- Email: [email protected]
Recycling of oceanic crust and sediments: the noble gas subduction barrier
Earth and Planetary Science Letters, 89 (1988) 173-183 Elsevier Science Publishers B.V., A m s t e r d a m - Printed in The Netherlands
173
[61
Recycling of oceanic crust and sediments" the noble gas subduction barrier T h o m a s S t a u d a c h e r a n d C l a u d e J. All+gre Laboratoire de G$ochimie et Cosmochimie, Institut de Physique du Globe et DOpartement des Sciences de la Terre, Universit~s Paris 6 et 7, 4, Place Jussieu, F-75252 Paris Cedex 05 (France) Received October 26, 1987; revised version received March 3, 1988 We have determined the concentrations and isotopic composition of noble gases in old oceanic crust and oceanic sediments and the isotopic composition of noble gases in emanations from subduction volcanoes. Comparison with the noble gas signature of the upper mantle and a simple model allow us to conclude that at least 98% of the noble gases and water in the subducted slab returns back into the atmosphere through subduction volcanism before they can be admixed into the earth's mantle. It seems that the upper mantle is inaccessible to atmospheric noble gases due to an efficient subduction barrier for volatiles.
1. Introduction
2. Sample location and preparation
The earth's mantle presently ejects large quantities of gas through volcanic emanations either directly into the open air or through submarine volcanism and by fumarolic or phreato-magmatic activity. A priori such gases can be primordial in origin or have been recycled from gas-rich material that has been reinjected into the interior of the Earth through subduction processes. In fact, annually, about 7 x 1016 g of oceanic crust and up to 1.3 x 1016 g of sediments are available for reinjection into the Earth's mantle through subduction processes (section 5). Both oceanic crust and sediments are highly enriched in water and atmospheric gases; accordingly, several authors [1-3] have claimed that the noble gas signature of the Earth's mantle should be essentially controlled by subducted atmospheric noble gases. However, noble gas data on fresh MORB glasses show high 4°mr//36Ar ratios up to 28,000 [4,5,7] and 129Xe/13°Xe ratios up to 7.55 [5-7]. The latter correspond to 129Xe e x c e s s e s of up to 17% over atmosphere. These enrichments, products of the 4°K and 1291 decay respectively have been maintained in the source region for MORB, that is, the upper mantle, for at least 4 X 1 0 9 years [6-8]. In this paper, we will explore the problem of rare gases recycling through quantitative estimates of rare gas content and isotopic ratios in subducted material.
Old oceanic crust. We analyzed two fine-grained basalt samples from Ninety East Ridge in the Indian Ocean (Leg 22 214-48-1 and 216-37-3). Exact locations and further information are found in [9]. Six samples come from DSDP Leg 52 corresponding to Jurassic oceanic crust in the Atlantic Ocean. Drill holes 418A and 417D are close together at about 2 5 ° 5 ' N and 68°3'W. Water depth at these sites is 5500 m. The age of the oceanic crust was determined to be 108 Ma [10]. Detailed petrographic descriptions of the samples appear in [11]. For this study, we used pillow basalt or massive basalt from different positions in the drill core (Table 1) to determine whether alteration by seawater, as reflected in noble gas concentration or isotopic composition, is depth dependent.
0012-821X/88/$03.50
© 1988 Elsevier Science Publishers B.V.
Seafloor sediments. We analyzed six samples of
seafloor sediments. Two samples (MD-34-D2) are young and brittle sediments from the South West Indian Ridge (for location see [12]). Four samples (Leg 66 487-13-4 and 487-2-2) are unconsolidated sediments from a water depth of 4770 m from the Mexican Trench [13] (see Table 1). Volcanic gases. Finally, we analyzed six volcanic gas samples from subduction-related volcanoes,
174 TABLE 1 Drill core location and description of Leg 52, 22 and 66 samples Sample
Depth below seafloor (m)
Description
Leg 52 418A-18-3
sediment massif basalt
0-208 344
417D-48-5
sediment massif basalt
0-343 536
-52-6
pillow basalt
562
-55-2
pillow basalt
589
-65-6
massif basalt
672
-68-6
massif basalt
700
Leg 22 216-37-3 214-48-1
basalt glassy basalt
465.5 441
Leg 66 487-2-2
gray mud
3.5
-13-4
gray mud
110.5
MD-34-D2
sediment
dredged
sparsely phyric, medium gray massif basalt of intersertal to intergranular structure sparsely phyric, massif basalt with a mottled, dark gray groundmass having a medium- to coarse-grained, subophitic texture sparsely to moderately phyric pillow basalt; plagioclase, clinopyroxene and olivine phenocrysts highly phyric, slightly altered pillow basalt; groundmass glassy to fine-grained highly phyric massif basalt with a moderately altered, coarse-grained groundmass moderately phyric massif basalt with a doleritic coarse-grained groundmass fine-grained medium gray basalt, pilotaxitic trachytic texture medium dark gray fine-grained glassy basalt gray hemipelagic mud; clay material are dominantly quartz feldspars and mica; calcareous and siliceous microfossils are present transition between gray hemipelagic and brown pelagic clay microfossils are exclusively siliceous
M o m o t o m b o (Nicaragua), Poas ( C o s t a Rica) a n d V u l c a n o (Italy). Preparation. A l l r o c k samples were w a s h e d in p u r e ethanol. Brittle s e d i m e n t samples were w r a p p e d in a l u m i n u m foil in o r d e r to avoid m a t e r i a l loss a n d possible c o n t a m i n a t i o n of o t h e r s a m p l e s within the extraction line. O c e a n i c crust samples are large chips of 5 - 1 0 m m size, o b v i a t i n g the need for a l u m i n u m foil containers; typical b l a n k s are given in Tables 2 a n d 3. A l l oceanic crust s a m p l e s a n d three s e d i m e n t s a m p l e s were b a k e d at = 150 ° C for 48 hours p r i o r to analysis. A l i q u o t s o f the sediments were k e p t in an u n b a k e d p a r t of the e x t r a c t i o n line in o r d e r to test if a d s o r b e d a t m o spheric n o b l e gases are lost d u r i n g baking. Then, stepwise h e a t i n g o r total fusion was p e r f o r m e d in our c o n v e n t i o n a l e x t r a c t i o n line [4,6]. Volcanic gases are originally stored in c o p p e r tubes a n d filled in small p y r e x glass a m p u l e s
( < 0 . 1 c m 3) o n l y several d a y s b e f o r e analysis. T h e y were p l a c e d in a stainless steel crusher [14] b a k e d at = 1 0 0 ° C overnight a n d crushed u n d e r u l t r a h i g h v a c u u m b y a single stroke with a stainless steel b a l l bearing, levitated b y a m a g n e t outside the crusher wall. T h e storage time u n d e r v a c u u m was less t h a n 24 hours. N o b l e gases were then p u r i f i e d o n h o t Ti-sponges, s e p a r a t e d in four n o b l e gas fractions ( H e a n d Ne, Ar, Kr, Xe) a n d m e a s u r e d in A R E S I B O I [4-6]. 3. Results and observations 3.1. Oceanic crust N o b l e gas c o n c e n t r a t i o n s a n d i s o t o p i c c o m p o s i t i o n s are listed in T a b l e 2. N o n e of the b a s a l t s a m p l e s show 4 H e c o n c e n t r a t i o n s that c o u l d be e x p e c t e d d u e to their geological age. Based on the K c o n t e n t ( T a b l e 5) a n d a s s u m i n g a K / U of
175 TABLE 2 Noble gas data for old oceanic crust
a
4He ( 10-6 )
4He 3He
20Ne ( 10-1° )
36Ar (10 -1° )
4OAr 36Ar
84Kr (10 -14 )
130Xe (10 -14 )
Leg 22 Ninety East Ridge 216-37-3
800°C 1500°C total
1.88_+0.01 0.14_+0.01 2.02+0.02
--- 2.0×108 =2.5×106 -- 3.3×107
0.6_+0.2 0.6-+0.2
4.0 _+0.1 11.9 _+0.1 15.9 - + 0 . 1
1,582_+17 744_+ 8 953_+14
1,581 _+ 82 7,593 _+ 241 9,174 _+ 323
31 160 191
_+ 1 _+ 3 _+ 3
214-48-1
800°C 1500°C total
3.14+0.01 0.06_+0.01 3.20_+0.02
= 5.0× = 3.0x10
= 4.0x107
1.1-+0.4 0.7_+0.3 1.8_+0.5
3.6 - + 0 . 1 23.2 _+0.2 26.9 _+0.2
790-+10 1,100_+13 1,060+16
2,675 _+ 173 11,350 + 660 14,025 + 683
76 247 323
_+ 1 _+10 _+10
10 7 6
Leg 52 Jurassic 418A-18-3
total
0.25+0.01
<10 7
1.4-+0.7
64.0 _+2.8
347_+ 4
27,800 _+1,670
417D-48-5
800°C 1500°C total
0.35+0.01 0.01+0.01 0.36+0.02
<10 7
1.6-+0.2
<10 7
-
<10 7
1.6-+0.2
22.5 +1.0 6.9 -+0.3 29.4 _+1.1
330-+ 4 388-+ 5 344-+ 6
7,310 _+ 445 5,150 + 315 12,460 _+ 545
212 250 462
_+ 7 +_ 9 _+11
800°C 1600°C total
0.29+0.01 0.025:0.01 0.31-+0.02
<10 7
0.9-+0.4 0.9-+0.4
6.7 -+0.3 6.0 +0.2 12.7 +0.4
372_+ 5 363-+ 4 368-+ 6
1,747 _+ 105 2,230 -+ 134 3,977 _+ 170
82 58 140
5:3 _+ 2 _+ 4
-55-2
800°C 1050°C 1500°C total
0.62+0.01 =5.0×10 7 0.31_+0.02 --1.7×107 0.055:0.02 =3.0×106 0.97_+0.03 =2.0x107
0.4-+0.2 0.4_+0.2
17.0 8.1 5.7 30.8
335_+ 346_+ 388_+ 359_+
456 454 328 1,238
-+17 _+16 _+12 5:27
-65-6
800°C 1500°C total
0.42+0.01 --2.5×106 0.06_+0.01 =l.4X106 0.48_+0.02 - - - - 2 . 3 x 1 0 6
0.6-+0.1 0.6_+0.1
15.7 _+0.6 17.9 _+0.7 33.6 _+0.9
317_+ 4 352_+ 4 336_+ 6
5,445 + 330 13,460 -+ 810 18,905 -+ 870
329 655 984
-+12 -+23 _+26
-68-6
total
0.41_+0.01 =1.7×106
-
41.6 _+1.8
359+ 4
22,120 _+1,330
987
_+34
298_+ 5
8.0_+
-52-6
Blank a
1500 ° C
0.105:0.01
---3.0×107 <107
-
0.3-+0.1
-+0.7 _+0.3 _+0.2 _+0.8
0.07-+0.03
4 4 5 8
9,210 7,440 5,715 22,365
___ _+ _+ _+
550 450 345 790
3.6
1 , 5 7 7 _+54
0.24_+ 0.12
Concentrations are given in cm3 STP/g; blanks are in cm3 STP. 4He/3He ratios are only estimates (see text).
