Himalayan uplift and osmium isotopes in oceans and rivers

Geochimica et Cosmochimica Acta, Vol. 63, No. 23/24, pp. 4005– 4012, 1999 Copyright © 1999 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/99 $20.00 ⫹ .00

Pergamon

PII S0016-7037(99)00305-1

Himalayan uplift and osmium isotopes in oceans and rivers M. SHARMA1,* G. J. WASSERBURG,2 A. W. HOFMANN,1 and G. J. CHAKRAPANI3 1

Max–Planck–Institut fu¨r Chemie, Postfach 3060, D-55020 Mainz, Germany The Lunatic Asylum of the Charles Arms Laboratory, Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA 3 Department of Earth Sciences, University of Roorkee, Roorkee 247 667, India

2

Abstract—Previous studies have shown that 187Os/188Os in seawater has become increasingly radiogenic over the last 40 Ma in a manner analogous to strontium. This rapid rise in the marine 187Os/188Os over the last 17 Ma has been attributed to an increase in the bulk silicate weathering rates resulting from the rise of the Himalayas and/or selective weathering and erosion of highly radiogenic organic rich ancient sediments. The key test of this hypothesis is the 187Os/188Os and the total osmium concentration of the Himalayan rivers. We report the concentration and isotopic composition of osmium in the Ganges, the Brahmaputra, and the Indus rivers. The 187Os/188Os of the Ganges close to its source (at Kaudiyal, 30°05⬘N, 78°50⬘E) is 2.65 and [Os] ⫽ 45 fM/kg. A second sample of the lower reaches of the Ganges at Patna (25°30⬘N, 85°10⬘E) gives 187 Os/188Os ⫽1.59 and [Os] ⫽ 171 fM/kg. The 187Os/188Os of the Brahmaputra at Guwahati (26°10⬘N, 91°58⬘E) is 1.07 and [Os] ⫽ 52 fM/kg. A sample of the Indus (Besham, 34°55⬘N, 72°51⬘E) has a 187Os/188Os of 1.2 and [Os] ⫽ 59 fM/kg. We infer that the Himalayas do not provide either a high flow of osmium or a highly radiogenic osmium component to the oceans. The overall trend for osmium and strontium could be explained by a regularly increasing input of global continental weathering sources but the Himalayas themselves appear not to be the dominant source. Copyright © 1999 Elsevier Science Ltd continental (rivers and aerosols) sources. On average, the 187 Os/188Os ratios of the mantle and cosmic materials are nearly identical (⬃0.13 with 187Re/188Os ⫽ 0.36) and about 10 times lower than that of average continental matter (1.26 with 187 Re/188Os ⫽ 48). Because of the large isotopic contrast between the mantle, the cosmic, and the continental sources, the Os isotopes in the oceans should be an ideal tracer of continental weathering. The mean residence time of Os in the oceans (␶៮ Os) is of the order of 104 years (Levasseur et al., 1998; Sharma et al., 1997). It is (1) short enough to assume achievement of a steady state on a million year time-scale, and (2) longer than the ocean mixing time (⬃1500 yr) such that Os isotopic evolution of the deep oceans may track changes induced by significant changes in the contributing sources. Thus, if a large continental region undergoes extensive weathering and erosion following an uplift, increased input of continental Os should be reflected in the long term record of Os isotopes in the chemical precipitates (Mn crusts, metalliferous sediments) that sequester Os in the oceans. The past Os isotopic record of the oceans, as preserved in the marine sediments, is complicated by Os inputs from sources other than seawater. Extraction of seawater component of Os from ocean sediments, therefore, uses differential acid leaching of the samples to measure the “hydrogenous Os” that is attributed to seawater. This process is fraught with difficulties (Oxburgh, 1998; Pegram et al., 1992; Pegram and Turekian, 1999; Peucker–Ehrenbrink et al., 1995; Ravizza, 1993). More reliable analyses of the hydrogenous component of marine sediments have indicated that the seawater 187Os/188Os has increased in the last 66 Ma, except between 29 and 17 Ma when there was a little change (Fig. 1b) (Pegram et al., 1992; Pegram and Turekian, 1999; Peucker–Ehrenbrink et al., 1995; Ravizza, 1993). The rate of increase of 187Os/188Os ratio in the last 17 Ma has been quite rapid (0.019/Ma; Fig. 1b) and has been interpreted to be a consequence of (1) increased weathering in

1. INTRODUCTION

We report the first measurements of the concentration and isotopic composition of osmium (Os) in the Ganges, the Brahmaputra, and the Indus rivers. As evident from the sediments deposited in the Indogangetic alluvium and the Bengal and the Indus fans, extensive weathering and erosion of the Himalayas and the Tibetan Plateau occurred with the uplift that began about 55 Ma ago. A significant increase in continental weathering appears to have taken place beginning 17 million years ago as reflected by a factor of 10 increase in sediment accumulation rate in the Indian Ocean (Turekian, 1996) and by the marine Sr isotopic record (Fig. 1a). At the same time, the ␦18O record of benthic foraminifera points to a dramatic global cooling trend, ostensibly related to a decrease in the level of greenhouse gases (most dominantly CO2) in the earth’s atmosphere. These changes have been attributed to the onset of a substantial uplift in the Himalayan region leading to a shift towards more weathering of continental silicates (Raymo and Ruddiman, 1992). We use Os isotopes in the Himalayan rivers to evaluate the extent to which these rivers may have contributed to the flow of ions required to cause the observed changes. Extensive discussion of various possible causes of the isotopic shifts in Sr and Os have been presented by many investigators (cf. Krishnaswami et al., 1992; Edmond, 1992; Richter et al., 1992; Godde´ris and Franc¸ois, 1995; Peucker–Ehrenbrink et al., 1995; Reisberg et al., 1997; McCauley and DePaolo, 1997; Turekian and Pegram, 1997). An initial report of this study was presented by Sharma et al. (1998). Radiogenic 187Os is produced from the ␤-decay of 187Re with a half-life of 42 billion years. The Os isotopic composition of the oceans is governed by inputs from mantle (hydrothermal and submarine weathering), cosmic (dust and meteorites), and

