isotopes are (1) ancientocean surface temperature at the time of carbonate precipita-tion, (2) ancient ocean chemical composition at the time of car-

Duncan, R. A., Backman, J., Peterson, L. C , et al., 1990. Proc. ODP, Sci.Results, 115: College Station, TX (Ocean Drilling Program).

2 Department of Geology, Duke University, Durham, NC 27708, U.S.A.3 Rosenstiel School of Marine and Atmospheric Sciences, University of Mi-

ami, 4600 Rickenbacker Causeway, Miami, FL 33149, U.S.A.

bonate precipitation, and (3) near-surface or deeper burial dia-genesis. All of these processes are important in some cases and,as will be seen, it is often very difficult to deduce the fractionalrole played by each. Each element or stable isotopic ratio re-sponds differently to each of the forcing variables. For a varietyof reasons, we think that a relative ranking of the presently re-ported parameters in terms of the fidelity of their recording ofpaleotemperature or paleochemistry (instead of diagenesis) is:CaCO3 > 5I3C > strontium > magnesium > iron > manga-nese > 518O. That is, manganese concentrations in carbonatesshould be a lot more susceptible to diagenetic alteration than to-tal carbonate variations. If minor elemental variations coincidewith variations of total carbonate or 513C, then we reason thatthey may be reflecting true paleoceanographic variations ratherthan postdepositional changes.

METHODSSamples at Site 709 were collected with the hydraulic piston

corer and the advanced hydraulic piston corer. A total of ap-proximately 220 samples were used in this study. Approximatelyone g of each sample was treated for 30 min with 30 ml of 3%H2O2 in a water bath maintained at 55°C. This procedure disag-gregated the samples and oxidized any organic carbon present inthe samples. The samples were then wet sieved with deionizedwater and the fine fraction (<63 fim) was retained. Sampleswere centrifuged for 30 min and the supernatant liquid was de-canted. A reductive cleaning step was performed next with amixed reagent freshly prepared by combining 40 ml of 0.3 M so-dium citrate and 5 ml of 1.0 M sodium bicarbonate. Sampleswere reacted for 30 min in a water bath at 55 °C with 9 ml ofmixed reagent plus 0.5 g of sodium dithionate. Samples werethen centrifuged and the supernatant was decanted. The sam-ples were rinsed three times with distilled water and centrifugedafter each rinse. Finally, the samples were dried and precisely

661

P. A. BAKER, M. J. MALONE, S. J. BURNS, P. K. SWART

weighed subsamples of about 50 mg were dissolved in 50 ml ofammonium acetate-acetic acid (pH = 5.5) solution for about 10min at room temperature. The samples were centrifuged for 30min and the supernatant liquid was saved for analysis.

All elemental analyses were conducted with standard meth-ods using a Perkin Elmer Model 5000 atomic absorption spec-trophotometer. Replicate analyses indicate the following relativeerrors: 1.7% for CaCO3, 2.2% for strontium, 3.4% for magne-sium, 6.4% for manganese, and 24.4% for iron. Since typicalinstrumental errors are about l%-2% for all elements, it isclear that sample heterogeneity is responsible for the large un-certainties in manganese and iron determinations. Uncertaintiesin CaCO3 determinations are included in all the other elementaldata.

Stable isotopic analyses were performed on fine-fraction sam-ples that had been treated in exactly the same fashion as thesamples for elemental analysis (except for the final dissolutionstep). Carbon and oxygen isotopic analyses were conducted withphosphoric acid dissolution using an automated on-line devicereacting at 90°C for 10 min. Extracted CO2 was analyzed on aMAT 251 stable isotope ratio mass spectrometer at the Univer-sity of Miami. All results are reported relative to the PDB stan-dard. External error, determined by analyses of 20 replicatesamples, is ±0.026%0 for 518O and ±O.O16%o for 513C.

Sample ages were derived from nannofossil zonation and theages for the zonal boundaries are presented in Shipboard Scien-tific Party (1988a, table 3). Within the zones, ages were deter-mined by interpolation assuming constant sedimentation rates.

RESULTS AND DISCUSSIONAll of the chemical and isotopic results from this study are

presented in Table 1. The most relevant data are shown in Fig-ures 1 through 7.

Carbonate ContentsCarbonate contents of the sediments are shown in Figure 1.

The vast majority of samples have carbonate contents greaterthan 80%, making them suitable for this study. These data com-pare quite well with the carbonate determinations of bulk sam-ples shown in Shipboard Scientific Party (1988b, p. 486, fig.20). Major features include a carbonate low around 47 Ma, anincrease in carbonate from 47 to 31 Ma, and two carbonatehighs at 31 and 26 Ma with an intervening low at 29 Ma. Thereis a carbonate low at 20 Ma followed by a high centered at 18Ma. This is followed by an interval of slumping between ap-proximately 13.0 and 17.5 Ma (Shipboard Scientific Party, 1988b,p. 468); within this interval there are two discrete fluctuations ofcarbonate content. Each is associated with a slump block. Thecarbonate contents are thought to be repeats of the top of theunderlying, nonslumped interval. Since 13 Ma, fluctuations havebeen somewhat smaller with a significant carbonate low show-ing up around 10 Ma (seen much better in the data of ShipboardScientific Party, 1988b) and a carbonate high around 5 Ma.There appears to be some correlation between the overall sedi-mentation rate and the carbonate content (higher sedimentationrates are associated with higher carbonate contents), but thesedimentation rates are not very well constrained at this resolu-tion.

Carbon IsotopesCarbon isotopic data are shown in Figure 2. This record is

similar to the carbon isotopic compositions of bulk carbonatesediments of the same age at Deep Sea Drilling Project (DSDP)Sites 525 and 528 reported by Shackleton and Hall (1984) andShackleton (1987). The first-order features of both data sets are:high, positive values of <513C from 47 to about 20 Ma, low ornegative values at present, and a period of decrease of 513C

from 10 Ma to the present. One notable feature, different frompreviously published curves is the 1.6%o decrease in 513C in theearly Miocene. Because this decrease is coincident with large in-creases in iron, manganese, and magnesium, however, we thinkthat it is produced by a diagenetic process that occurs duringslumping (discussed below). It should be noted that the isotopicchange is also coincident with a sharp decrease in the carbonatemass accumulation rate (Peterson and Backman, this volume)at neighboring (nonslumped) sites.

It is evident that there is also fine structure in the carbon iso-topic record at Site 709. Before 20 Ma there are about five iso-topic excursions, each lasting about 5 m.y. There are two majorpositive excursions in this period. These occur during the earlyOligocene (about 32 Ma) and just after the Oligocene/Mioceneboundary (about 23 Ma). There are three major negative excur-sions in the same period. They are centered at about 36 Ma(near the Eocene/Oligocene boundary), 30 Ma, and 26 Ma.

Shackleton (1987) interpreted the decrease in the 513C of bulkcarbonate sediments since the middle Miocene as a global oce-anic signal reflecting a long-term decrease in the amount of or-ganic carbon buried in sediments. Similarly, the positive excur-sions before 20 Ma may represent shorter time periods of glob-ally increased burial of organic carbon. Vincent and Berger(1985) identified a positive carbon isotopic excursion of about1 per mil lasting about 4 m.y. centered on 15.5 Ma (also ob-served by Shackleton, 1987). This shift was seen in both benthicand planktonic foraminifers and thus is considered to reflect aglobal oceanic change. Vincent and Berger (1985) suggested thatthis shift was produced by the increased burial of organic car-bon beneath regions of high productivity. At Site 709 there is apositive isotopic excursion of about 1 per mil in about the sametime interval, although because of the slumping in the 13-17.5Ma interval at this site, the exact stratigraphic position of thisisotopic peak is unknown.