12,700 [15] we w o u l d e x p e c t b e t w e e n 0.6 a n d 6.4 x 10 - 6 c m 3 S T P / g 4 H e . T h i s i m p l i e s a loss o f 25-80% of radiogenic 4He. The 4He/3He ratios a r e o n l y e s t i m a t e s d u e to t h e l o w 3 H e c o n c e n t r a tion and relatively high HD + correction on mass 3. O n l y s a m p l e s f r o m t h e N i n e t y E a s t R i d g e s h o w H e c o n c e n t r a t i o n s h i g h e r t h a n 1 x 10 -6 c m 3 S T P / g ; t h e s e s a m p l e s also e x h i b i t t h e h i g h e s t 4 H e / 3 H e r a t i o s o f - 4 x 10 v w h i c h is p r o b a b l y d u e to h i g h e r U c o n c e n t r a t i o n s in t h e s e s a m p l e s , b a s e d o n t h e h i g h m e a s u r e d K c o n t e n t s ( T a b l e 5). Neon concentrations are low and typical 2°Ne/36Ar ratios are lower than the atmospheric o r d e e p - s e a w a t e r v a l u e (Fig. 1). 36Ar c o n c e n t r a t i o n s are r a t h e r c o n s t a n t a n d v a r y o n l y b e t w e e n 13 a n d 65 × 1 0 - l o c m 3 S T P / g w i t h a m e a n v a l u e o f
32 × 10 - 1 ° c m 3 S T P / g . T h e N i n e t y E a s t R i d g e b a s a l t s s h o w t h e h i g h e s t 4 ° A r / 3 6 A r r a t i o s u p to 1580 d u e to t h e i r e x c e p t i o n a l l y h i g h K c o n c e n t r a t i o n s ( T a b l e s 2 a n d 5). B a s a l t s a m p l e s f r o m L e g 52 h a v e a m e a n 4 ° A r / 3 6 A r r a t i o o f 350 _+ 12. U s i n g only temperature fractions with the highest 4°Ar/36Ar ratios produces a slightly higher mean o f 370 _+ 20. N o a p p a r e n t r e l a t i o n s h i p exists b e t w e e n e i t h e r 36Ar c o n c e n t r a t i o n o r 4 ° A r / 3 6 A r r a t i o w i t h d e p t h in t h e drill c o r e ( T a b l e 2); w e a s s u m e this to b e t r u e for t h e e n t i r e p a r t o f t h e o c e a n i c crust that has been altered by seawater and cont a m i n a t e d b y n o b l e gases d i s s o l v e d in s e a w a t e r , n o t m e r e l y t h e u p p e r 360 m t h a t w e a n a l y z e d . A l s o a p p e a r i n g in T a b l e 5 a r e K - A r ages o f the oceanic crust samples. Only one sample from Leg
176
I'X / 3eAr/,~, ('X/36Arl , .
55-2) may be due to exchange with deep-sea water-derived argon during alteration. Therefore, we can assume that the K-Ar ages of 65 and 36 Ma for Ninety East Ridge basalts may also be lower limits, even if these ages are in reasonable agreement with the ages of the oceanic crust as determined by ocean floor magnetic anomalies [16]. As for Ar, the Kr and Xe concentrations show no large variations. The isotopic compositions of both cannot be distinguished from atmospheric gases, with the exception of a small excess of 5 × 10 14 c m 3 S T P / g of ] 3 6 X e in the 1500°C fraction of sample 214-48-1 which is certainly due to spontaneous fission of 238U. Such noble gas concentrations confirm data by D y m o n d and Hogan [17] on two holocrystalline basalts. The only exceptions are their very small 4He concentrations and a 4°Ar/36Ar isotopic ratio which is identical to the atmospheric value, certainly due to their young geological ages.
I
1000
i,/I
100
10
mantkD dsw ak
oceanic st
0.1
Iments
0.01 20
36
He Ne Ar
84
Kr
130
Xe
Fig. 1. Noble gas abundance pattern relative to atmosphere and normalized to 36Ar for oceanic crust and oceanic sediments. The noble gas pattern for mean mantle [8] and deep sea water (dsw) are given for reference.
3.2. Seafloor sediments Table 3 gives noble gas data for sediment samples. Due to the small sample size, the isotopic composition of He and Ne was not measurable and Ne concentrations are only estimates. The 36Ar concentrations vary between 2 and 4 × 10 -8
52 (417D-68-6; K-Ar age = 105 Ma) yields an age which is identical to that of the oceanic crust at that site. The remaining basalts have lost large amounts of radiogenic 4°Ar (up to 77% for 417DTABLE 3 Noble gas data for oceanic sediments a 4He (10 -6)
20Ne (10 -1°)
36Ar (10 -1°)
500°C 1500°C total
0.44±0.20 0.97±0.41 1.41±0.46
17 17
165 53 218
± 8 ±13 ±15
350± 8 672± 33 428 ±23
448 400 848
± 25 ± 22 ± 33
7.4±0.3 7.8±0.3 15.2±0.4
unbaked
1500 o C
-
216
± 14
426±12
882
± 49
17.0±0.7
487-2-2 baked unbaked
1500°C 1500°C
1.65±0.47 1.25±0.43
18 ±10 0.8 ± 0.8
320 360
±19 +20
356± 8 396± 8
942 1560
± 52 ± 90
36.4±0.1 66.8±0.3
Indian Ocean MD-34-D2 baked 1500°C unbaked 1500 o C
0.37±0.17 0.20 ± 0.10
1.4 + 1.0 -
188 315
+10 ± 17
320± 6 317_+ 6
2026 3040
±110 ±350
45.4±1.7 78.1±2.8
Aluminum blank 500°C 1500°C
0.1 ±0.01 0.1 ±0.01
0.04± 0.01 0.30± 0.10
Leg 66 487-13-4 baked
a
± 7 ± 7
-
aOAr 36Ar
0.003± 0.001 0.25 + 0.03
Concentrations are given in cm3 STP/g; blanks are in cm 3 STP.