* Author to whom correspondence should be addressed. 4005

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Os and Os concentration of the Himalayan rivers with that of the average continental crust and other major rivers, we can assess whether they have singular characteristics that could govern the seawater evolution. Direct measurements of the Os isotopic composition of the Himalayan rivers have not previously been made. Indirect measurements on sediment leaches from the Ganges at Patna and Varanasi suggest that the 187Os/188Os of the river is high (⫽1.94) (Pegram et al., 1994). This is also suggested by analysis of a sediment collected in Bangladesh (⫽2.6) (PiersonWickmann et al., 1998). Analyses of Mn nodules preserved in the paleosoils of the Ganges flood plain also suggest that waters draining into some segments of this river in the past contained highly radiogenic Os (Chesley et al., 1998). These observations have led workers to deduce that the Os input from the Ganges into the oceans should be much more radiogenic than the average upper continental crust (Pegram et al., 1994; Pegram et al., 1992; Pierson–Wickmann et al., 1998; Turekian and Pegram, 1997). In contrast, the 187Os/188Os ratio of the Brahmaputra and the Indus are assumed to be influenced more by the weathering and erosion of ophiolitic assemblages (with 187Os/ 188 Os ⬍⬍ 1) exposed along the Indus–Tsangpo suture zone (Fig. 2). The sediment leach data also appear to suggest a low value for the Brahmaputra (⫽ 0.6) (Pierson–Wickmann et al., 1998). Analyses of Mn nodules preserved in the paleosoils of the Indus flood plain suggest that the 187Os/188Os ratio of the waters draining into this river has been similar to the average upper continental crust (Chesley et al., 1998). 2. SAMPLING AND ANALYTICAL METHODS

Fig. 1. (a) Variation in the seawater Sr isotopic composition of over the last 70 Ma based on compilations by McCauley and DePaolo (1997). (b) Variation in the seawater Os isotopic composition over the last 70 Ma based on compilations by Peucker-Ehrenbrink et al. (1995) (solid line) compared with the data from a single core (dashed line with solid circles) (Pegram and Turekian, 1999). (c) R-value over the last 70 Ma corresponding to the two-component model discussed in the text. The following parameters were used: (187Os/188Os)C ⫽ 1.26, (187Os/ 188 SW Os)H ⫽ 0.126, COs ⫽ 57 fm/kg, (187Os/188Os)SW ⫽ 1.05, (87Sr/ 86 Sr)C ⫽ 0.7119, (87Sr/86Sr)H ⫽ 0.7030, (87Sr/86Sr)SW ⫽ 0.70915. ROs was calculated using the data from (Pegram and Turekian, 1999). Both ROs and RSr are within a factor of two of each other. In detail, however, ROs decreases by a factor of four over the last 17 Ma, whereas RSr remains constant (see inset).

the Himalayas, and/or (2) increased weathering of ancient organic rich sediments (black shales) possessing extremely radiogenic 187Os/188Os (Pegram et al., 1992; Ravizza, 1993; Ravizza and Esser, 1993). Increased weathering of black shales in the Himalayas may have also influenced the Os isotopes in the oceans during the Neogene (Pegram et al., 1992; Peucker– Ehrenbrink et al., 1995; Turekian and Pegram, 1997). If the Himalayan orogeny is directly responsible for the observed increase in the seawater 187Os/188Os in the last 17 Ma, we can evaluate its impact by examining the Os isotopes in major Himalayan rivers. Also, by comparing the present-day 187Os/