One of the most obvious means of increasing the amount ofcarbon buried in marine sediments is to increase the area ofshallow-water sedimentation by eustatic sea-level rise. Positive<513C excursions, therefore, should be associated with high eu-static sea levels. This prediction seems to be borne out in apply-ing the Cenozoic eustatic sea-level cycle chart of Haq et al.(1987) converted to the Leg 115 time scale. High sea-level standsoccurred from 50 to 45 Ma, 42 to 40 Ma, 35 to 30 Ma, 24 to23 Ma, and 17 to 14 Ma. Each of these intervals, including theexcursion discussed by Vincent and Berger (1985), is also a timeof higher <513C.

These conclusions must be weighed carefully. Miller and Fair-banks (1985) found quite different timing of "global" carbonisotopic cycles in the Oligocene and Miocene epochs, and theyconcluded that these cycles were independent of eustatic sea-level changes. Shackleton (1987) concluded that the possible as-sociations between 513C events and sea level could not "be testedadequately until the sea-level record is better known and under-stood." Shackleton (1987) did not have available to him thechronology and sea-level cyclicity chart published by Haq et al.(1987). In fact, since the Paleocene, there does seem to be co-incidence between sea-level highstands (Haq et al., 1987) andShackleton's (1987) positive 513C excursions.

StrontiumStrontium concentrations of ancient marine carbonate sedi-

ments were measured previously by many workers. It has beenproposed that these concentrations were influenced by severalfactors, including leaching of strontium from noncarbonatesources during sample preparation, variations of the oceanic ra-tio of dissolved strontium-to-calcium (Sr/Ca) (e.g., Graham etal., 1982; Renard, 1986), variations of oceanic temperature (Kins-man, 1969), and diagenesis (e.g., Baker et al., 1982). We think

662

CARBONATE FINE-FRACTION COMPOSITION

Table 1. Elemental and stable isotopic compositions of fine fraction samples from Site 709.

Core, section

115-709 A-

1H-11H-21H-31H-41H-61H-72H-12H-22H-32H-63H-13H-23H-33H-43H-53H-64H-14H-24H-34H-44H-56H-16H-16H-16H-16H-26H-36H-46H-56H-67H-17H-27H-37H-47H-57H-68H-18H-28H-38H-48H-58H-69H-19H-29H-39H-49H-59H-610H-110H-210H-310H-410H-610H-711H-111H-211H-311H-411H-511H-612H-112H-212H-412H-512H-613H-113H-213H-313H-413H-513H-614H-114H-214H-314H-414H-5

Depth(mbsf)

0.982.483.985.488.489.88

11.0812.5814.0818.5820.6822.1823.6825.1826.6828.1830.3831.8833.3834.8835.8

49.6849.6849.6849.6851.1852.6854.1855.6857.1859.2860.7862.2863.7865.2866.7868.9870.4871.9873.4874.9875.4878.5880.0881.5883.0884.5886.0888.2889.7891.2892.7895.7896.9

97.8899.38

100.88102.38103.88105.38107.58109.08112.08113.58115.08117.18118.68120.18121.68123.18124.68126.78128.28129.78131.28132.78

Wt(mg)

49.850.050.850.850.050.949.949.050.949.349.750.650.050.850.450.550.050.051.049.849.950.050.149.350.350.750.050.551.050.749.750.750.949.949.549.951.049.050.650.551.050.049.250.951.050.950.049.850.950.050.251.049.950.850.749.850.450.049.949.950.350.750.550.850.349.849.250.549.350.550.850.550.049.050.050.0

CaCO3

(mg)

44.7743.7744.7747.2743.3941.0239.6441.3944.1443.0241.6444.1443.7740.6442.3943.0230.6332.6433.0142.7741.0244.1443.7745.1445.5245.8945.5245.5245.1445.1443.7743.7744.7744.7742.0242.0245.1443.0246.1444.1446.1443.0243.7742.7744.1445.1443.0244.5242.0238.2742.0241.3942.0244.5245.8944.1443.7742.7742.0241.6443.7743.3942.7741.3942.7742.0141.3946.5242.3945.1445.5244.1442.3942.3942.3943.39

CaCO3

(wt%)

89.9087.5488.1393.0586.7980.5879.4484.4786.7287.2683.7987.2487.5480.0084.1185.1861.2865.2864.7383.8682.2088.2987.3691.2790.4990.5291.0490.1488.5289.0488.0686.3387.9589.7284.8884.2088.5287.7991.1987.4190.4886.0388.9684.0286.5588.6986.0389.3982.5576.5383.7081.1684.2087.6390.5288.6486.8485.5484.2083.4587.0185.5984.6981.4885.0284.3784.1392.1285.9989.3989.6087.4184.7886.5184.7886.79

Sr(ppm)

1798201121222084206320482220202919941953223317102011185820762080274228032545184717921869191917831812183018451692170617501759180518431843188019991861195316691869155017551634184718241673199916961880201217971776191616061722162015991672174921131634161318471691188218811945177316511584176815861486152215221567

Mn(ppm)

268217212159311280202387284372276204366381295321457444515444414261274255253240242297244277308286265279297286266291314272271209286257227244174225274284238302238303240249228234262324354415468664736559568645720565593510330271201196

Fe(ppm)

480605681635553549744640544802516510765554401337212276197281219295331354286490319406476476605583447380345535354395553510347523388304272410395573607444476290309382436419514526345516297242245314281250193226283177351419342401377369

Mg(ppm)

905685748497772987757858872872949770982

13161050872

129011031091982

1110997

1017964956904923

10331008997

10401040994905952940853872

1084997

1062930

11121040952

1019860

10331047108410951148976

1123654

1053114210761047672

10511060947

12561052964930946

10261063114211551345132113211129

513C

0.350.130.411.09

-0 .22-0 .39

0.250.610.250.220.220.320.48

-0.230.210.400.280.310.340.38

-0 .160.640.640.640.640.000.541.040.210.560.82

-0.010.350.560.81

-0 .050.800.620.660.301.010.540.280.611.040.650.580.500.920.780.930.810.990.771.091.760.740.861.180.990.811.020.970.990.990.901.411.290.950.930.941.561.301.620.770.54

518O

-0.20-0 .60-0 .02-0 .09-0 .38-0.11-0 .50-0 .45-0.25-0.13-0.03

0.09-0 .77-0 .32-0.58-0.41-0.58-0.53-0 .94-0 .29-0 .50

0.190.190.190.190.040.170.59

-0 .330.070.27

-0.21-0.15

0.040.15

-0 .400.31

-0 .390.690.030.16

-0 .300.280.230.200.24

-0 .300.00

-0.03-0 .24

0.03-0 .44-0.81

0.160.000.02

-0 .240.00

-0 .090.040.06

-0 .05-0 .10-0 .09

0.15-0.23

0.010.470.470.610.450.190.330.230.13

-0 .04

Age(Ma)

0.110.270.370.460.790.951.081.251.421.942.142.282.422.562.702.833.043.183.323.463.544.404.404.404.404.494.584.664.734.814.914.995.175.145.225.305.415.495.565.645.725.745.905.986.066.136.216.296.406.486.616.777.087.207.307.467.617.777.938.088.318.478.779.0810.0610.9311.1111.2811.4511.6211.7912.0312.2012.3712.5512.72

663


Table 1 (continued).