299+_10 295± 5
84Kr (10 12)
130Xe (10 12)
0.02± 0.08±
0.01 0.02
(4.0±2.0)X10 (2.3±0.8)×10
4 3
177
cm 3 STP/g. The measured 36Ar concentrations for baked and unbaked samples from Leg 66 agree within error. Although the results from the more brittle sediment MD-34-D2 indicate that baking resulted in the loss of some Ar, this difference is not reflected in isotopic composition as both baked and unbaked fractions yield the same low 4°Ar/ 36Ar ratio of - 320. Leg 66 sediments show slightly higher 4°Ar/36Ar ratios of 350-430 for whole rock samples and ratios up to 670 for the 1500°C fraction of Leg 66 487-13-4. While Kr and Xe systematically show somewhat higher concentrations in unbaked samples, the isotopic ratios are all atmospheric. The noble gas concentrations in Table 3 are up to 30 times smaller than concentrations reported by Podosek et al. [18] in a seafloor sediment; however, their 4°Ar/36Ar ratio of 390 is in agreement with our data. Further sediment data were recently published by Matsuda and Nagao [19]. Their noble gas concentrations in East Pacific sediments are higher by a factor of 5-10. Their 4°Ar/36Ar ratios of - 300 are significantly lower. 3.3. Volcanic gases
Noble gas data for He, Ne and Ar from volcanic gases appear in Table 4. Absolute amounts are not significant because the exact volumes of the glass ampules are unknown. Furthermore, Kr and Xe were not measured because these gases were trapped on a cooled glass surface during filling of the ampules. The He and Ar isotopic ratios are noteworthy. Momotombo and Poas are typical subduction-related volcanoes. The 4 H e / 3 He ratios of 113,000-163,000 show clearly that significant amounts of mantle He are involved during the
generation of arc magma. These results are in agreement with previous studies on gases from arc volcanoes [20] and in natural gases from subduction zones [21]. Recently, Poreda [22] published even much larger helium variations in back arc basalt glasses from Mariana Trough and Lau Basin. The 4 H e / 3 H e ratios range between about atmospheric value and 66,000, the later ratio being certainly due to a low-radiogenic hotspot component. Poreda also showed that the high water content and the D / H ratio ( S D = - 3 2 . 2 to -46%) are quite different from normal MORB values (SD = -77%o) and that part of these volatiles should be due to the water-rich descending slab. Further basalt glass samples from Mariana Trough analyzed by Sano et al. [23] show a more uniform 4 H e / 3 H e ratios of 97,000-106,500, thus significantly more radiogenic than normal MORB. The 4°Ar/36Ar ratios ranging between 360 and 3869, however, are much lower than for typical MORB glasses [4,5,7,8] and show that an important atmospheric noble gas component must be present in these glasses, possibly due to subducted oceanic crust a n d / o r sediments. The 4°Ar/36Ar ratios in our gas samples range from the atmospheric value up to 330 for ENS-177. For gases with low ratios, atmospheric contamination during gas-collection is obvious. Higher ratios of 330 are close to typical values for oceanic crust and sediments (Tables 2 and 3) but may still be contaminated to a certain degree by atmosphere. Similar ratios of 100,000 and 330 for 4 H e / 3 H e and 4°Ar/36Ar, respectively, have been found in a large plagioclase phenocryst for the arc volcano Zao in Japan (B. Marti, personal communication)
TABLE 4 lie, Ne and Ar data for volcanic gases a 4He
4He
20Ne
36Ar
40Ar
(10 6)
3He
(10 - 8)
(10 - 8)
36Ar
ENS-90 Vulcano ENS-93 Vulcano ENS-162 Momotombo ENS-177 Momotombo ENS-179 Momotombo ENS-186 Poas
0.64_+ 0.01 2.60 _+0.01 0.13 + 0.01 0.21 _+0.01 0.34 _+0.01
144,000 _+5,000 125,000 _+3,700 163,300 _+12,200
1.19 3.97 4.31 7.09 1.96 2.44
0.50 + 0.01 1.45 ___0.03 3.69+0.08 14.69_+0.32 1.38_+0.01 2.95 _+0.06
316 + 8 327 + 4 300+6 330+6 294_+6 293 + 6
Blank ampule
0.08 _+0.01
(1.8 _+0.6) x 10 4
285 _+7
112,800 _+6,500 126,700 _+6,240 -
a Measured gas contents and blanks are given in cm 3 STP.
+ 0.03 + 0.11 +0.11 +0.10 _+0.02 _+0.08
0.010 _+0.003
178 TABLE 5 K, 4°Ar * a n d K - A r ages for o c e a n i c c r u s t a n d o c e a n i c sediments Weight (g)
K (ppm)
4°Ar * (10 6 cm3/g)
K-Ar age (106 a)
1.3463 1.1449
4073 12542
1.043 1.765
65 36
1.4557 1.2335 1.0629 0.9875 1.1009 1.3441
850 336 2017 406 628
0.333 0.143 0.092 0.196 0.136 0.264
98 69 25 84 105
0.0196 0.0145 0.0181 0.0213 0.0575 0.0354
-
2.889 2.819 1.936 3.618 0.461 0.677
Oceanic crust Leg 22 Ninety East Ridge 216-37-3 214-48-1
Leg 52 Jurassic 418A-18-3 417D-48-5 -52-6 -55-2 -65-6 -68-6
Oceanic sediments Leg 66 487-13-4 -2-2 MD-34-D2
baked unbaked baked unbaked baked unbaked
-
and in natural bubbling water from the Nigorikawa basin in Japan [24]. It seems that during subduction large amounts of mantle He (up to 99.7% of 3He and 70% of 4He) are picked up in the partial melting zone of the subducted plate and outgassed through volcanic emanations. On the other hand, the smaller diffusion constant for Ar a n d / o r its relatively high concentration in the subducted plate may explain the similarity between the Ar isotopic compositions of volcanic gases and fumaroles and those for subducted oceanic crust and sediments. During devolatization a n d / o r melting of the slab, sufficient reinjected Ar is added to the source region of arc volcanoes to dominate their Ar isotopic characteristics. Nevertheless we cannot exclude that volcanic gases are additionally contaminated by gases from the back-arc crustal rocks through which ascending magma must traverse. 4. Summary (1) The 4°mr/36Ar ratio in old oceanic crust is about 350; however, it can locally increase up to 1600 or even higher.
(2) Radiogenic 4°Ar and 4He are not quantitatively retained in the oceanic basalts. In Jurassic rocks from Leg 52 up to 77% of 4 ° A r * and up to 80% of radiogenic 4He are lost possibly due to alteration effects. (3) No correlation exists between depth in the drillcore and the noble gas concentration or isotopic composition of samples from Leg 52. Oceanic crust may be contaminated by atmospheric noble gas dissolved in seawater down to several thousands of meters. (4) The isotopic ratios of 4°Ar/36Ar in oceanic sediments of 350-400 are not very different from those of oceanic crust. Younger sediments, certainly having higher water contents, show slightly lower ratios. (5) Noble gas concentrations in such sediments are about 10 times higher than those in oceanic crust. The relative abundance patterns in both, however, are identical (Fig. 1). Compared to the upper mantle, Ne shows a strong relative depletion and Xe a slight enrichment. (6) The highest 4 ° A r / / 3 6 A r ratios in gases from subduction volcanoes of 330 are nearly the same as those in the subducted oceanic plate indicating the contribution of volatiles from the subducted material to the mantle source of the arc magmatism. However, He shows an overwhelming mantle component with up to more than 99% of 3He being mantle-derived. 5. The subduction barrier for volatiles
In the next section, we will consider the reinjection of oceanic crust and oceanic sediments based on the data obtained. Injection of 4°K and 238U will increase the isotopic ratios of 4°Ar/36Ar and 136Xe/13°Xe, respectively. Reinjection of atmospheric noble gases would lower these isotopic ratios in mantle material. The upper mantle is in fact a complex system and a large number of processes should be taken into account: (a) mantle outgassing, for example [4,8]; (b) continental crust formation, for example [25,26]; (c) chemical fractionation between radiogenic and radioactive isotopes through partial melting; (d) reinjection of radioactive isotopes; and (e) reinjection of atmospheric gases. As a first approximation, we will not consider all of the above processes simultaneously but,
179
rather, we will discuss the reinjection of gases on a short time scale, thus neglecting, for example, the effect of reinjection of radioactive parent isotopes through subduction or their depletion in the upper mantle resulting from the formation of the continents. Even if presently we are not able to decide whether reinjection or depletion of radioactive elements into or from the mantle are dominant, we can show that compared to reinjection of atmospheric noble gases into the upper mantle, such phenomenon is of secondary importance. For example, let us suppose that the subduction rate of oceanic crust was constant within the last 1 AE and as well noble gases and uranium (0.08 p p m [15]) were subducted. Then in the same time span the fissiogenic 136Xe due to subducted 238U would be less than 0.13%o of the reinjected atmospheric 136Xe. Note, that for 129Xe the parent isotope 1291 is extinct and the problem of reinjection or depletion of the latter does not exist. Let us define ct = (1"/1) and/~ = ( P / I ) to be the isotopic ratio of a noble gas and the corresponding chemical ratio. I * is the abundance of the radiogenic isotope, P that of the parent isotope and I that of the stable, non-radiogenic isotope. We write the variation of the isotopic ratio in the upper mantle due to the radioactive decay and reinjection as: d
~ 0 ~ m = S~k~m + (o/r -- O~m)L(t )
(1)
where subscript "m" stands for upper mantle and " r " for reinjected material. S takes into account the branching for the radioactive decay or spontaneous fission * 4 ° K ~ 4°Ar: S -
X-----L--~ 0.1048 XB + Xe
238U-~4136Xe: S = _ _ h s f
. 136y = 3.70 × 10 - s
h~ + Xsf
L(t) is the transfer function for reinjected noble gases:
* X~=4.962×10 -1° a 1; X =1.54×10 10 a - l ; Xe = 0.581 :<10 - l ° a - l ; ~-sf= 8-47× 10-17 a - l ; 136y= 0.0673
For our approximation, we will use a (]~m) * * which is independent of continental crust formation over the time period that we consider. In fact, continental growth within the last 109 years is small [26]. Secondly, we will use L(t)= X. L to be constant, which means that as much oceanic crust is formed as oceanic crust is reinjected through subduction zones. X is a subduction barrier factor and stands for the fact that only a part of the noble gases in oceanic crust or sediments will be finally mixed into the upper mantle, while the rest ( 1 - X) will be lost before and during compaction of the sediments a n d / o r during subduction-related m a g m a t i c / m e t a s o m a t i c activity. We will define: L = m r-
[Ilo,:+s (i)m
(2)
where m r is the amount of subducted material per unit time, [I]oc+ s is the weighted concentration of the reference noble gas isotope in oceanic crust and sediments and ( I ) m Js the total amount of the corresponding isotope in the upper mantle. Integration of equation (1) gives us: (o~m - 0tr) t = (0~m -- Olr)0 • e - x L ' SX [e_X/ XL,] +/~0 XE--_ )~ --e-
(3)
where subscript t indicates the present-day ratio, and 0 the ratio t years ago. For a first estimate let us use: d
(O~m) = 0
(41
Such an assumption implies that otm is constant or that the variation in a due to subducted noble gases is equivalent and inverse to the variation due to radioactive decay. Of course such an assumption is crude for the following reasons: (1) When the main part of the present atmosphere was formed from the mantle, the isotopic ratios of 4°Ar/36Ar a n d 136Xe/13°Xe were ~< 296 and 2.17, respectively. Today, however, ratios of 4°Ar/36Ar as high as 28,000 [4,5,7,8] have been measured in mantle-derived rocks. Obviously, the radioactive part of equation (1) must be more important than the subduction part or such