Samples of three major Himalayan rivers, the Ganges, the Brahmaputra, and the Indus were collected in acid-cleaned, high-density, polyethylene bottles. The Ganges and the Brahmaputra were sampled in November 1997. These samples were filtered by using 0.45-␮m filters and acidified with ultra-pure HCl to a pH of 2.0 within 12 h of collection. The Indus sample was collected in May 1998. This sample was filtered and acidified in the laboratory about 2 weeks after collection. The locations of these samples are marked in Figure 2. Of the two Ganges samples, one was collected close to the source (at Kaudiyal, 30°05⬘N, 78°50⬘E) and other near the mouth of the river (at Patna, 25°30⬘N, 85°10⬘E). These sites were chosen to assess the Os contribution from the Ganges floodplain. Before descending into the floodplain, the Ganges (and its precursors) flows through an area consisting of carbonates, shales (pyritous at places), quatzites, and low- to highgrade gneisses (Sarin et al., 1992). The Ganges at Kaudiyal was sampled to examine to the influence of this variegated packet of lithologies on the Os dissolved in the Ganges headwaters. The Brahmaputra sample was collected upstream of the city of Guwahati (26°10⬘N, 91°58⬘E). The Indus sample comes from Besham (34°55⬘N, 72°51⬘E) along the Karakoram highway in Pakistan. The Brahmaputra and the Indus samples were collected to assess the impact of ophiolite weathering on the Os dissolved in the water. We consider the Brahmaputra (at Guwahati) and the Ganges (at Patna) samples to be representative of these rivers before their confluence and descent into the Bay of Bengal. As the samples were collected during medium to low flow stages, the measured Os concentration is probably higher than during the flood season, but this does not affect our conclusions. Determination of Os isotopes in waters is a challenging analytical problem due to low concentration and multiple oxidation states of Os in waters. Whereas the development in negative thermal ionization mass spectrometry has provided a sensitive technique to measure subfemtomolar quantities of Os, the issue of tracer–sample equilibration for water analysis has been a difficult one to resolve. Sharma et al. (1997) used a ‘cold’ SO2 reduction step to equilibrate tracer with sample but noted a lack of complete equilibration. Levasseur et al. (1998) developed a ‘hot’ Br2–CrO3–H2SO4 oxidation technique that

Osmium Isotopes in the Himalayan Rivers

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Fig. 2. Generalized map of the Indian subcontinent marking sample locations. The first number in the parentheses against each location gives the measured Os concentration and the second the Os isotopic composition.

permits full tracer-sample equilibration. They showed that although they could confirm the isotopic composition in seawater found by Sharma et al. (1997), the concentration of Os in seawater was 3 times higher than their reported value. It is apparent that a large portion of Os in seawater is bound to organic matter making tracer–sample equilibration nearly impossible at low temperatures. In approaching the measurement of Os in the Himalayan river samples, we have used a variation of ‘hot’ oxidation technique. This procedure, in brief, is as follows. Between 50 and 100 mL of water is taken in a glass tube and frozen in dry ice-alcohol mixture. The tracer is then added into the tube and frozen following which 100 ␮L of 40% CrO3 in HNO3 is added. The tube is then sealed and placed in an oven at 180°C for 40 h. Under these conditions, Os is oxidized to OsO4 and there is complete tracer– sample equilibration (see below). The tube is then chilled to 4°C and scored open. The liquid is poured into a distillation apparatus and OsO4 distilled and trapped in chilled HBr. The hexbromoosmate fraction is then dried and further cleaned using micro-distillation (Birck et al., 1997). The total 188Os blank of this procedure is 0.013 ⫾ 0.006 fM with 187 Os/188Os of ⫽ 0.27 ⫾ 0.14. The total Os yield of this procedure is 90%. Os isotopes were measured using procedures outlined in Sharma et al. (1997). The kinetics of isotopic exchange between the tracer and sample Os is rather slow even at elevated temperatures (Levasseur et al., 1998). Using the Ganges sample at Patna, experiments were carried out to determine the duration for which the sample should be kept in the oven at 180°C to ensure complete tracer-sample equilibration (Levasseur et al., 1998). Samples were kept in oven from 4 to 84 h following which Os was separated and measured. The results of these experiments are presented in Figure 3 in a time (duration of oxidation step) versus concentration (in fM/kg) plot. In such a plot the measured concentration should increase initially and then become constant with time. However, complications in such a plot may arise at lower time steps if the temperature achieved inside the glass tube is less than the final temperature. For the apparatus used, we determined that a minimum of 4 h is necessary to achieve a temperature of 180°C in the liquid inside the tube (“zero-time”). It is apparent in Figure 3 that the measured concentration shows the expected increase with the time the sample is kept in oven and is nearly constant for times of from 40 to 84 h. Although self-consistent, these experiments cannot per se demonstrate complete equilibration as we need to independently know the concentration of Os in this sample! To address this issue, we used a chemical

separation protocol developed for equilibrating silicate/sulfide/metal Os with tracer Os in a closed system (Shirey and Walker, 1995). This

Figure 3. Plot showing the time required to equilibrate Os in the water samples with 190Os tracer isotope. The kinetics are slow and require oxidation of Os to OsO4 in a closed system at elevated temperatures (Levasseur et al., 1998). Seven aliquots of one sample were taken and oxidized at 180°C for times ranging from 4 to 84 h. The measured concentration of the samples did not change after 40 h. Another aliquot (9 g) was processed using the protocols that allow equilibration of Os in sulfides/silicates/metal samples with the tracer and is identical to that for samples oxidized over 40 h. The “zero time experiment” refers to the aliquot that was removed from the oven immediately after the internal temperature had reached 180°C.