Core, section

115-709A- (Cont.)

14H-615H-115H-215H-315H-415H-515H-616H-116H-216H-316H-416H-516H-617H-117H-217H-317H-417H-517H-618H-118H-218H-318H-418H-518H-619H-119H-219H-319H-420H-120H-220H-320H-420H-520H-621H-121H-221H-321H-421H-521H-521H-521H-521H-6

I15-709B-

21X-121X-221X-321X-421X-521X-622X-122X-222X-322X-422X-522X-623X-123X-223X-323X-323X-323X-323X-423X-524X-124X-224X-324X-424X-524X-625X-125X-225X-3

Depth(mbsf)

134.28136.48137.98139.48140.98142.48143.98146.18147.68149.18150.68152.18153.68155.83157.33158.83160.33161.83163.33165.58167.08168.58170.08171.58173.08175.28176.78178.28179.6184.88186.38187.88189.38190.88192.38194.48195.98197.48198.98200.48200.48200.48200.48201.98

188.18189.68191.18192.68194.18195.68197.78199.28200.78202.28203.78205.28207.48208.98210.48210.48210.48210.48211.98213.48217.08218.58220.08221.58223.08224.58226.78228.28229.78

Wt(mg)

50.049.850.049.250.850.049.949.649.849.950.050.550.049.950.050.250.150.050.951.049.150.149.149.049.949.649.850.050.850.449.849.950.549.349.749.550.049.850.349.850.150.049.649.7

50.949.350.249.250.550.051.049.950.750.850.950.450.050.950.750.050.550.949.850.850.049.950.050.950.249.850.750.250.9

CaCO3

(mg)

40.8943.0243.0243.3944.1444.5244.5241.3945.1443.0244.7743.3943.0236.1438.8939.6444.1443.3944.5244.7743.3943.3942.0239.8934.8938.2738.8939.6437.5245.5244.1443.7746.5243.3944.5244.1446.8944.5245.5246.5246.5245.8946.8944.77

46.1444.1444.7743.0244.7746.8947.2744.1446.8946.5245.5245.8947.8945.5144.1442.3942.0241.6441.0242.0240.2739.6441.0244.7743.3941.6444.1443.3944.77

CaCO3

(wt%)

81.7886.3886.0388.2086.9089.0489.2183.4590.6586.2189.5485.9386.0372.4277.7878.9788.1186.7987.4687.7888.3886.6185.5781.4169.9277.1578.0979.2873.8590.3188.6487.7192.1288.0289.5789.1893.7989.3990.4993.4192.8591.7994.5490.08

90.6689.5489.1887.4388.6593.7992.6888.4692.4991.5789.4391.0695.7989.4387.0784.7883.2081.8182.3682.7180.5379.4482.0387.9586.4483.6287.0786.4487.95

Sr(ppm)

14921709170916131620152716511691150617091485161316272089189016781461156718871686174017751797161720491868198020812146201017901805184916941853179019511774184518921924191818771966

18211744196617901876168518621869175919241966191819421846194820292047211320972178222222582011204419362017218619821966

Mn(ppm)

220232291207861752472532521511313346349789669568498507494469323876940840946614476555453406340366376403382306266247253215215196192201

271396357232391299243283192247242305115297261271274300305297372290268313288336306323335

Fe(ppm)

25733727915027238229244753239538039239566443750541933450544730046119088

530562437227

27176113114140150180113139236

5514014014275

235

282272447116145171106113277226

77283

5226414730715519212215532326512229015031227227778

Mg(ppm)

12961069976

103714051269126914741174110411951210118613141234133712011152977

10611221968

10001203147612281119971986780895971849

1060921

1008864

1000945881915904863871

964974905988927

1034899940981946923981793989986

1215952997

1109964981933865860922877770853771

513C

0.550.550.370.641.481.561.341.200.900.940.760.450.491.321.010.750.530.661.050.670.560.360.590.520.790.840.911.191.221.951.821.682.071.781.801.312.311.311.081.181.181.181.181.34

1.861.872.521.751.761.692.031.561.261.241.261.241.201.231.321.321.321.321.191.281.551.441.611.411.211.131.501.301.01

518O

0.040.33

-0 .13-0 .01-0 .07-0 .02-0.03-0 .18-0 .36-0 .24-0 .03

0.08-0 .04

0.03-0 .16

0.420.150.13

-0 .060.060.50

-0 .060.040.190.100.040.08

-0 .15-0 .20

0.050.190.15

-0 .020.300.230.050.280.020.10

-0 .16-0 .16-0 .16-0 .16-0.08

0.360.140.170.500.170.660.310.05

-0 .010.080.00

-0.14-0 .04-0.15

0.040.040.040.04

-0 .06-0.04

0.100.200.13

-0 .060.07

-0 .09-0 .17-0.17-0 .26

Age(Ma)

12.8913.1413.3613.6113.8514.1014.3414.7014.9515.1915.4415.6815.9316.2816.5316.7717.0217.2617.5117.8818.1218.3718.6219.0219.4119.9920.3820.7821.1222.5122.9123.2623.5023.7423.9824.3224.5624.8025.0425.2825.2825.2825.2825.54

23.3123.5523.7924.0324.2724.5124.8425.0825.3325.5925.8526.1126.4826.7427.0027.0027.0027.0027.2527.5128.1328.3128.4728.6328.7928.9529.1829.3329.49

664



Core, section

115-709B- (Com.)

25X-425X-525X-626X-126X-226X-326X-426X-426X-426X-427X-127X-227X-327X-427X-5

115-709C-

26X-126X-226X-326X-426X-526X-627X-127X-227X-328X-128X-228X-329X-129X-229X-329X-429X-529X-630X-130X-130X-130X-130X-230X-330X-43OX-531X-131X-231X-331X-431X-532X-132X-232X-332X-432X-532X-633X-133X-233X-333X-433X-534X-134X-234X-334X-434X-534X-635X-135X-335X-435X-435X-435X-435X-535X-636X-136X-2

Depth(mbsf)

231.28232.78234.28236.48237.98239.48240.98240.98240.98240.98246.08247.58249.08250.58252.08

238.38239.88241.38242.88244.38245.88247.98249.48250.98257.68259.18260.68267.38268.88270.38271.88273.88274.88277.08277.08277.08277.08278.58280.08281.58283.08286.78288.28289.78291.28292.78296.38297.88299.38300.88302.38303.88306.08307.58309.88310.58312.08315.78317.28318.78320.18321.78323.28325.38328.38329.88329.88329.88329.88331.38332.9335.08336.58

Wt(nig)

50.051.050.450.449.950.051.049.950.650.250.849.550.750.050.0

49.949.950.650.650.550.350.750.150.350.350.649.949.950.249.850.050.650.050.250.449.850.050.050.850.050.550.550.050.150.350.550.050.650.050.050.049.949.850.349.850.550.050.550.050.250.050.049.950.050.149.950.650.449.949.850.050.050.2