** /.tm =~K/36Ar =10,500; ]~m=238U/130xe =1-4×106 [8].
180 TABLE 6 N o b l e g a s s i g n a t u r e a n d c o n c e n t r a t i o n s in the u p p e r m a n t l e [8], o c e a n i c s e d i m e n t s , a n d o c e a n i c crust. A t m o s p h e r e is g i v e n for reference a Reservoir
4 He
4He
Upper mantle
1.6×10 -5
82,000
Oceanic sediments Oceanic crust
1.0 × 1 0 - 6 1.0)<10 -6
---107
Atmosphere
0.153
20 N e
36Ar
40Ar
3He
722,500
s4 K r
130Xe
129Xe 130Xe
130Xe
2.7×10-13
6.6-7.6
2.2-2.55
1.6 × 1 0 - 9 1.6×10 -1°
4.3 × 1 0 - 1 1 7.4×10 -12
6.48 6.48
2.17 2.17
0.0207
0.000112
6.48
2.17
36Ar 7.9×10
10
7.0 × 1 0 - 1 o =1.0×10 -1° 0.525
3.5×10
10
2.7 × 10 - s 3.2)<10 - 9 =- 1
10,000 28,000 375 350 296
3.2×10
11
136 Xe
a C o n c e n t r a t i o n s a r e given in c m 3 S T P / g .
elevated ratios in the upper mantle could not be obtained today. (2) Today about 8 x 1 0 1 6 g of material are reinjected into the upper mantle every year. In the past, however, when mantle temperatures were higher and mantle convection rates correspondingly greater, the rates of formation of oceanic crust and subduction of material into the upper mantle were certainly quantitatively larger too. (3) On the other hand, formation of 4°Ar* from 4°K was more important in the past due to the higher abundance of the parent isotope (T1/2 = 1.25 x 1 0 9 a ) . However, the rate of formation of 1 3 6 X e by the spontaneous fission of 23SU was not proportionally as great due to the longer half-life of 238U ( = 4.5 X 10 9 a). Therefore we consider the subduction barrier factor X (equation (6)) which results from the estimate in the equation (4) as an upper limit. In fact X should be smaller than the value we will calculate below. Let us now try to estimate L. Presently, about 7.0 × 1016 g of oceanic crust are formed every year *, the same amount should be subducted. Unfortunately, we do not know exactly to what depth oceanic crust is charged with atmospheric gases. We assume here 40-80% is in fact contaminated by atmospheric noble gases. Together with oceanic crust, oceanic sediments are subducted too. As sediments contain about ten times the noble gas content of oceanic basaltic crust, they are particularly important in determining the
m=60,1300km length of ridges
x
5km
× 2×4cm/aX2.9g/cm
depth of oceanic crust
spreading rate
density
3
mass of reinjected noble gases. Oceanic sediments represent a mass of about 1.0 × 10 24 g [27]. The amount of oceanic crust rnoc is about 4.5 × 1 0 24 g * *. This means that on average 18% by weight of oceanic crustal material (crust and sediments) is sediments. We will not take into account here that the distribution of oceanic sediments is not uniform but depends on the age of the oceanic crust. Young ridge areas which are usually further away from subduction zones have had less time to accumulate significant amounts of sediments and vice versa. An important complication arises, because the amount of subducted sediments is not known. Estimates, based on geochemical studies of island arc basalts range from no sediment reinjection [28] up to subduction of hundreds of meter thick sediment layer [29]. As we cannot distinguish, based on noble gases, between the two theories, we will use both to bracked the amount of sediment reinjection. In case one we only allow reinjection of oceanic crust but no reinjection of sediments, in case two we will reinject oceanic crust and all of the oceanic sediments, as estimated above. We then get the following balance of subducted rocks: o c e a n i c c r u s t m: altered oceanic crust ma: s e d i m e n t s ms:
7.0 X 1016 g / a (4.2_+1.4)×1016 g/a b e t w e e n 0 a n d 1.3 x 1016 g / a
Using the mean noble gas concentrations in Table 6, we estimate the total amounts of noble gases that should be subducted according to the equa-
** mo~=3.1×10
s km 2 ×
surface of the oceanic crust
5kin
× 2.9g/cm 3
depth of oceanic crust
density
181 tion: 'I = m r [ i l ] o c + s = ma'Ioc + ms'I s
(5)
or 4He: (4.2-5.5) X 101° cm 3 S T P / a 2°Ne: (4.2-13.3) x 10 6 c m 3 S T P / a 36mr: (1.3-4.9) X 108 cm 3 S T P / a 84Kr: (6.7-27.5) x 10 6 c m 3 S T P / a 13°Xe: (3.1-8.7) x 105 cm 3 S T P / a All~gre et al. [8] estimated noble gas quantities ( I ) m in the upper mantle of 3.8 X 1017cm3 STP of 36mr and 3.8 x 1014 cm 3 STP of 13°Xe. The value L in equation (2) will then be (3.42-12.9)x 10 - 1 ° and (8.16-22.9)X 10 -1° for 36Ar and 13°Xe, respectively. 5.1. Volatile reinjection without subduction barrier Let us briefly describe what would be the effects if all such gases in the subducted matter were to be admixed with the upper mantle. Assuming a constant subduction rate during 109 years about 9.7 × 1015 cm 3 STP of 2°Ne, 3.4 × 1017cm3 STP of 36Ar, 1.9 × 1016 c m 3 STP of 84Kr, 6.4 × 1014 c m 3 STP of 13°Xe and 1.7 × 10 24 c m 3 o f H 2 0 would be subducted. Such quantities, however, correspond to 1%, 90%, 110% and 170% of the total 2°Ne, 3 6 A r , 8 4 K r and 13°Xe respectively, that are presently stored in the upper mantle [8] and about 1.5 times the water in the world's oceans. The corresponding gas circulation would never allow the maintenance of high argon or xenon isotopic ratios within the mantle. Accordingly, these values (especially that for H 2 0 ) are completely unrealistic. 5.2. The subduction barrier factor Let us estimate now the quantity of gases that are liberated from the slab before mixing with the upper mantle. From equation (4) we get: S~k/.t m =
-- (o/r - -
Otm)LX
or:
X-
SA~rn
(~m--~r)L
(6)
Assuming that there was no significant change in the noble gas isotopic ratios of the atmosphere and using a r = 370 and 2.17 for 4°Ar/36Ar and 136Xe/13°Xe, respectively, we get a subduction barrier factor X of 0.04 + 0.02 for argon and 0.017 + 0.008 for xenon. This means that the sub-
ducted slab loses at least 98% of its gases before the material is admixed to the upper mantle. Of course, such an estimate is quite crude due to our assumption that d a / d t = 0. Isotopic ratios of Ar and Xe in the upper mantle are definitely high [4,5,8,30] so the subduction component in equation (1) must be much smaller than the radiogenic one. Consequently, the subduction barrier factor X must be even much smaller than our estimated value of - 0.017. We should comment here that our oceanic sediments are in fact dry samples and do not necessarily represent water-laden deep-sea sediments. Published noble gas data for seafloor sediments show Xe concentrations up to 30 times higher than our highest value for oceanic sediments [18,19]. Furthermore, shales may contain even more noble gases than oceanic sediments, especially in their wet state [31-33]. However, the only consequence of higher noble gas concentrations in the subducted matter would be an even smaller subduction barrier factor of 0.002 or less. The gas loss from the subduction layer occurs in two stages. First, sediments undergo compaction which results in the release of water and, certainly, noble gases. Vents, resembling the black and white smokers of the East Pacific Rise have been discovered along the subduction zone off Oregon [34] and indications for interstitial water venting from the down going oceanic plate off Japan [35] are probably the results of this process. Secondly, partial melting takes place in the diving slab, creating subduction volcanism through which large amounts of gases are released back into the atmosphere. It is worth to point out here a large difference in the geodynamic behaviour of different volatiles. While noble gases and water, once outgassed from the mantle stay nearly quantitatively in the atmosphere-hydrosphere, Javoy et al. [36] estimated that carbon, outgassed from the mantle essentially as CO 2 has only a residence time in the atmosphere-hydrosphere of 200 M a before it is reinjected into the mantle in the form of carbonates. 5.3. Xenon isotopic ratios and the subduction barrier factor We will finally demonstrate how the 129Xe// 13°Xe and 136Xe/a3°Xe ratios depend on the subduction barrier factor X.