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M. Sharma et al. Table 1. Osmium isotopes in the dissolved loads of the Himalayan rivers compaed with other rivers. Drainage (km3/yr)

Ganges Patna

Average Kaudiyal

Weight (kg)

(Total 188Os)ca (fmol)

Concentration (fmol/kg)

0.11354 0.05407 0.06623 0.10602 0.10661

2.1 1.1 1.3 2.0 2.0

0.10883 0.045825

(187Os/188Os)mb

(187Os/188Os)ca

(187Os/186Os)ca

165.3 ⫾ 0.4 176.9 ⫾ 0.9 172.5 ⫾ 0.7 165.5 ⫾ 0.5 173.3 ⫾ 0.5 171 ⫾ 5

1.586 ⫾ 0.007 1.521 ⫾ 0.007 1.573 ⫾ 0.003 1.619 ⫾ 0.002 1.572 ⫾ 0.005

1.59 ⫾ 0.01 1.54 ⫾ 0.01 1.59 ⫾ 0.01 1.63 ⫾ 0.01 1.58 ⫾ 0.01 1.59 ⫾ 0.03

13.3 ⫾ 0.1 12.8 ⫾ 0.1 13.3 ⫾ 0.1 13.6 ⫾ 0.1 13.2 ⫾ 0.1 13.2 ⫾ 0.3

0.52 0.21

46.9 ⫾ 0.4 43 ⫾ 1 45 ⫾ 2

2.610 ⫾ 0.005 2.58 ⫾ 0.01

2.63 ⫾ 0.03 2.73 ⫾ 0.09 2.6 ⫾ 0.06

22.3 ⫾ 0.3 22.8 ⫾ 0.7 22.5 ⫾ 0.5

0.10537 0.06963

0.67 0.44

52.7 ⫾ 0.5 52.1 ⫾ 0.7 52 ⫾ 1

1.080 ⫾ 0.003 1.050 ⫾ 0.005

1.10 ⫾ 0.01 1.07 ⫾ 0.01 1.07 ⫾ 0.02

9.1 ⫾ 0.1 9.0 ⫾ 0.1 9.1 ⫾ 0.1

0.12058 0.05120

0.86 0.35

60.8 ⫾ 0.4 57 ⫾ 1 59 ⫾ 1 (45)c (15)c (14)c (44)c

1.190 ⫾ 0.003 1.181 ⫾ 0.004

1.20 ⫾ 0.01 1.21 ⫾ 0.02 1.20 ⫾ 0.02 1.25 1.72 1.05 1.28

10.1 ⫾ 0.1 10.1 ⫾ 0.2 10.1 ⫾ 0.2 10.4 14.4 8.8 10.7

590 (459)

(24)

Average Brahmaputra

626 Guwahati Average

Indus

175 Besham Average

Mississippic Columbiac Connecticutc Vistulac

580 172 4 33

Corrected for procedural blank of 0.013 ⫾ 0.006 fM of 188Os with 187Os/188Os of 0.27 ⫾ 0.14. Measured ratio corrected for isotope fractionation and for tracer contributions to the mixture. The power law was used for mass fractionation correction, assuming that the mass fractionation entails the OsO⫺ 3 molecular ion. The data are normalized for mass-dependent isotope fractionation by assuming the arbitrary reference value 192Os/188Os ⫽ 3.08267. for reduction of the Os oxide data, the following oxygen isotopic composition was used: 17O/16O ⫽ 0.00037 and 18O/16O ⫽ 0.002047. Isotopic composition of the Os normal used for data reduction: 186Os/188Os ⫽ 0.11973, 187 Os/188Os ⫽ 0.10677, 189Os/188Os ⫽ 1.21969, 190Os/188Os ⫽ 1.98146, 192Os/188Os ⫽ 3.08267. c Data using low temperature equilibration technique. May be subject to revision (Sharma and Wasserburg, 1997). a

b

procedure introduces relatively larger blanks but is adequate for this purpose. About 9 g sample was taken in a Carius tube and frozen. Following this the tracer solution, HCl and HNO3 were poured in and the tube sealed shut and kept in oven for 60 h at 240°C. The Carius tube was then chilled to 4°C and scored open. Following this, OsO4 was extracted by using liquid Br2 (Birck et al., 1997). The result of this sample is also plotted in Figure 3. The concentration of samples kept in oven for more than 40 h is within error of this independent estimate. Following these experiments, all samples reported were kept in oven for 40 h. 3. RESULTS AND INTERPRETATION

The results are listed in Table 1, which also gives the Os isotopic composition and concentration of the other rivers measured so far (Sharma and Wasserburg, 1997). The average concentrations and 187Os/188Os ratios are shown on the location map in Figure 2. There is about 2–3% variability in the Os concentration and isotopic composition of the Ganges at Patna. A similar variability is displayed by the other samples. The Os concentration in the Himalayan rivers varies from 45 fM/kg to 171 fM/kg and the 187Os/188Os from 1.07 to 2.65. Significantly, the Os concentration of the rivers draining the Himalayan highlands (the Ganges at Kaudiyal, the Brahmaputra, and the Indus) is between 45 and 59 fM/kg. Moreover, close to its source, the Ganges is quite radiogenic with low concentration of Os, becoming substantially less radiogenic with high Os concentration after it has traversed most of its flood plain. The significance of this change, which may be seasonal and depen-