CaCO3

(mg)

44.7745.8943.7744.5245.5246.1446.8945.8947.2747.8946.1444.7745.1444.7743.77

44.1445.8946.1444.7747.8946.1444.1443.7743.0242.0246.1444.7744.7745.1434.8938.2738.2736.5143.7743.3943.0243.0244.1445.1444.7644.7643.7645.1443.7644.7740.6441.6444.1442.7744.1441.3932.1439.6444.7744.1444.1444.5142.3941.3942.3939.6445.1443.3944.7741.0239.8942.0241.6441.0239.8941.0239.6441.02

CaCO3

(wt%)

89.5489.9986.8488.3391.2292.2991.9591.9793.4295.4190.8390.4489.0489.5487.54

88.4691.9791.1988.4794.8491.7487.0787.3685.5283.5391.1989.7289.7289.9370.0676.5376.5373.0387.1986.1086.8886.0388.2988.8689.5488.6587.5490.2987.3689.0080.4883.2888.0085.5388.2982.7864.4179.6089.0088.6487.4189.0483.9482.7884.4579.2890.2986.9689.5481.8779.9483.0482.6282.2080.1082.0379.2881.71

Sr(ppm)

199921461885188716151669179117541745168120591999202719211919

179017981821172020991983215219191953188017121921187619492107197319211863209020632080208014611750200014861348115213481318155014651257134413371305140017661206122311441011106211231062996676749849841953904949926990792820926

Mn(ppm)

279327343248253271245272264240142190188123206

28325121724620919522719525638136822333531050232732737028628826729024924425730226325525127928320419323420415735821419019319390153205177189122196145158138178156134163134164183

Fe(PPm)

2681091831461761412242611902517656

26678

297

1475476

11216717311324011619014114511214472

170131685758

116151204111179235

8055

240145123216113117181157109126787957

1461181571538876

1151122463246061

12512225

122

Mg(ppm)

804752891854923878832795772773715715764804845

816774726648606726759111802

1069899838893820

1290120213201452674807732732

12911086385530

1337149513371340115611891201135611551268163310221094130313371499151015701510169014181821

?162115791535153715481529168217661646

613C

1.261.601.111.400.861.211.381.381.381.381.411.811.891.531.96

1.181.331.491.711.911.671.871.801.851.141.331.431.281.361.201.221.301.53.26.26.26

1.261.612.101.281.211.511.891.891.781.63.66.69.65.57.80.70

1.06.76.58.78.85.78.74.60.70.65.72

1.601.771.761.761.761.761.811.641.71.79

518O

-0.13-0.29-0 .16-0 .06

0.220.060.380.380.380.38

-0.05-0.11

0.230.01

-0 .02

0.090.010.040.55

-0.13-0 .25

0.050.070.010.050.12

-0.12-0.13-0.25

0.000.120.250.58

-0.75-0.75-0.75-0 .75

0.500.46

-0 .79-0 .26

0.090.280.18

-0 .04-0 .50-0.03

0.310.300.000.49

-0 .34-0 .24-0 .22

0.220.230.740.130.220.220.710.670.940.170.160.330.330.330.330.250.470.520.31

Age(Ma)

29.6529.8129.9730.2030.3730.5430.7030.7030.7030.7031.2831.4531.6231.7831.95

30.4130.5830.7530.9231.0931.2631.4931.6631.8332.5832.7532.9233.6733.8634.0834.3034.5934.7435.6235.6235.6235.6236.0736.3136.5536.7737.2037.3837.5537.7337.9038.3238.5038.6738.8539.0339.2039.4639.6340.0040.1640.5141.3641.7042.0542.3742.7443.0843.5744.2644.5144.5144.5144.5144.7044.8945.1745.36

665



Depth Wt CaCO3

Core, section (mbsf) (mg) (mg)CaCO3 Sr Mn Fe Mg(wt<%) (ppm) (ppm) (ppm) (ppm) 61SO

Age(Ma)

115-709C- (Com.)

36X-337X-137X-237X-337X-437X-537X-6

338.08344.88346.38347.88349.38350.88352.38

49.849.950.350.850.050.951.0

41.0237.5238.8934.3933.6440.6442.39

82.3675.1877.3267.6967.2879.8583.12

792920836874

1026886637

183227219218163160177

61936429746283

1646154617361634168015131757

1.711.581.541.591.661.551.48

0.250.040.020.180.340.330.60

45.5546.4146.6046.7946.9847.1847.37

ooCQO

7 0 -

60

1 0 20 30 40

Age (Ma)

Figure 1. Three-point moving average of CaCO3 (wt%) in the fine fraction of sediments vs. age.

that we can rule out noncarbonate sources of strontium duringleaching because of our sample cleaning procedures.

Changes in the Sr/Ca ratio of carbonate sediments can beascribed either to changes in the distribution coefficient of stron-tium in calcite (Dc), thought to be primarily caused by tempera-ture changes, or to the Sr/Ca ratio in seawater, (Sr/Ca)5VV:

(Sr/Ca) = Dc (Sr/Ca)5(V

Experimental work has shown that strontium uptake duringnonbiogenic calcite growth is not strongly temperature depen-dent (Katz et al., 1972; Baker et al., 1982); however, there isconflicting evidence on the importance of temperature depen-dence on the strontium content of foraminiferal calcite. Delaneyet al. (1985) found little correlation between seawater tempera-ture and strontium concentrations in either core-top planktonic

foraminifers or in laboratory-cultured foraminifers. On the otherhand, Cronblad and Malmgren (1981) observed large and con-sistent glacial/interglacial variations of strontium and magne-sium in foraminifers from Pleistocene sediments near the PolarFront in the southern Indian Ocean. The positive correlationthey observed between warm-water climatic indexes and stron-tium concentrations in planktonic foraminifers strongly suggestsa temperature control on the strontium.

Graham et al. (1982) suggested that strontium concentrationvariations preserved in ancient sediments may reflect originalvariations in the Sr/Ca ratio in ancient seawater. On a muchlarger time scale, Renard (1986) attempted to correlate stron-tium and magnesium concentration changes in bulk pelagic car-bonate sediments from many different localities and ascribedthese changes to Sr/Ca and Mg/Ca secular variation in seawa-ter. Recently, Boyle and Keigwin (1982) and Lea and Boyle

666


a.

o

Age (Ma)

Figure 2. Three-point moving average of the carbon isotopic ratios in the carbonate fine fraction vs. age.

(1989) have convincingly shown that variations in seawater bar-ium and cadmium concentrations are recorded by benthic fora-minifers. Although neither strontium nor magnesium show verysignificant nonconservative behavior in modern-day seawater,past variations in their concentrations, relative to calcium con-centrations, can be expected to be recorded by carbonate-pro-ducing organisms.