182 10
la°Xe
8 ~.L~
0
7
~
,
2
-
,
1 Time (Ga)
(a)
4.0
\
13'Xe
lS°Xe 3 5 "
,
'
~00.5
0
,\ "~
30
0 2.5
2
1
0
Time (Gal Fig. 2. Evolution of (a) 129Xe/13°Xe and (b) 136Xe/13°Xe calculated back in time, using the subduction barrier factor X as a parameter and imposing present-day values of 7.55 and 2.55, respectively. (b)
F r o m equation (3) we find:
O/me = (O:m __ Otr) t eXL, +0/%
X LSX _ XP,o[e
where a m is the isotopic ratio t years ago. For 129Xe/13°Xe we can simplify (equation (7)) as S = 0. In Fig. 2a and b we have calculated the theoretical evolution of the isotopic ratios 129Xe/13°Xe and 136Xe/13°Xe within the earth's mantle using X as a parameter. For each value of X, the intersection of the curve with the Y-axis gives us the respective ratio that would be required in the upper mantle 2 × 10 9 years ago to, ultimately, produce the greater-than-atmospheric ratios that are observed today in mantle-derived rocks [5,6]. The ]29Xe/13°Xe ratio (Fig. 2a) is always decreasing with time, due to the fact that the parent isotope 129I with its small half-live (7"1/2 = 18 Ma) has completely decayed by the time of our calculation and reinjection of atmospheric xenon will always dilute ]29Xe excesses in the M O R B source. Fig. 2b portrays the evolution of 136Xe/13°Xe. For small values of the subduction barrier factor
X, the ratio is increasing, due to the production of ]36Xe by spontaneous fission of 23Su. F o r large values of X, however, the dilution by reinjected atmospheric xenon is more important than the production of fissiogenic 136Xe. Fig. 2a and b demonstrate that for large values of X it would be i m p o s s i b l e (a) to f o r m large excesses of 136Xe/13°Xe by spontaneous fission of 238U, and (b) to m a i n t a i n large 129xe/a3°xe and 136Xe/13°Xe ratios in the mantle, as measured in M O R B glasses, over a time span of the order of 10 9 years. Let us use an example. If X were 0.3, then we would need 129Xe/13°Xe and 136Xe/13°Xe ratios in the mantle 2 × 10 9 years ago of > 9.5 and > 3.3 respectively in order to be able to measure ratios of 7.5 and 2.6 in M O R B glasses today. Even if an extrapolation over 4.5 × 10 9 years is difficult because of the numerous abovementioned processes operating during mantle evolution, such estimates show us that we would need unreasonably high xenon excesses in the early earth, in order to reconcile them with the actual data. Consequently we propose that subduction zone volcanism forms a very efficient " s u b d u c t i o n barrier for volatiles" and keeps the earth's mantle almost inaccessible to atmospheric noble gases. Acknowledgements
D. Rousseau kindly measured the K contents of the oceanic crust samples. M. Javoy donated the volcanic gas ampules. We thank B.J. Pegram and S.H. Richardson for improving the English and C. Mercier and M. Sennegon for typing the manuscript. We thank the a n o n y m o u s reviewers for their critical remarks. This is I P G Contribution No. 1022. References
1 M. Ozima, F.A. Podosek and G. Igarashi, Terrestrial xenon isotope constraints on the early history of the Earth, Nature 315, 471-474, 1985. 2 O.K. Manuel and D.D. Sabu, The noble gas record of the terrestrial planets, Geochem. J. 15, 245-267, 1981. 3 T. Kirsten, H. Richter and D. Storzer, Abundance patterns of rare gases in submarine basalts and glasses (abstract), Meteoritics 16(4), 341, 1981. 4 Ph. Sarda, Th. Staudacher and C.J. All+gre, 4°Ar/36Ar in MORB glasses: constraints on atmosphere and mantle evolution, Earth Planet. Sci. Lett. 72, 357-375, 1985.
183 5 Th. Staudacher, Ph. Sarda, S. Richardson, C.J. All6gre, I. Sagna and L. Dimitriev, Noble gas isotopes from a topographic high on the MAR at 14 ° N, in preparation. 6 Th. Staudacher and C.J. Allrgre, Terrestrial xenology, Earth Planet. Sci. Lett. 60, 389-406, 1982. 7 C.J. Allrgre, Th. Staudacher, Ph. Sarda and M. Kurz, Constraints on evolution of earth's mantle from rare gas systematics, Nature 303, 762-766, 1983. 8 C.J. Allrgre, Th. Staudacher and Ph. Sarda, Rare gas systematics: formation of the atmosphere, evolution and structure of the Earth's mantle, Earth Planet. Sci. Lett. 81, 127-150, 1986/87. 9 Ch.C. Borch, J.G. Sclater, S. Gartner, Jr., R. Hekinian, D.A. Johnson, B. McGowran, A.C. Pimm, R.W. Thompson, J.J. Veevers and L.S. Waterman, Initial Reports of the Deep Sea Drilling Project, Vol. XXII, U.S. Government Printing Office, Washington, D.C., 1972. 10 T. Donnelly, J. Francheteau, W. Bryan, P. Robinson, M. Flower and M. Salisbury, Initial Reports of the Deep Sea Drilling Project, Vol. LI, LII, LIII, U.S. Government Printing Office, Washington, D.C., 1980. 11 C. Bollinger, Etude grochimique des phrnocristaux de plagioclases des roches ocraniques, Thesis, Universit6 Paris 7, 1979. 12 B. Hamelin and C.J. Allrgre, Large scale regional units in the depleted upper mantle revealed by an isotope study of the South-West Indian Ridge, Nature 315, 196-199, 1985. 13 J. Watkins, J.C. Moore, S.B. Bachman, F. Beghtel et al., Initial Reports of the Deep Sea Drilling Project, Vol. LXVI, U.S. Government Printing Office, Washington, D.C., 1979. 14 Th. Staudacher, M.D. Kurz and C.J. Allrgre, New noble gas data on glass samples from Loihi seamount and Hualalai and on dunite samples from Loihi and Rrunion Island, Chem. Geol. 56, 193-205, 1986. 15 K.P. Jochum, A.W. Hofmann, E. Ito, H.M. Seufert and W.M. White, K, U and Th in mid-ocean ridge basalt glasses and heat production, K / U and K / R b in the mantle, Nature 306, 431-436, 1983. 16 J.G. Sclater, B. Parson and C. Jaupart, Oceans and continents: similarities and differences in the mechanisms of heat loss, J. Geophys. Res. 86, 11,535-11,552, 1981. 17 J. Dymond and L. Hogan, Noble gas abundance patterns in deep-sea basalts--primordial gases from the mantle, Earth Planet. Sci. Lett. 20, 131-139, 1973. 18 F.A. Podosek, M. Honda and M. Ozima, Sedimentary noble gases, Geochim. Cosmochim. Acta 44, 1875-1884, 1980. 19 J.-I. Matsuda and K. Nagao, Noble gas abundance in a deep-sea sediment core from eastern equatorial Pacific, Geochem. J. 20, 71-80, 1986. 20 H. Craig, J.E. Lupton and Y. Horibe, A mantle helium component in circum pacific volcanic gases: Hakone, the Marianas and Mount Larsen, in: Terrestrial Rare Gases,
21
22
23
24
25
26
27
28
29
30
31
32 33
34
35
36
E.C. Alexander and M. Ozima, eds., pp. 3-16, Cent. Acad. Publ. Japan, Tokyo, 1978. Y. Sano, Y. Nakamura and H. Wakita, Areal distribution of 3He/4He ratios in the Tohoku district, Northeastern Japan, Chem. Geol. 52, 1-8, 1985. R. Poreda, Helium-3 and deuterium in back-arc basalts: Lau Basin and the Mariana Trough, Earth Planet. Sci. Lett. 73, 244-254, 1985. Y. Sano, Y. Nakamura, H. Wakita and T. Ishii, Light noble gases in basalt glasses from Mariana Trough, Geochim. Cosmochim. Acta 50, 2429-2432, 1986. K. Nagao, N.S. Takaoka and O. Matsubayashi, Isotopic anomalies of rare gases in the Nigorikawa geothermal area, Hokkaido, Japan, Earth Planet. Sci. Lett. 44, 82-90, 1979. C.J. Allrgre, S.R. Hart and J.F. Minster, Chemical structure and evolution of the mantle and continents determined by inversion of Nd and Sr isotopic data, II. Numerical experiments and discussion, Earth Planet. Sci. Lett. 66, 191-213, 1983. C.J. Allrgre and C. Jaupart, Continental tectonics and continental kinetics, Earth Planet. Sci. Lett. 74, 171-186, 1985. R.M. Garrels and F.T. MacKenzie, Sedimentary rocks, in: Evolution of Sedimentary Rocks, W.W. Norton&Company, New York, N.Y., 1971. R.J. Arculus, Island arc magmatism in relation to the evolution of the crust and mantle, Tectonophysics 75, 113-133, 1981. R.W. Kay, Volcanic arc magmas: implications of a melting-mixing model for element recycling in the crust-upper mantle system, J. Geol. 88, 497-522, 1980. C.J. Allrgre and Th. Staudacher, Terrestrial xenology, II, Japan-U.S. Seminar on Terrestrial Rare Gases, Yellowstone, Abstracts, pp. 1-4, 1986. R.A. Canalas, E.C. Alexander and O.K. Manuel, Terrestrial abundances of noble gases, J. Geophys. Res. 73, 3331-3334, 1968. D. Phinney, 36Ar, Kr and Xe in terrestrial materials, Earth Planet. Sci. Lett. 16, 413-420, 1972. T.J. Bernatowicz, F.A. Podosek, M. Honda and F.E. Kramer, The atmospheric inventory of xenon and noble gases in shales: the plastic bag experiment, J. Geophys. Res. 89, 4597-4611, 1984. E. Suess, B. Carson, S.D. Ritger, J.C. Moore, L.D. Kulm and G.R. Cochrane, Biological communities at vent sites along the subduction zone off Oregon, Bull. Biol. Soc. Washington 6, 475-484, 1985. J. Boulrgue, J.T. Iiyama, J.L. Charlou and J. Jedwab, Nankal Trough, Japan Trench and Kuril Trench: geochemistry of fluid samples by submersible "Nautile", Earth Planet. Sci. Lett. 83, 363-375, 1987. M. Javoy, F. Pineau and C.J. Allrgre, Carbon geodynamic cycle, Nature 300, 171-173, 1982.