dent on the relative flows of various tributaries, is not clear at present. For the discussion below, we note that the Os isotopic composition of the upper Ganges is 2.65 and that in the lower reaches (Patna), it is 1.59, which should approach the value of 187 Os/188Os that is fed into the ocean. This value is somewhat more radiogenic than the canonical average upper continental crust (⫽1.26) (Esser and Turekian, 1993) and the range of values determined for the other rivers (Table 1). The Os isotopic composition of the Ganges estimated from the sediment leaches (Pegram et al., 1994; Pierson–Wickmann et al., 1998) is within the range of the directly measured values. Pierson–Wickmann et al. (1998) analyzed rock samples from the principal lithologies in the Himalayas as well as river sediments. They found that while the rocks and sediments derived from the Tethyan Sedimentary Series and the High Himalayan Crystalline Series contain low amounts of osmium (10 to 100 ppt) with 187Os/188Os ⫽ 0.9 to 2.6, the Lesser Himalayan shales are very radiogenic (187Os/188Os ⫽ 5 to 7.8) and contain between 100 and 550 ppt Os. Indeed, the high 187 Os/188Os of the Ganges at Kaudiyal could result from small contributions of such organic rich sedimentary rocks. The observed elevated 187Os/188Os ratios also could result from incongruent weathering (Pegram et al., 1994). Nonetheless, it is evident that a highly radiogenic Os source does not dominate the Himalayan rivers. An important source of nonradiogenic Os (187Os/188Os ⬍⬍1) in the High Himalayas is the ophiolites

Osmium Isotopes in the Himalayan Rivers

exposed along the Indus–Tsangpo suture zone (Fig. 2). We also do not observe a dominance of this source in either the Indus or the Brahmaputra that flow through this region. The above results indicate that, on average, all three Himalayan rivers have 187Os/188Os ratios close to other rivers measured so far (Table 1) and to the average upper continental crust (Esser and Turekian, 1993). The concentrations are also roughly compatible with the other rivers (Levasseur, 1999). The significance of these results is that the Os dissolved in these rivers appears to reflect averaging of large segments of upper continental crust by erosion and does not point to the dominance of a singular lithology such as peridotites associated with ophiolites or black shales. 4. DISCUSSION

4.1. Himalayan Osmium Contributions Several workers have concluded that the change in the marine Os record during the past 17 Ma (Fig. 1b) must be due to the Himalayan orogeny, resulting from either increased weathering intensity or from the erosion of highly radiogenic black shales (Pegram et al., 1992; Peucker–Ehrenbrink et al., 1995; Turekian and Pegram, 1997). Another view is that the observed increase in the seawater 187Os/188Os ratio is the result of a general increase in the weathering of organic matter not necessarily related to the Himalayan uplift (Ravizza, 1993; Ravizza and Esser, 1993). The present-day seawater 187Os/ 188 Os ratio is the highest in the last 17 Ma. It follows that if selective weathering and erosion of black shales from the Himalayas were the main cause of the observed increase in the seawater 187Os/188Os ratio, the present-day rivers (neglecting the effect of the residence time) would be extremely radiogenic. However, as concluded above, a highly radiogenic Os source does not dominate the Himalayan rivers. If increased weathering intensity in the Himalayan region itself has contributed significantly during the past 17 Ma to the increased Os flow into the oceans, the following conditions should be satisfied. These conditions are identical to those invoked by Richter et al. (1992) for evaluating the Sr contribution from the Himalayas. They are: (1) the increase in the marine 187Os/188Os should occur during the time of uplift in the Himalayas; (2) the flow of Os from rivers draining this region should, on average, have an isotopic composition much more radiogenic than the other sources of continental Os and/or carry a significant fraction of total Os being added to the oceans; and (3) the amount of Os eroded since the uplift must be sufficient to account for the integrated Os flow. The rapid rise in the marine 187Os/188Os ratio during the last 17 Ma follows closely a period of rapid uplift of the western Tibetan Plateau (cf. Harrison et al., 1992). A nearly constant and continuous rise in the 187Os/188Os in the last 17 Ma would require a regular increase in erosion rates during this time. The inferred increase in sediment accumulation rate in the Indian Ocean (Turekian, 1996) appears to satisfy this requirement. As the Himalayan rivers, on average, have 187Os/188Os close to the other rivers as well as the upper continental crust, the main issue then is that of the amount of Os provided by Himalayan weathering relative to the total amount of Os buried during the last 17 Ma. The concentration of Os in the oceans is 57 fM/kg (Levasseur et al., 1998). Assuming a quasi-steady state and a ␶៮ Os of