Graham et al. (1982) proposed two major mechanisms forchanging oceanic Sr/Ca ratios: (1) variable hydrothermal inputof calcium to the oceans at the ridge crests and (2) variable ratesof strontium removal resulting from changes in the proportionof high-strontium aragonite vs. low-strontium calcite sedimen-tation. The ratio of aragonite to calcite should increase duringhigh sea-level stands when more shallow carbonate banks areflooded. Graham et al. (1982) concluded that both processesplayed important roles in determining Cenozoic Sr/Ca ratios.Delaney and Boyle (1986) concluded that for the past 40 m.y.both Li/Ca and Sr/Ca ratios in foraminifers changed very littleand that lack of variability in the Li/Ca ratio implied that therehas been no significant variability in oceanic hydrothermalfluxes. They also concluded that over the past 100 m.y. there hasbeen less than a factor of 2 decrease in hydrothermal chemicalfluxes. On the other hand, Renard (1986) has reported numer-ous larger variations of Sr/Ca and Mg/Ca ratios in bulk pelagiccarbonate samples and once again attributed these variations toboth hydrothermal and sea-level variations.

Increased hydrothermal activity should bring about an in-crease in oceanic-dissolved strontium but a larger increase inoceanic-dissolved calcium; hence, the Sr/Ca ratio of seawater

should decrease (Palmer and Edmond, 1989). Likewise, increasedhydrothermal activity is thought to cause global sea-level rise, inwhich case more aragonite sedimentation would be expected onshallow shelves and bank tops and the pelagic Sr/Ca of seawa-ter should decrease.

Strontium concentration is a useful indicator of carbonatediagenesis because foraminifers and, especially, nannofossils pre-cipitate calcite that is enriched in strontium relative to equilib-rium calcite. During burial diagenesis, some of the biogenic cal-cite is dissolved and reprecipitated. In the process, strontium isreleased into the pore fluids (Baker et al., 1982) and diagenesisresults in a lower concentration of strontium in older carbonatesediments.

The potential importance of diagenesis in determining theobserved strontium concentrations (Fig. 3) can be evaluated withthe pore-water strontium analyses for Site 709 reported by Swartand Burns (in this volume). From their data, we can estimate apresent-day, near-surface, dissolved strontium gradient of about4.2 x 1011 mol/cm4. Using Fick's Law and a diffusion coeffi-cient in the sediments of about 3 x 10 ~6 cmVs, we can calcu-late a diffusive flux of strontium out of sedimentary pore watersinto the overlying seawater of 4.0 x 10"3 mol/cmVm.y. Com-pare this to the rate of accumulation of strontium in the sedi-ments (at 75% porosity, 90% CaCO3, 2100 ppm Sr, 10 m/m.y.sedimentation rate): 1.94 x 10~2 mol/cm2/m.y. (note that thisis cm2 of pore water). Thus, about 20% of all the strontium thataccumulates on the seafloor diffuses back into overlying seawa-ter (and, at the very least, 20% of the carbonate has been re-crystallized). However, as shown previously (Baker, 1985; Rich-

667


3000

2000-

Q .

1000-

Age (Ma)

Figure 3. Three-point moving average of the strontium concentrations in the carbonate portion of the fine fraction vs. age.

ter and DePaolo, 1988), this number is probably closer to 50%recrystallized within about 10 m.y. of deposition. Almost all ofthis strontium is lost from the upper 100 m of sediment, sincebelow this depth there is no observed pore-water strontium con-centration gradient (Swart and Burns, this volume). If we as-sume that the present-day average strontium concentration inthe carbonate fine fraction of 2100 ppm was typical of sedi-ments for the last several million years, then these calculationsgive us the result that after about 7-m.y. (the age at 100 m belowseafloor [mbsf]) of diagenesis there should be a decrease toabout 1680 ppm. In fact, this is approximately the magnitude ofconcentration change observed in sediments from 7 Ma to thepresent.

The factor that seems most likely to change the dissolvedstrontium gradient out of the sediments and the resulting diage-netic strontium loss of ancient sediments is bottom-water (hence,pore-water) temperature. Temperature can affect the strontiumdistribution coefficient in calcite and the diffusion coefficientof strontium in sediments, but the most thermally activated pro-cess is the rate of calcite recrystallization. For example, assum-ing that a 10°C rise in bottom-water temperature can produce adoubling of the rate of calcite dissolution and reprecipitation,then we could expect roughly a doubling of the rate of stron-tium release to the pore waters and a doubling of the diffusiveflux of strontium to overlying seawater. The net result should belower strontium concentrations in the sediments during periodsof higher bottom-water temperature.

There is a feedback mechanism to the above-mentioned pro-cess. At the present time, calcite produced during diagenesis be-

low 100 mbsf should have about 1160 ppm strontium (given theaverage Sr/Ca ratio of pore waters observed below 100 mbsfand a distribution coefficient of strontium in calcite of 0.04;Baker, 1985). Clearly, except in the Eocene, diagenetic calciteprobably constitutes only about 50% of the total fine fraction(assuming for the moment an original strontium content of2000 ppm). If rates of recrystallization are sufficiently acceler-ated, the Sr/Ca ratio of pore waters would be elevated, the dia-genetic calcite would have a higher strontium content, and therewould be a decreased loss of strontium from the sediments.

In summary, diagenesis has been demonstrably important ininfluencing the strontium content of the carbonate sediments atSite 709, but original variations of the strontium concentrationsof the carbonate, caused by paleotemperature changes or changesin seawater Sr/Ca ratios, could persist. In the simple case where50% of the biogenic calcite has been replaced by diagenetic cal-cite of a fixed composition, the original concentration variationwould be halved by diagenesis. Accordingly, to crudely decon-volve the strontium curve shown in Figure 3, we should do thefollowing:

1. Fix the modern-day ratio at the observed value.2. Back to 7 Ma, adjust the observed values for linearly in-

creasing amounts of diagenetic calcite to 50% calcite of 1160ppm strontium by 7 Ma.

3. Before 7 Ma, any carbonate with greater than 1580 ppmoriginally had >2000 ppm and any carbonate with < 1580 ppmoriginally had <2000 ppm. If we assume a 50% addition ofdiagenetic carbonate of 1160 ppm, then the following simple al-

668


EQ.a.

1800

1600-

1400-

1200-

1000^

800-

600-

400 i

200-

50

Age (Ma)

Figure 4. Three-point moving average of the magnesium concentrations in the carbonate portion of the fine fraction vs.age.

gorithm can be used to calculate the original strontium contentof carbonate (Sr,) from the present-day observed value (Sr):

2 x Sr - 1160 = Sr, (in ppm).

4. Note that we have not accounted for any variation ofrates of diagenetic strontium loss and we have assumed steady-state conditions in the pore waters throughout.

The deconvolved strontium concentration curve (Fig. 8) isquite similar to the results published by Graham et al. (1982) forSr/Ca ratios of planktonic foraminifers. The most striking fea-ture of both studies, also noted by Renard (1986) and apparentin Delaney and Boyle (1988), is the late Eocene strontium in-crease. This feature coincides with the late Eocene initiation ofthe rise in the 87Sr/86Sr ratio in seawater (Elderfield, 1986). Italso coincides with the decrease in seafloor spreading rates (Pit-man, 1978) and the lowering of sea level from middle Eocenehighs (Haq et al., 1987).

Elsewhere, strontium concentrations display less dramatic var-iations, but there seems to be some correlation between high sealevel stands and low strontium concentrations, and vice versa.For example, some of the lowest sea levels of the Neogene werethought to have occurred during the middle and late Miocenecentered at about 10 Ma. A strontium peak also occurs at thistime. Low sea levels in the late Oligocene coincide with highstrontium concentrations. Low sea levels in the late Plioceneand Pleistocene coincide with high strontium concentrations.