173
[61
Recycling of oceanic crust and sediments" the noble gas subduction barrier T h o m a s S t a u d a c h e r a n d C l a u d e J. All+gre Laboratoire de G$ochimie et Cosmochimie, Institut de Physique du Globe et DOpartement des Sciences de la Terre, Universit~s Paris 6 et 7, 4, Place Jussieu, F-75252 Paris Cedex 05 (France) Received October 26, 1987; revised version received March 3, 1988 We have determined the concentrations and isotopic composition of noble gases in old oceanic crust and oceanic sediments and the isotopic composition of noble gases in emanations from subduction volcanoes. Comparison with the noble gas signature of the upper mantle and a simple model allow us to conclude that at least 98% of the noble gases and water in the subducted slab returns back into the atmosphere through subduction volcanism before they can be admixed into the earth's mantle. It seems that the upper mantle is inaccessible to atmospheric noble gases due to an efficient subduction barrier for volatiles.
1. Introduction
2. Sample location and preparation
The earth's mantle presently ejects large quantities of gas through volcanic emanations either directly into the open air or through submarine volcanism and by fumarolic or phreato-magmatic activity. A priori such gases can be primordial in origin or have been recycled from gas-rich material that has been reinjected into the interior of the Earth through subduction processes. In fact, annually, about 7 x 1016 g of oceanic crust and up to 1.3 x 1016 g of sediments are available for reinjection into the Earth's mantle through subduction processes (section 5). Both oceanic crust and sediments are highly enriched in water and atmospheric gases; accordingly, several authors [1-3] have claimed that the noble gas signature of the Earth's mantle should be essentially controlled by subducted atmospheric noble gases. However, noble gas data on fresh MORB glasses show high 4°mr//36Ar ratios up to 28,000 [4,5,7] and 129Xe/13°Xe ratios up to 7.55 [5-7]. The latter correspond to 129Xe e x c e s s e s of up to 17% over atmosphere. These enrichments, products of the 4°K and 1291 decay respectively have been maintained in the source region for MORB, that is, the upper mantle, for at least 4 X 1 0 9 years [6-8]. In this paper, we will explore the problem of rare gases recycling through quantitative estimates of rare gas content and isotopic ratios in subducted material.
Old oceanic crust. We analyzed two fine-grained basalt samples from Ninety East Ridge in the Indian Ocean (Leg 22 214-48-1 and 216-37-3). Exact locations and further information are found in [9]. Six samples come from DSDP Leg 52 corresponding to Jurassic oceanic crust in the Atlantic Ocean. Drill holes 418A and 417D are close together at about 2 5 ° 5 ' N and 68°3'W. Water depth at these sites is 5500 m. The age of the oceanic crust was determined to be 108 Ma [10]. Detailed petrographic descriptions of the samples appear in [11]. For this study, we used pillow basalt or massive basalt from different positions in the drill core (Table 1) to determine whether alteration by seawater, as reflected in noble gas concentration or isotopic composition, is depth dependent.
0012-821X/88/$03.50
© 1988 Elsevier Science Publishers B.V.
Seafloor sediments. We analyzed six samples of
seafloor sediments. Two samples (MD-34-D2) are young and brittle sediments from the South West Indian Ridge (for location see [12]). Four samples (Leg 66 487-13-4 and 487-2-2) are unconsolidated sediments from a water depth of 4770 m from the Mexican Trench [13] (see Table 1). Volcanic gases. Finally, we analyzed six volcanic gas samples from subduction-related volcanoes,
174 TABLE 1 Drill core location and description of Leg 52, 22 and 66 samples Sample
Depth below seafloor (m)
Description
Leg 52 418A-18-3
sediment massif basalt
0-208 344
417D-48-5
sediment massif basalt
0-343 536
-52-6
pillow basalt
562
-55-2
pillow basalt
589
-65-6
massif basalt
672
-68-6
massif basalt
700
Leg 22 216-37-3 214-48-1
basalt glassy basalt
465.5 441
Leg 66 487-2-2
gray mud
3.5
-13-4
gray mud
110.5
MD-34-D2
sediment
dredged
sparsely phyric, medium gray massif basalt of intersertal to intergranular structure sparsely phyric, massif basalt with a mottled, dark gray groundmass having a medium- to coarse-grained, subophitic texture sparsely to moderately phyric pillow basalt; plagioclase, clinopyroxene and olivine phenocrysts highly phyric, slightly altered pillow basalt; groundmass glassy to fine-grained highly phyric massif basalt with a moderately altered, coarse-grained groundmass moderately phyric massif basalt with a doleritic coarse-grained groundmass fine-grained medium gray basalt, pilotaxitic trachytic texture medium dark gray fine-grained glassy basalt gray hemipelagic mud; clay material are dominantly quartz feldspars and mica; calcareous and siliceous microfossils are present transition between gray hemipelagic and brown pelagic clay microfossils are exclusively siliceous
M o m o t o m b o (Nicaragua), Poas ( C o s t a Rica) a n d V u l c a n o (Italy). Preparation. A l l r o c k samples were w a s h e d in p u r e ethanol. Brittle s e d i m e n t samples were w r a p p e d in a l u m i n u m foil in o r d e r to avoid m a t e r i a l loss a n d possible c o n t a m i n a t i o n of o t h e r s a m p l e s within the extraction line. O c e a n i c crust samples are large chips of 5 - 1 0 m m size, o b v i a t i n g the need for a l u m i n u m foil containers; typical b l a n k s are given in Tables 2 a n d 3. A l l oceanic crust s a m p l e s a n d three s e d i m e n t s a m p l e s were b a k e d at = 150 ° C for 48 hours p r i o r to analysis. A l i q u o t s o f the sediments were k e p t in an u n b a k e d p a r t of the e x t r a c t i o n line in o r d e r to test if a d s o r b e d a t m o spheric n o b l e gases are lost d u r i n g baking. Then, stepwise h e a t i n g o r total fusion was p e r f o r m e d in our c o n v e n t i o n a l e x t r a c t i o n line [4,6]. Volcanic gases are originally stored in c o p p e r tubes a n d filled in small p y r e x glass a m p u l e s
( < 0 . 1 c m 3) o n l y several d a y s b e f o r e analysis. T h e y were p l a c e d in a stainless steel crusher [14] b a k e d at = 1 0 0 ° C overnight a n d crushed u n d e r u l t r a h i g h v a c u u m b y a single stroke with a stainless steel b a l l bearing, levitated b y a m a g n e t outside the crusher wall. T h e storage time u n d e r v a c u u m was less t h a n 24 hours. N o b l e gases were then p u r i f i e d o n h o t Ti-sponges, s e p a r a t e d in four n o b l e gas fractions ( H e a n d Ne, Ar, Kr, Xe) a n d m e a s u r e d in A R E S I B O I [4-6]. 3. Results and observations 3.1. Oceanic crust N o b l e gas c o n c e n t r a t i o n s a n d i s o t o p i c c o m p o s i t i o n s are listed in T a b l e 2. N o n e of the b a s a l t s a m p l e s show 4 H e c o n c e n t r a t i o n s that c o u l d be e x p e c t e d d u e to their geological age. Based on the K c o n t e n t ( T a b l e 5) a n d a s s u m i n g a K / U of