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4 ⫻ 104 year (Sharma et al., 1997), this gives an annual flow rate of 1995 mol Os/yr into and out of the oceans. The average Os concentration of the Himalayan rivers is 100 fM/kg (Table 1). As these rivers constitute 3.3% of the annual river flow, they supply 6.2% of the present inflow of Os to the oceans. Moreover, as the average 187Os/188Os of the rivers is 1.44, they contribute about 13% of the present increase in the seawater 187 Os/188Os. It follows that a factor of 2 increase in either the Os flow or the isotopic composition of the Himalayan river input will not significantly affect the seawater Os isotopic evolution. Clearly, the flow of Os to the deep sea delivered by the Himalayan rivers has large uncertainties, as it critically depends on the assumption that: (1) Os dissolved in the riverine transport is not trapped within the estuaries; and (2) there is no release of Os from the riverine sediment during or after its burial in the oceans. By examining the Os isotopic composition of sediments in the Long Island Sound, Williams et al. (1997) suggested that a large amount of Os gets trapped within the estuaries. In contrast, direct measurements of the Columbia River estuary appear to suggest that Os mixes conservatively in the estuaries (Porcelli et al., 1998). This issue is at present far from being resolved. In any case, a smaller delivery of Os does not weaken the argument pertaining to the transport of riverine Os from the Himalayas. We can also estimate an upper bound on the Os flow from the Himalayas from the bulk sediment deposited in the Indogangetic plain, Ganges–Brahmaputra delta, and the Bengal and the Indus Fans (⫽4.4 ⫻ 1019 kg) (Johnson, 1994). This amount includes sediment eroded before 17 Ma and thus represents an upper limit of the amount eroded during the last 17 Ma. Assuming that the sediment released 100% of its Os before or after deposition and that it contained, on average, 263 pM/kg of Os (Esser and Turekian, 1993), we find that the total Os that could have been supplied to the oceans by the Himalayan rivers (in dissolved form plus from the sediments) in the last 17 Ma is 40% of the total integrated flow during this time. If the 187 Os/188Os of the dissolved and suspended load were 1.44, it would contribute to about 41% of the increase in the seawater 187 Os/188Os 17 Ma ago. It follows that there is insufficient Os from the Himalayan drainage to account for the seawater evolution curve. Leaching experiments by Pegram et al. (1994) on the Ganges sediments suggest that they contain a component of labile Os that may be somewhat more radiogenic (187Os/188Os ⬃ 1.94) than the average riverine 187Os/188Os of 1.44. The leachable Os fraction constitutes ⬃10% of the Os in bulk sediment (Esser and Turekian, 1993). This suggests that the release of labile Os with 187Os/188Os ⬃ 1.94 would contribute to about 9% of the increase in the seawater 187Os/188Os 17 Ma ago. Reisberg et al. (1997) argue that the leachable radiogenic Os carried by the Himalayan river sediment is completely released to the oceans before deposition. As shown above, the bound from the sediments deposited by the Himalayan rivers suggests that a secondary release of Os while significant is not sufficient to drive the seawater evolution of Os. We, therefore, conclude that the total amount of Os with the corresponding isotopic composition provided from the weathering and erosion of the Himalayas cannot explain the marine Os isotopic evolution with time.

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Is it, then, possible that large inputs of highly radiogenic Os from rivers other than those from the Himalayas have significantly influenced the marine Os isotopic record over the past 17 Ma? Only a handful of large rivers have been analyzed so far and none of them appears to be extremely radiogenic (Levasseur, 1999; Sharma and Wasserburg, 1997). Rivers draining ancient cratons or areas dominated by black shales could possibly possess highly radiogenic Os. Indeed, rivers flowing through the Baltic Shield appear to have very radiogenic Os isotopes (Peucker–Ehrenbrink and Ravizza, 1996). These rivers are, however, too small to significantly influence the marine Os record (see below). Large rivers draining ancient organic matter (black shale), which is highly enriched in Os with average Re/Os about 15 times that of the average upper continental crust, should possess very high Os isotopic ratios. As Re is highly mobile under oxidizing conditions, such rivers should also be enriched in Re. For large rivers such as the Amazon and the Orinoco, we note that in comparison to the Himalayan rivers they are depleted in Re by a factor of 2 to 3 (Colodner et al., 1993). Hence, it is likely that these rivers do not provide Os much more radiogenic than the average continental crust (see also Levasseur, 1999). These observations are consistent with our previous conclusion that, on average, the rivers appear to supply osmium 187Os/188Os close to the average upper continental crust (Sharma and Wasserburg, 1997). However, we emphasize that the above inference is based on a small data set on large rivers. For example, many small rivers passing through ancient black shales could potentially supply large amounts of very radiogenic osmium. However, such rivers have not been represented in any sampling so far. 4.2. Time Evolution of Marine Osmium What are, then, the key factors that have controlled the temporal evolution of seawater Os? If the time scale for changes of the input is much greater than 104 yr, then on a 106-yr time scale, any phase lag is negligible and the quasisteady state Os isotopic ratio in seawater (SW) at time ␶ is given C C ␶ ␶ by: (187Os/188Os)SW ⫽ XOs (␶)(187Os/188Os)C ⫹ (1 ⫺ XOs (␶)) C C C H (187Os/188Os)H, where XOs (␶) ⫽ J188 ( ␶ )/( J ( ␶ ) ⫹ J ( 188Os 188Os ␶)) ⫽ Os H C i 188 1/(1 ⫹ J188 Os Os ( ␶ )/J188 Os ( ␶ )) and J188 Os represents flow of from reservoir i ⫽ Continent (C), Hydrothermal plus Cosmic (H); all the terms except (187Os/188Os)H are functions of time. The two pertinent variables governing the marine Os isotopic H C composition are, therefore, ROs (␶) ⬅ J188 Os/J188Os and 187 188 ␶ ( Os/ Os) C . Considering first the implications of the isotopic shifts inferred for the time interval between the present and 17 Ma ago, let us assume that the observed shift is due only to changes in fraction of Os coming from the continents. Using the nominal reference values: (187Os/188Os)C ⫽ 1.26, (187Os/188Os)H ⫽ C 0.126, and (187Os/188Os)SW ⫽ 1.05, this would require XOs to change from 0.52 (17 Ma ago) to 0.81 today. In other words, ROs should decrease by a factor of 4, from 0.91 to 0.23. On the other hand, if we were to assume the isotopic shift were due only to changes in the isotopic composition of the crustal component, this would require (187Os/188Os)C to change from 0.85 to 1.26 (today). If this shift in the crustal component were due to a shift of 187Os/188Os in a suite of rivers contributing a fraction “f ” of the net riverine Os input, then this would require