This implies that sea level is important, but it does not distin-guish between the hydrothermal and shallow-water aragoniticcauses.

What fraction of the strontium concentration change canreasonably be explained by variations in the aragonite/calciteratio? This question was originally addressed by Graham et al.(1982) and it is their model, with a minor adaptation, that weapply below to look again at this problem.

The oceanic mass balance for strontium can be written as:

dSr/dt = FD + FR - Fs,

where F is the flux of strontium to the ocean and subscripts D,R, and S represent the diagenetic, riverine, and carbonate sedi-mentation fluxes of strontium, respectively. The value, FD, wasdiscussed by Elderfield and Gieskes (1982). As noted by Gra-ham et al. (1982), Fs is given by:

FS = fAFA + (1 - /AWC*

where FA and Fc are the sedimentary fluxes of aragonite andcalcite, respectively, and/4 is the fraction of aragonite.

Furthermore,

Fs = fADA(Sr/Csi)swJ - fA)Dc(Sr/Ca)swJ,

where / is the burial flux of carbonate sediments. We have as-sumed values of DA = 1.00 and Dc = 0.23 from present-day

669


800

600-

EQ.Q.

4 0 0 -

200-

40 50

Age (Ma)

Figure 5. Three-point moving average of the iron concentrations in the carbonate portion of the fine fraction vs. age.

values of the Sr/Ca ratio of seawater, aragonitic benthic greenalgae and corals, and bulk pelagic carbonate sediments. Herewe take

FD = (0.2)Dc(Sr/Ca)sw/

because, as discussed above, about 20% of the strontium inmodern carbonate sediments is returned to overlying seawater(and possibly higher in aragonitic sediments). Then, at steadystate

FR/J(Sr/Csi)sw = 0.77/4 + 0.184.

At the present time,

(FD + FR)/J = 0.0025,

which can be adjusted to make the present-day fA come out atthe most appropriate value. Substituting for FD, we find that

FR/J = 0.0021,

which is very close to the estimate of Delaney and Boyle (1988).We can assume, for the sake of discussion, that this ratio

(FR/J) has been constant throughout time. This allows us toevaluate independently the importance of changes in/4:

fA = 0.00273/(Sr/Ca)5VV - 0.239.

The results of this calculation are given in Table 2. One can seethat late Eocene values of 1000 ppm strontium (Sr/Ca =.00114) imply that about 30% of the total CaCO3 was depositedas high-strontium aragonite. During the late Oligocene sea-levellow stand (around 28 Ma), the high values of Sr/Ca imply avalue of/4 = 0 (i.e., there was no high-strontium aragonite;most shallow carbonate-producing banks must have been sub-aerially exposed). The model is adjusted to make modern valuesof high-strontium aragonite equal a reasonable value of totalCaCO3 sedimentation.

With regard to strontium concentration variations, we con-clude the following. First, diagenesis is clearly of importance indecreasing the strontium concentrations of carbonate sediments.Nevertheless, real paleoceanographically controlled variationsare preserved to some degree. The reasonable correlation be-tween high sea-level stands and low strontium values suggestssome kind of sea-level control, most likely through a shelf/ba-sin fractionation mechanism. High sea-level stands resulted inlarger fractions of aragonitic sediments and less strontium inseawater. As pointed out by Delaney and Boyle (1986), a similarshelf/basin fractionation mechanism can also account for thesecular changes observed in the carbonate compensation depth(CCD). Variations in hydrothermal fluxes may also producechanges in the oceanic Sr/Ca ratio, but Delaney and Boyle (1986)concluded that these changes were insufficient to account formore than a fraction of the observed strontium variations. Pa-leotemperature changes deduced from the oxygen isotopic rec-ord are also probably insufficient to account for the observedvariations of strontium.

670


1000

800-

600-

Q.a.

400-

200-

Age (Ma)

Figure 6. Three-point moving average of the manganese concentrations in the carbonate portion of the fine fraction vs.age.

Oxygen Isotopes

If it is correct, as deduced above, that diagenetic calcite con-stitutes about 50% of the carbonate sediments, then this shouldbe reflected in the oxygen isotopic composition of the fine frac-tion. In Figure 9, the approach of the carbonate sediments toapparent oxygen isotopic equilibration with pore waters is illus-trated. For the sake of these calculations it was assumed that (1)the 518O of pore waters was constant at O.O%o (SMOW) (near thepresent-day values measured by P. K. Swart, pers. comm.,1989), (2) the bottom-water temperature was 2°C, and (3) theaverage geothermal gradient at this site was about 35°C/km.The oxygen isotopic fractionation factor was calculated fromthe equation of O'Neil et al. (1969). From these calculations, wecan safely conclude that the course of burial diagenesis through-out the history of Site 709 would have involved an increase in<518O of the carbonate sediments. This is particularly evident ifwe assume, following Savin (1977), that before 15 Ma theoceans were about l%o lower in <518O than at present (of course,the pore waters before 15 Ma also would have been isotopicallylighter).

A comparison of the oxygen isotopic data (Fig. 7) with thestrontium concentration data (Fig. 3) reveals an inverse correla-tion: almost invariably, lower strontium values are associatedwith higher oxygen isotopic compositions. This relationship ismost clearly revealed in the Eocene chalks. Because similar co-incident variations would be expected to result from temperature

changes (colder temperatures should result in heavier oxygen iso-topic ratios and lower strontium concentrations) and diageneticchanges, it is not straightforward to differentiate between diage-netic and paleoclimatic causes for the observed relationship.

MagnesiumThe Mg/Ca ratios in carbonate sediments respond to the

same factors that influence Sr/Ca ratios, namely, the tempera-ture of precipitation, the Mg/Ca ratio of ancient seawater, anddiagenesis. Cronblad and Malmgren (1981) showed an apparentpositive correlation between near-surface, ocean-water tempera-ture and magnesium concentrations in planktonic foraminifers.The range of observed concentrations was from about 400 to800 ppm. Temperature controls, therefore, should bring about acorrelation between strontium and magnesium concentrations(which we did not observe). If sea-level changes can cause a sec-ular variation of aragonite/calcite ratios, they should also bringabout a similar variation of the ratio high-magnesium calcite/calcite. Because high-magnesium calcite is usually much lessabundant in shallow-water carbonate sediments than aragonite,however, these variations would not be very effective in alter-ing seawater Mg/Ca values. Hydrothermal input should have astrong influence on the seawater Mg/Ca ratio and in the samedirection as the seawater Sr/Ca ratio. Because of the longer resi-dence time of magnesium than strontium, however, maxima (orminima) in Mg/Ca ratios should follow the maxima (or min-ima) of Sr/Ca ratios by a few million years (e.g., Renard, 1986),

671


Q.

CO

Age (Ma)

Figure 7. Three-point moving average of the oxygen isotopic ratios in the carbonate fine fraction vs. age.

and variations should generally be positively correlated. Finally,diagenesis should bring about an increase in the Mg/Ca and adecrease in the Sr/Ca ratio of pelagic carbonate sediments (e.g.,Baker et al., 1982), thus producing an inverse correlation be-tween strontium and magnesium concentrations.