175 TABLE 2 Noble gas data for old oceanic crust
a
4He ( 10-6 )
4He 3He
20Ne ( 10-1° )
36Ar (10 -1° )
4OAr 36Ar
84Kr (10 -14 )
130Xe (10 -14 )
Leg 22 Ninety East Ridge 216-37-3
800°C 1500°C total
1.88_+0.01 0.14_+0.01 2.02+0.02
--- 2.0×108 =2.5×106 -- 3.3×107
0.6_+0.2 0.6-+0.2
4.0 _+0.1 11.9 _+0.1 15.9 - + 0 . 1
1,582_+17 744_+ 8 953_+14
1,581 _+ 82 7,593 _+ 241 9,174 _+ 323
31 160 191
_+ 1 _+ 3 _+ 3
214-48-1
800°C 1500°C total
3.14+0.01 0.06_+0.01 3.20_+0.02
= 5.0× = 3.0x10
= 4.0x107
1.1-+0.4 0.7_+0.3 1.8_+0.5
3.6 - + 0 . 1 23.2 _+0.2 26.9 _+0.2
790-+10 1,100_+13 1,060+16
2,675 _+ 173 11,350 + 660 14,025 + 683
76 247 323
_+ 1 _+10 _+10
10 7 6
Leg 52 Jurassic 418A-18-3
total
0.25+0.01
<10 7
1.4-+0.7
64.0 _+2.8
347_+ 4
27,800 _+1,670
417D-48-5
800°C 1500°C total
0.35+0.01 0.01+0.01 0.36+0.02
<10 7
1.6-+0.2
<10 7
-
<10 7
1.6-+0.2
22.5 +1.0 6.9 -+0.3 29.4 _+1.1
330-+ 4 388-+ 5 344-+ 6
7,310 _+ 445 5,150 + 315 12,460 _+ 545
212 250 462
_+ 7 +_ 9 _+11
800°C 1600°C total
0.29+0.01 0.025:0.01 0.31-+0.02
<10 7
0.9-+0.4 0.9-+0.4
6.7 -+0.3 6.0 +0.2 12.7 +0.4
372_+ 5 363-+ 4 368-+ 6
1,747 _+ 105 2,230 -+ 134 3,977 _+ 170
82 58 140
5:3 _+ 2 _+ 4
-55-2
800°C 1050°C 1500°C total
0.62+0.01 =5.0×10 7 0.31_+0.02 --1.7×107 0.055:0.02 =3.0×106 0.97_+0.03 =2.0x107
0.4-+0.2 0.4_+0.2
17.0 8.1 5.7 30.8
335_+ 346_+ 388_+ 359_+
456 454 328 1,238
-+17 _+16 _+12 5:27
-65-6
800°C 1500°C total
0.42+0.01 --2.5×106 0.06_+0.01 =l.4X106 0.48_+0.02 - - - - 2 . 3 x 1 0 6
0.6-+0.1 0.6_+0.1
15.7 _+0.6 17.9 _+0.7 33.6 _+0.9
317_+ 4 352_+ 4 336_+ 6
5,445 + 330 13,460 -+ 810 18,905 -+ 870
329 655 984
-+12 -+23 _+26
-68-6
total
0.41_+0.01 =1.7×106
-
41.6 _+1.8
359+ 4
22,120 _+1,330
987
_+34
298_+ 5
8.0_+
-52-6
Blank a
1500 ° C
0.105:0.01
---3.0×107 <107
-
0.3-+0.1
-+0.7 _+0.3 _+0.2 _+0.8
0.07-+0.03
4 4 5 8
9,210 7,440 5,715 22,365
___ _+ _+ _+
550 450 345 790
3.6
1 , 5 7 7 _+54
0.24_+ 0.12
Concentrations are given in cm3 STP/g; blanks are in cm3 STP. 4He/3He ratios are only estimates (see text).
12,700 [15] we w o u l d e x p e c t b e t w e e n 0.6 a n d 6.4 x 10 - 6 c m 3 S T P / g 4 H e . T h i s i m p l i e s a loss o f 25-80% of radiogenic 4He. The 4He/3He ratios a r e o n l y e s t i m a t e s d u e to t h e l o w 3 H e c o n c e n t r a tion and relatively high HD + correction on mass 3. O n l y s a m p l e s f r o m t h e N i n e t y E a s t R i d g e s h o w H e c o n c e n t r a t i o n s h i g h e r t h a n 1 x 10 -6 c m 3 S T P / g ; t h e s e s a m p l e s also e x h i b i t t h e h i g h e s t 4 H e / 3 H e r a t i o s o f - 4 x 10 v w h i c h is p r o b a b l y d u e to h i g h e r U c o n c e n t r a t i o n s in t h e s e s a m p l e s , b a s e d o n t h e h i g h m e a s u r e d K c o n t e n t s ( T a b l e 5). Neon concentrations are low and typical 2°Ne/36Ar ratios are lower than the atmospheric o r d e e p - s e a w a t e r v a l u e (Fig. 1). 36Ar c o n c e n t r a t i o n s are r a t h e r c o n s t a n t a n d v a r y o n l y b e t w e e n 13 a n d 65 × 1 0 - l o c m 3 S T P / g w i t h a m e a n v a l u e o f
32 × 10 - 1 ° c m 3 S T P / g . T h e N i n e t y E a s t R i d g e b a s a l t s s h o w t h e h i g h e s t 4 ° A r / 3 6 A r r a t i o s u p to 1580 d u e to t h e i r e x c e p t i o n a l l y h i g h K c o n c e n t r a t i o n s ( T a b l e s 2 a n d 5). B a s a l t s a m p l e s f r o m L e g 52 h a v e a m e a n 4 ° A r / 3 6 A r r a t i o o f 350 _+ 12. U s i n g only temperature fractions with the highest 4°Ar/36Ar ratios produces a slightly higher mean o f 370 _+ 20. N o a p p a r e n t r e l a t i o n s h i p exists b e t w e e n e i t h e r 36Ar c o n c e n t r a t i o n o r 4 ° A r / 3 6 A r r a t i o w i t h d e p t h in t h e drill c o r e ( T a b l e 2); w e a s s u m e this to b e t r u e for t h e e n t i r e p a r t o f t h e o c e a n i c crust that has been altered by seawater and cont a m i n a t e d b y n o b l e gases d i s s o l v e d in s e a w a t e r , n o t m e r e l y t h e u p p e r 360 m t h a t w e a n a l y z e d . A l s o a p p e a r i n g in T a b l e 5 a r e K - A r ages o f the oceanic crust samples. Only one sample from Leg
176
I'X / 3eAr/,~, ('X/36Arl , .
55-2) may be due to exchange with deep-sea water-derived argon during alteration. Therefore, we can assume that the K-Ar ages of 65 and 36 Ma for Ninety East Ridge basalts may also be lower limits, even if these ages are in reasonable agreement with the ages of the oceanic crust as determined by ocean floor magnetic anomalies [16]. As for Ar, the Kr and Xe concentrations show no large variations. The isotopic compositions of both cannot be distinguished from atmospheric gases, with the exception of a small excess of 5 × 10 14 c m 3 S T P / g of ] 3 6 X e in the 1500°C fraction of sample 214-48-1 which is certainly due to spontaneous fission of 238U. Such noble gas concentrations confirm data by D y m o n d and Hogan [17] on two holocrystalline basalts. The only exceptions are their very small 4He concentrations and a 4°Ar/36Ar isotopic ratio which is identical to the atmospheric value, certainly due to their young geological ages.
I
1000
i,/I
100
10
mantkD dsw ak
oceanic st
0.1
Iments
0.01 20
36
He Ne Ar
84
Kr
130
Xe
Fig. 1. Noble gas abundance pattern relative to atmosphere and normalized to 36Ar for oceanic crust and oceanic sediments. The noble gas pattern for mean mantle [8] and deep sea water (dsw) are given for reference.
3.2. Seafloor sediments Table 3 gives noble gas data for sediment samples. Due to the small sample size, the isotopic composition of He and Ne was not measurable and Ne concentrations are only estimates. The 36Ar concentrations vary between 2 and 4 × 10 -8
52 (417D-68-6; K-Ar age = 105 Ma) yields an age which is identical to that of the oceanic crust at that site. The remaining basalts have lost large amounts of radiogenic 4°Ar (up to 77% for 417DTABLE 3 Noble gas data for oceanic sediments a 4He (10 -6)
20Ne (10 -1°)
36Ar (10 -1°)
500°C 1500°C total
0.44±0.20 0.97±0.41 1.41±0.46
17 17
165 53 218
± 8 ±13 ±15
350± 8 672± 33 428 ±23
448 400 848
± 25 ± 22 ± 33
7.4±0.3 7.8±0.3 15.2±0.4
unbaked
1500 o C
-
216
± 14
426±12
882
± 49
17.0±0.7
487-2-2 baked unbaked
1500°C 1500°C
1.65±0.47 1.25±0.43
18 ±10 0.8 ± 0.8
320 360
±19 +20
356± 8 396± 8
942 1560
± 52 ± 90
36.4±0.1 66.8±0.3
Indian Ocean MD-34-D2 baked 1500°C unbaked 1500 o C
0.37±0.17 0.20 ± 0.10
1.4 + 1.0 -
188 315
+10 ± 17
320± 6 317_+ 6
2026 3040
±110 ±350
45.4±1.7 78.1±2.8
Aluminum blank 500°C 1500°C
0.1 ±0.01 0.1 ±0.01
0.04± 0.01 0.30± 0.10
Leg 66 487-13-4 baked
a
± 7 ± 7
-
aOAr 36Ar
0.003± 0.001 0.25 + 0.03
Concentrations are given in cm3 STP/g; blanks are in cm 3 STP.