that there be a shift of ⌬(187Os/188Os)/f in the isotopic composition of these rivers. If f were 6.2 ⫻ 10⫺2 (⫽ the proportion of the Himalayan rivers), this would require a shift of ⌬(187Os/ 188 Os) ⫽ 6.6 over the past 17 Ma in this suite of rivers. It is evident that shifts in the riverine Os isotopic composition would have to be dramatic even if major, but not dominant, rivers were considered the cause of the observed variations. Clearly, for the estimated flow of dissolved Os, increasingly high 187Os/188Os from the Himalayan rivers cannot drive the seawater Os isotopic composition. Further, an unreasonably large shift of ⌬(187Os/188Os) ⫽ 1.0 is still inferred if f were 0.40 (⫽ dissolved Os delivered by the Himalayan rivers plus the upper bound on the Os released from the sediments). We conclude that the input of increasingly high 187Os/188Os from sources that do not comprise a large or dominant input cannot be the basis of the increase of 187Os/188Os in the oceans. It follows that large global changes in ROs (t) and/or (187Os/ 188 ␶ Os)C are required to produce the observed effects. The relative roles of these two scenarios are difficult to assess. If estimated fluxes of Ir and 3He in the oceanic sediments can be used to assess the contribution of cosmic dust to the oceans, then this contribution has increased several-fold during the last 17 million years, with most of the increase occurring in the last 3 million years (Farley, 1995; Kyte et al., 1993). This would cause ROs to increase rather than decrease. If seafloor spreading rate is used to estimate the hydrothermal flow of osmium, we note that it has also increased (by ⬃10%) during the last 17 Ma (Engebretson et al., 1992). Thus, the estimated increases in the cosmic accretion rate and the seafloor spreading rate would increase ROs and are therefore in the opposite direction to that required from the observations. We note that the mean continental elevation has increased by about 15% during the last 17 Ma (Harrison, 1994). A proportional relationship between mean continental elevation and flow of riverine Os to the oceans would thus be insufficient to cause the observed change in ROs. This requires that if global weathering and erosion were the cause of the observed shift in the marine Os isotopic ratio, the relationship between the riverine Os flow and continental elevation would have to be strongly non-linear. On the other hand, if a global increase in ␶ (187Os/188Os)C were the cause of the observed isotopic shift, it would require that none of the rivers sampled so far has yielded a representative 187Os/188Os flowing into the oceans, or that there is a non-riverine source of radiogenic Os component. 4.3. Osmium–Strontium Correlation We explore the above cases by comparing the marine Os record with the well-established shifts in marine 87Sr/86Sr and by posing the question: Why is there a crude positive correlation between the Os and Sr records for the last 40 Ma (Fig. 1)? Previous studies (Peucker–Ehrenbrink et al., 1995; Turekian and Pegram, 1997) have suggested that the correlation between 187 Os/188Os and 87Sr/86Sr ratios over the last 17 Ma is due to the regularly increasing weathering supply of more radiogenic Os and Sr from the Himalayas. Richter et al. (1992) have argued that the observed increase in seawater 87Sr/86Sr since 40 Ma (Fig. 1b) is due to the increased Sr flow from the Himalayan rivers. Edmond (1992) has pointed out that only the Himalayan type orogenies, where ancient rocks are remobilized, can pro-