In comparing Figures 3 and 4, we observe a tendency forstrontium and magnesium to be inversely correlated. This isbest observed at the top of the core (in the last 3 m.y.) wherediagenetic recrystallization of calcite occurs at a maximum rateand near the bottom of the core (from about 35 to 47 Ma). Wetherefore think that a major portion of the changes observed inthe Mg/Ca and Sr/Ca ratios of the carbonate fine fraction areof diagenetic origin. Nevertheless, as mentioned previously, dia-genesis should not eliminate original variations; rather, it shouldonly dampen their observed amplitudes.

Iron and ManganeseIron (Fig. 5) and manganese (Fig. 6) are both transition met-

als with similar ionic radii and charges; thus, both would be ex-pected to have similar distribution coefficients in calcite. Bothelements display variable redox states in seawater and shallowpore waters. Manganese (II) is thought to be the dominant re-dox state of manganese in seawater (even though it is thermo-dynamically unstable), whereas iron (III) is thought to be domi-nant for iron (Bruland, 1983). Neither element is well mixed inseawater. Generally, manganese is more abundant than iron,particularly in surface waters (Bruland, 1983). Both iron andmanganese are greatly enriched in pore waters as a result of mi-crobial oxidation of organic carbon. Although low levels of

manganese and iron may be incorporated into biogenic calcite,it seems much more likely that diagenesis will play the greatestrole in their incorporation into carbonate sediments. Becausemanganese is reduced more easily and rapidly than oxidizediron, pore-water maxima of these elements are attained at dif-ferent sub-bottom depths (e.g., Froehlich et al., 1979). Manga-nese is also generally more abundant than iron in pore waters(Froehlich et al., 1979; Sawlan and Murray, 1983). At Site 709there is good evidence for sulfate reduction in the pore waters(Swart and Burns, this volume) and pyrite precipitation in thesediments; hence, dissolved iron concentrations in the pore wa-ters should only be high in the upper few meters of the sedimentcolumn.

From the above discussion, we conclude that it is most likelythat iron and manganese concentrations in pelagic carbonatesediments are controlled by diagenetic processes. Their micro-bial reduction allows them to be solubilized and made availablefor incorporation into reprecipitated diagenetic calcite or per-haps into trace amounts of a mixed iron, manganese, magne-sium, and calcium carbonate phase. Because of their slightlydifferent redox chemistries, peak concentrations of these ele-ments in the carbonate phase may be offset in terms of sub-bot-tom depth, and concentration peaks for both elements may notexactly coincide.

The manganese data (Fig. 6) may be the best means for as-sessing the nature of the slumping process described by D. Rioand H. Okada (Shipboard Scientific Party, 1988b, p. 468). Theydetermined that the interval between Sections 115-709C-15H-4(136 mbsf) and 115-709C-18H-3 (164 mbsf) was disturbed. Al-

672


4000

3000 -

EQ.Q.

(I)

2000-

1000 -

40 5 0

Figure 8. Three-point moving average of the deconvoluted strontium values (line) and strontium three-point moving aver-age (dashed line) vs. age.

Table 2. The aragoniticfraction of the total car-bonate sedimentation cal-culated to be necessary toexplain the observed Sr/Ca ratio of pelagic car-bonate sediments.

(Sr/Ca)5lv

0.00300.00500.00800.00870.01000.0110

(Sr/Ca)

0.000690.001150.001840.002000.002300.00250

/A

0.670.310.100.070.030.01

Note: (Sr/Ca)5VV = the ratioin seawater (0.0087 is themodern-day value); Sr/Ca = the ratio observedin pelagic carbonate sed-iments; and fA = thefraction of aragonite ofthe carbonate sediments.

though the whole interval was stratigraphically disordered, theyconcluded that individual nannofossil assemblages were coher-ent. Thus, reworking must have resulted from discrete slumpblocks. S. Robinson (pers. comm., 1989; Shipboard ScientificParty, 1988b) used magnetic susceptibility to correlate between

Holes 709A and 709C. From his work, we can place the top ofthe slumped interval in Hole 709A at about 135 mbsf. Presum-ably, the bottom of this interval in 709A is at about 163 mbsf.This depth interval corresponds to a time interval from about13.0 to 17.5 Ma.

We interpret the large increase in manganese concentration(Fig. 6) at 19.0 Ma and subsequent decrease at 18.1 Ma as thetop of the intact sedimentary record. This corresponds closely,but not perfectly, with a large iron concentration peak and, im-portantly, with a sharp decrease in the 13C/12C ratio. The rapidburial of this sediment by several meters of slumped sedimentcreated an effective barrier to the diffusion of dissolved oxygeninto the pore waters and rapidly accelerated the rates of manga-nese and iron reduction and, possibly, of sulfate reduction. Thetwo following manganese cycles (with peaks at 13.8 and 16.3Ma in Fig. 6) are interpreted as repeated sequences with similardiagenetic histories. If the enhanced concentrations of manga-nese, iron, and magnesium actually indicate the precipitation ofone or two separate mixed carbonate mineral phases during earlydiagenesis, these phases would be expected to have very low car-bon isotopic compositions. The presence of small amounts ofsuch minerals could explain the associated decreases of 513Cover the same depth interval. Two phases are proposed becauseiron and manganese correlate very closely over the slumped in-terval, but they do not correlate closely with magnesium andstrontium. Small amounts of dolomite were identified in smearslides of sediments from the slumped interval and near the baseof the hole from 44.5 Ma to the core bottom.

673


2 -

Q)D.

oco ^ 'VA/W\A/I/V

Age (Ma)

Figure 9. Oxygen isotopic composition of calcite in equilibrium with present-day pore waters, three-point moving av-erage of the oxygen isotopic composition of the carbonate fine fraction (dotted line), and presumed initial isotopiccomposition of the fine fraction vs. age. See text for further details.

CONCLUSIONS

A quantitative evaluation of the mechanisms involved in pro-ducing minor element concentration variations in deep-sea car-bonate sediments is not yet possible. Such determinations willrequire a careful comparison of several sets of chemical datameasured on deep-sea carbonate sediments from different oceansand in different environmental settings. Nevertheless, we candraw some useful conclusions from the present study.

A correlation is observed between high sea-level stands andhigh values of 513C. It is suggested that this supports the hy-pothesis proposed by Shackleton (1987) that <513C is controlledby the fraction of organic carbon that is buried: more organiccarbon is buried during high stands and oceanic-dissolved in-organic carbon becomes isotopically heavier. Carbon isotopicchanges mirror the long-term (Eocene to Present) decrease insea level resulting from slower rates of seafloor spreading. Su-perimposed on this long-term change are several shorter termvariations.

There appears to be some correlation between high sea-levelstands and high CaCO3 contents in the sediment (this relation-ship is discussed more thoroughly elsewhere in this volume).This relationship implies a shelf-basin fractionation mechanismfor carbonate accumulation: more carbonate is deposited on theshelves during high stands and the CCD rises.

Because much of the shelf carbonate is aragonite, more stron-tium is buried during sea-level highstands and oceanic Sr/Ca ra-

tios decrease. As a result, we observed an inverse relationshipbetween sea level and strontium concentration.

Diagenesis plays an important role in many of the observedparameters. Burial diagenesis clearly affects strontium and mag-nesium concentrations and undoubtedly oxygen isotopic ratiosas well. Microbially mediated oxidation of organic carbon inthe slumped sediments causes early diagenetic reactions such asreduction of iron and manganese oxides, sulfate reduction, andprobably precipitation of mixed carbonate minerals. This causeslarge variations of iron, manganese, magnesium, and carbonisotopic ratios within the slumped interval.

ACKNOWLEDGMENTSWe thank the crew and the rest of the shipboard technical

and scientific parties, especially the co-chief scientists, J. Back-man and R. A. Duncan. Obviously, the help of all of these peo-ple was essential for the present study. We acknowledge the sup-port by USSAC and NSF-OCE-87-18382 (to P. A. Baker).

REFERENCES

Baker, P. A., 1985. Pore-water chemistry of carbonate-rich sediments,Lord Howe Rise, southwest Pacific Ocean. In Kennett, J. P., von derBorch, C. C , et al., Init. Repts. DSDP, 90: Washington (U.S. Govt.Printing Office), 1249-1256.

Baker, P. A., Gieskes, J. M., and Elderfield, H., 1982. Diagenesis ofcarbonates in deep-sea sediments—evidence from Sr/Ca ratios andinterstitial dissolved Sr2+ data. J. Sediment. Petrol., 52:71-82.

674


Boyle, E. A., and Keigwin, L. D., 1982. Deep circulation of the NorthAtlantic over the last 200,000 years: geochemical evidence. Science,218:784-787.

Bruland, K. W., 1983. Trace elements in sea-water. In Riley, J. P., andChester, R. (Eds.), Chemical Oceanography (Vol. 8): London (Aca-demic Press), 157-220.

Cronblad, H. G., and Malmgren, B. A., 1981. Climatically controlledvariation of Sr and Mg in Quaternary planktonic foraminifera. Na-ture, 291:61-64.

Delaney, M. L., Be, A.W.H., and Boyle, E. A., 1985. Li, Sr, Mg, andNa in foraminiferal calcite shells from laboratory culture, sedimenttraps, and sediment cores. Geochim. Cosmochim. Ada, 49:1327-1341.

Delaney, M. L., and Boyle, E. A., 1986. Lithium in foraminiferal shells:implications for high-temperature hydrothermal circulation fluxesand oceanic crustal generation rates. Earth Planet. Sci. Lett., 80:91-105.

, 1988. Tertiary paleoceanic chemical variability: unintendedconsequences of simple geochemical models. Paleoceanography, 3:137-156.

Elderfield, H., 1986. Strontium isotope stratigraphy. Palaeogeogr., Pa-laeoclimatol., Palaeoecol., 57:71-90.

Elderfield, H., and Gieskes, J. M., 1982. Sr isotopes in interstitial wa-ters of marine sediments from Deep Sea Drilling Project cores. Na-ture, 300:493-497.

Froehlich, P. N., Klinkhammer, G. P., Bender, M. L., Luedtke, N. A.,Heath, G. R., Cullen, D., Dauphin, P., Hammond, D., Hartman,B., and Maynard, V., 1979. Early oxidation of organic matter in pe-lagic sediments of the eastern equatorial Atlantic: suboxic diagene-sis. Geochim. Cosmochim. Ada, 43:1075-1090.

Graham, D. W., Bender, M. L., Williams, D. F., and Keigwin, L. D.,1982. Strontium-calcium ratios in Cenozoic planktonic foraminif-era. Geochim. Cosmochim. Ada, 46:1281-1292.

Haq, B. U., Hardenbol, J., and Vail, P. R., 1987. Chronology of fluctu-ating sea levels since the Triassic. Science, 235:1156-1167.

Katz, A., Sass, E., Starinsky, A., and Holland, H. D., 1972. Strontiumbehavior in the aragonite-calcite transformation: an experimentalstudy at 40°-98°. Geochim. Cosmochim. Ada, 36:481-496.

Kinsman, D. J., 1969. Interpretation of Sr2+ concentrations in carbon-ate minerals and rocks. J. Sediment. Petrol., 39:486-508.

Lea, D., and Boyle, E. A., 1989. Barium content of benthic foraminif-era controlled by bottom-water composition. Nature, 338:751-753.

Miller, K. G., and Fairbanks, R. G., 1985. Oligocene to Miocene carbonisotope cycles and abyssal circulation changes. In Sundquist, E. T.,and Broecker, W. S. (Eds.), The Carbon Cycle and Atmospheric

CO2: Natural Variations Archean to Present. Am. Geophys. UnionMonogr., 32:469-486.

O'Neil, J. R., Clayton, R. N., and Mayeda, T. K., 1969. Oxygen isotopefractionation in divalent metal carbonates. J. Chem. Phys., 51:5547-5558.

Palmer, M. R., and Edmond, J. M., 1989. The strontium isotope budgetof the modern ocean. Earth Planet. Sci. Lett., 92:11-26.

Pitman, W. C , 1978. Relationship between eustacy and stratigraphic se-quences of passive margins. Geol. Soc. Am. Bull., 89:1389-1403.

Renard, M., 1986. Pelagic carbonate chemostratigraphy (Sr, Mg, 18O,13C). Mar. Micropaleontol., 10:117-164.

Richter, F. M., and DePaolo, D. J., 1988. Diagenesis and Sr isotopeevolution of seawater using data from DSDP 590B and 575. EarthPlanet. Sci. Lett., 90:382-394.

Savin, S. M., 1977. The history of the Earth's surface temperature dur-ing the past 100 million years. Ann. Rev. Earth Planet. Sci., 5:319-355.

Sawlan, J. J., and Murray, J. W., 1983. Trace metal remobilization in theinterstitial waters of red clay and hemipelagic marine sediments.Earth Planet. Sci. Lett., 64:213-230.

Shackleton, N. J., 1987. The carbon isotope record of the Cenozoic:history of organic carbon burial and of oxygen in the ocean and at-mosphere. In Brooks, J., and Fleet, A. J. (Eds.), Marine PetroleumSource Rocks: London (Blackwell Sci. Publ.). Geol. Soc. Spec. Publ.,26:423-434.

Shackleton, N. J., and Hall, M. A., 1984. Carbon isotope data fromLeg 74 sediments. In Moore, T. C , Jr., Rabinowitz, P. D., et al.,Init. Repts. DSDP, 74: Washington (U.S. Govt. Printing Office),613-619.

Shipboard Scientific Party, 1988a. Explanatory notes. In Backman, J.,Duncan, R. A., et al., Proc. ODP, Init. Repts., 115: College Station,TX (Ocean Drilling Program), 17-42.

Shipboard Scientific Party, 1988b. Site 709. In Backman, J., Duncan,R. A., et al., Proc. ODP, Init. Repts., 115: College Station, TX(Ocean Drilling Program), 459-588.

Vincent, E.. and Berger, W. H., 1985. Carbon dioxide and polar coolingin the Miocene: the Monterey Hypothesis. In Sundquist, E. T., andBroecker, W. S. (Eds.), The Carbon Cycle and Atmospheric CO2:Natural Variations Archean to Present. Am. Geophys. UnionMonogr., 32:455-486.

Date of initial receipt: 12 January 1989Date of acceptance: 5 January 1990Ms 115B-178

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36. minor element and stable isotopic ......minor element and stable isotopic composition of the...

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