Osmium Isotopes in the Himalayan Rivers

duce rivers enriched in Sr and containing extremely radiogenic Sr to drive the seawater Sr isotopic composition. However, the data from Krishnaswami et al. (1992) and Trivedi et al. (1995) show that the net output of the Brahmaputra, Indus, and Ganges, while enriched in Sr by a factor of 1.5 over world rivers, is not exceptionally radiogenic with 87Sr/86Sr being only 0.7181 as compared to the average input of 0.7119. Modeling by Krishnaswami et al. (1992) indicates that the Himalayan drainage can only account for 15–30% of the observed shift in the marine Sr isotopic record (see also Godde´ris and Franc¸ois, 1995). From the above discussion it is evident that the observed Os–Sr correlation has not resulted from the Himalayan riverine input. In the following, we pursue a two-component model to compare the observed Os–Sr correlation. As the residence time for Sr in the oceans is ⬃3 Ma (Palmer and Edmond, 1989), the present treatment is approximate since we neglect the phase lag. This should be a reasonable approximation for time scales 6 to 9 Ma. Assuming a quasi-steady state for Sr and an isotopic composition of marine carbonates close to seawater, the temporal isotopic evolution of marine 87Sr/86Sr is given by an ␶ equation symmetrical to the one for Os: (87Sr/86Sr)SW ⬇ C C 87 86 ␶ 87 86 XSr (␶)( Sr/ Sr)C ⫹ (1 ⫺ XSr(␶))( Sr/ Sr)H. Here, all symbols have the same connotation as above except H which represents only the hydrothermal inflow and no cosmic dust inflow for Sr. The equation for R is given by: R␶i ⫽ (1/XC i ) ⫺1 ⫽ SW Sw H j (⌫C i ⫺ ⌫i )/(⌫i (␶) ⫺ ⌫i ), where ⌫ i represents the isotopic composition of element i ⫽ Os, Sr in reservoir j ⫽ C, H, SW. We now explore the relationship between ROs and RSr. Assuming that there is a complete transfer of Os and Sr from riverine and hydrothermal sources to the oceans, we get: ROs/ H C H C H ˙ H C ˙ C H ˙ H C ˙ C RSr ⫽ (JOs /JOs )/(JSr /JSr) ⫽ (COs W /COs W )/(CSr W /CSr W ), j where Ci is the average concentration of i (⫽ Os, Sr) in j (C ⫽ ˙ j represents the water rivers, H ⫽ hydrothermal flow), and W flow rate from j. Both Os and Sr are apparently buffered in high temperature hydrothermal systems (Albare`de et al., 1981; Elderfield and Schultz, 1996; Piepgras and Wasserburg, 1985; Sharma et al., 1999). If the flows JH i are completely due to high H sw H temperature hydrothermal sources, then CSr ⬇ CSr and COs ⬇ SW SW C SW C COs , and ROs/RSr ⫽ COs /COs )/(CSr CSr), where the superscript SW C SW ⫽ seawater. Using COs ⫽ 57 fmol/kg, COs ⬇ 100 fmol/kg, SW C CSr ⫽ 90 ␮mol/kg, and CSr ⫽ 0.89 ␮mol/kg, we find that ROs/RSr ⬃ 6 ⫻ 10⫺3. This requires that the predominant source of the ‘H’ component for Os in the oceans cannot be from high temperature hydrothermal fluids but must be from cosmic dust infall. The above calculation suggests that the Os isotopic budget in the oceans reflects a balance between continental and cosmic dust infall (not high temperature hydrothermal sources). It follows that while the low 87Sr/86Sr at 40 to 50 Ma (Fig. 1b) can be from high hydrothermal input, the correlated 187Os/ 188 Os value during this time cannot be from this source, but must be due to low continental input. The partial coupling to the Os and Sr systems is provided by continental inputs. Thus, the observed correlation between Os and Sr isotopes must be due to increasing continental inputs which will effectively reduce both ROs and RSr. This suggests that a global increase in weathering may have produced the observed trends in the marine records of both Os and Sr. The above calculations and conclusions are strictly true only if the flows JH i are completely from high temperature ridge-

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crest fluids. As the low temperature hydrothermal flow rate possibly exceeds the high temperature fluid flow rate by a factor of ⬃100 (e.g., Elderfield and Schultz, 1996), the former may contribute significantly to JH i . This will be testable when the Os and Sr data for the low temperature ridge-flank fluids become available. The big problem with this model is that it requires a factor of four decrease in ROs over the last 17 million years while RSr remains essentially constant (Fig. 1c). Furthermore, it is not obvious how a 15% increase in elevation, a 10% increase in seafloor spreading rate and a factor of ⬃5 increase in cosmic dust flux over this time get translated into a factor of four decrease in ROs. 5. CONCLUSIONS

The net Himalayan river flow is neither enriched in Os nor extremely radiogenic compared to the other rivers measured so far. Moreover, the rivers have isotopic compositions similar to the estimates of the upper continental crust, indicating that they average out large segments of the continent containing diverse lithologies of varied ages. It is apparent that the Himalayan uplift itself has not directly provided Os to the oceans in amounts significant enough to cause the variations observed in the marine Os record over the last 17 Ma. Our modeling suggests that both the Sr and Os signals in seawater reflect continental input, but for Sr the non-radiogenic Sr is from hydrothermal sources while for Os, the dominant non-radiogenic source is from cosmic infall, unless contributions from low temperature hydrothermal fluids play a major role. If dominant non-radiogenic Os is from cosmic infall, this would effectively de-couple the behavior of these two systems with regard to hydrothermal sources but provide a correlated trend with continental erosion contributions without requiring highly radiogenic Os and Sr inputs under certain circumstances. The overall trend for osmium and strontium could then be explained by a regularly increasing input of global continental weathering sources. The uplift of the Tibetan Plateau may certainly have an impact on atmospheric and oceanic circulation and the global weathering cycle (Rind et al., 1997), however, our results indicate that the seawater Os isotopic composition is not directly coupled to the erosion of the Himalayas themselves. Acknowledgments—Funding for this work was provided by a Max– Planck Postdoctoral Fellowship to M.S. We are grateful to G. Contin for collecting the sample of the Indus river in Pakistan. We thank K. Turekian and J. Edmond for giving critical comments. Caltech G.P.S. Contribution 8607 (1000). Work supported by DOE grant DE-FG0388ER13851. REFERENCES Albare`de F., Michard A., Minster J. F., and Michard G. (1981) Sr-87– Sr-86 Ratios in Hydrothermal Waters and Deposits From the East Pacific Rise At 21-Degrees-N. Earth Planet. Sci. Lett. 55, 229 –236. Birck J. L., Roy Barman M., and Capmas F. (1997) Re–Os isotopic measurements at the femtomole level in natural samples. Geostandards Newslett. 20, 19 –27. Chesley J. T., Ruiz J., and Quade J. (1998) The 187Os/188Os record of Himalayan palaeorivers: Himalayan tectonics and ocean chemistry. Mineralogical Magazine, 62A, 323–324. Colodner D., Sachs J., Ravizza G., Turekian K., Edmond J., and Boyle E. (1993) The Geochemical Cycle of Rhenium—a Reconnaissance. Earth Planet. Sci. Lett. 117, 205–221. Edmond J. M. (1992) Himalayan tectonics, weathering processes, and

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