299+_10 295± 5
84Kr (10 12)
130Xe (10 12)
0.02± 0.08±
0.01 0.02
(4.0±2.0)X10 (2.3±0.8)×10
4 3
177
cm 3 STP/g. The measured 36Ar concentrations for baked and unbaked samples from Leg 66 agree within error. Although the results from the more brittle sediment MD-34-D2 indicate that baking resulted in the loss of some Ar, this difference is not reflected in isotopic composition as both baked and unbaked fractions yield the same low 4°Ar/ 36Ar ratio of - 320. Leg 66 sediments show slightly higher 4°Ar/36Ar ratios of 350-430 for whole rock samples and ratios up to 670 for the 1500°C fraction of Leg 66 487-13-4. While Kr and Xe systematically show somewhat higher concentrations in unbaked samples, the isotopic ratios are all atmospheric. The noble gas concentrations in Table 3 are up to 30 times smaller than concentrations reported by Podosek et al. [18] in a seafloor sediment; however, their 4°Ar/36Ar ratio of 390 is in agreement with our data. Further sediment data were recently published by Matsuda and Nagao [19]. Their noble gas concentrations in East Pacific sediments are higher by a factor of 5-10. Their 4°Ar/36Ar ratios of - 300 are significantly lower. 3.3. Volcanic gases
Noble gas data for He, Ne and Ar from volcanic gases appear in Table 4. Absolute amounts are not significant because the exact volumes of the glass ampules are unknown. Furthermore, Kr and Xe were not measured because these gases were trapped on a cooled glass surface during filling of the ampules. The He and Ar isotopic ratios are noteworthy. Momotombo and Poas are typical subduction-related volcanoes. The 4 H e / 3 He ratios of 113,000-163,000 show clearly that significant amounts of mantle He are involved during the
generation of arc magma. These results are in agreement with previous studies on gases from arc volcanoes [20] and in natural gases from subduction zones [21]. Recently, Poreda [22] published even much larger helium variations in back arc basalt glasses from Mariana Trough and Lau Basin. The 4 H e / 3 H e ratios range between about atmospheric value and 66,000, the later ratio being certainly due to a low-radiogenic hotspot component. Poreda also showed that the high water content and the D / H ratio ( S D = - 3 2 . 2 to -46%) are quite different from normal MORB values (SD = -77%o) and that part of these volatiles should be due to the water-rich descending slab. Further basalt glass samples from Mariana Trough analyzed by Sano et al. [23] show a more uniform 4 H e / 3 H e ratios of 97,000-106,500, thus significantly more radiogenic than normal MORB. The 4°Ar/36Ar ratios ranging between 360 and 3869, however, are much lower than for typical MORB glasses [4,5,7,8] and show that an important atmospheric noble gas component must be present in these glasses, possibly due to subducted oceanic crust a n d / o r sediments. The 4°Ar/36Ar ratios in our gas samples range from the atmospheric value up to 330 for ENS-177. For gases with low ratios, atmospheric contamination during gas-collection is obvious. Higher ratios of 330 are close to typical values for oceanic crust and sediments (Tables 2 and 3) but may still be contaminated to a certain degree by atmosphere. Similar ratios of 100,000 and 330 for 4 H e / 3 H e and 4°Ar/36Ar, respectively, have been found in a large plagioclase phenocryst for the arc volcano Zao in Japan (B. Marti, personal communication)
TABLE 4 lie, Ne and Ar data for volcanic gases a 4He
4He
20Ne
36Ar
40Ar
(10 6)
3He
(10 - 8)
(10 - 8)
36Ar
ENS-90 Vulcano ENS-93 Vulcano ENS-162 Momotombo ENS-177 Momotombo ENS-179 Momotombo ENS-186 Poas
0.64_+ 0.01 2.60 _+0.01 0.13 + 0.01 0.21 _+0.01 0.34 _+0.01
144,000 _+5,000 125,000 _+3,700 163,300 _+12,200
1.19 3.97 4.31 7.09 1.96 2.44
0.50 + 0.01 1.45 ___0.03 3.69+0.08 14.69_+0.32 1.38_+0.01 2.95 _+0.06
316 + 8 327 + 4 300+6 330+6 294_+6 293 + 6
Blank ampule
0.08 _+0.01
(1.8 _+0.6) x 10 4
285 _+7
112,800 _+6,500 126,700 _+6,240 -
a Measured gas contents and blanks are given in cm 3 STP.
+ 0.03 + 0.11 +0.11 +0.10 _+0.02 _+0.08
0.010 _+0.003
178 TABLE 5 K, 4°Ar * a n d K - A r ages for o c e a n i c c r u s t a n d o c e a n i c sediments Weight (g)
K (ppm)
4°Ar * (10 6 cm3/g)
K-Ar age (106 a)
1.3463 1.1449
4073 12542
1.043 1.765
65 36
1.4557 1.2335 1.0629 0.9875 1.1009 1.3441
850 336 2017 406 628
0.333 0.143 0.092 0.196 0.136 0.264
98 69 25 84 105
0.0196 0.0145 0.0181 0.0213 0.0575 0.0354
-
2.889 2.819 1.936 3.618 0.461 0.677
Oceanic crust Leg 22 Ninety East Ridge 216-37-3 214-48-1
Leg 52 Jurassic 418A-18-3 417D-48-5 -52-6 -55-2 -65-6 -68-6
Oceanic sediments Leg 66 487-13-4 -2-2 MD-34-D2
baked unbaked baked unbaked baked unbaked
-
and in natural bubbling water from the Nigorikawa basin in Japan [24]. It seems that during subduction large amounts of mantle He (up to 99.7% of 3He and 70% of 4He) are picked up in the partial melting zone of the subducted plate and outgassed through volcanic emanations. On the other hand, the smaller diffusion constant for Ar a n d / o r its relatively high concentration in the subducted plate may explain the similarity between the Ar isotopic compositions of volcanic gases and fumaroles and those for subducted oceanic crust and sediments. During devolatization a n d / o r melting of the slab, sufficient reinjected Ar is added to the source region of arc volcanoes to dominate their Ar isotopic characteristics. Nevertheless we cannot exclude that volcanic gases are additionally contaminated by gases from the back-arc crustal rocks through which ascending magma must traverse. 4. Summary (1) The 4°mr/36Ar ratio in old oceanic crust is about 350; however, it can locally increase up to 1600 or even higher.
(2) Radiogenic 4°Ar and 4He are not quantitatively retained in the oceanic basalts. In Jurassic rocks from Leg 52 up to 77% of 4 ° A r * and up to 80% of radiogenic 4He are lost possibly due to alteration effects. (3) No correlation exists between depth in the drillcore and the noble gas concentration or isotopic composition of samples from Leg 52. Oceanic crust may be contaminated by atmospheric noble gas dissolved in seawater down to several thousands of meters. (4) The isotopic ratios of 4°Ar/36Ar in oceanic sediments of 350-400 are not very different from those of oceanic crust. Younger sediments, certainly having higher water contents, show slightly lower ratios. (5) Noble gas concentrations in such sediments are about 10 times higher than those in oceanic crust. The relative abundance patterns in both, however, are identical (Fig. 1). Compared to the upper mantle, Ne shows a strong relative depletion and Xe a slight enrichment. (6) The highest 4 ° A r / / 3 6 A r ratios in gases from subduction volcanoes of 330 are nearly the same as those in the subducted oceanic plate indicating the contribution of volatiles from the subducted material to the mantle source of the arc magmatism. However, He shows an overwhelming mantle component with up to more than 99% of 3He being mantle-derived. 5. The subduction barrier for volatiles
In the next section, we will consider the reinjection of oceanic crust and oceanic sediments based on the data obtained. Injection of 4°K and 238U will increase the isotopic ratios of 4°Ar/36Ar and 136Xe/13°Xe, respectively. Reinjection of atmospheric noble gases would lower these isotopic ratios in mantle material. The upper mantle is in fact a complex system and a large number of processes should be taken into account: (a) mantle outgassing, for example [4,8]; (b) continental crust formation, for example [25,26]; (c) chemical fractionation between radiogenic and radioactive isotopes through partial melting; (d) reinjection of radioactive isotopes; and (e) reinjection of atmospheric gases. As a first approximation, we will not consider all of the above processes simultaneously but,
179
rather, we will discuss the reinjection of gases on a short time scale, thus neglecting, for example, the effect of reinjection of radioactive parent isotopes through subduction or their depletion in the upper mantle resulting from the formation of the continents. Even if presently we are not able to decide whether reinjection or depletion of radioactive elements into or from the mantle are dominant, we can show that compared to reinjection of atmospheric noble gases into the upper mantle, such phenomenon is of secondary importance. For example, let us suppose that the subduction rate of oceanic crust was constant within the last 1 AE and as well noble gases and uranium (0.08 p p m [15]) were subducted. Then in the same time span the fissiogenic 136Xe due to subducted 238U would be less than 0.13%o of the reinjected atmospheric 136Xe. Note, that for 129Xe the parent isotope 1291 is extinct and the problem of reinjection or depletion of the latter does not exist. Let us define ct = (1"/1) and/~ = ( P / I ) to be the isotopic ratio of a noble gas and the corresponding chemical ratio. I * is the abundance of the radiogenic isotope, P that of the parent isotope and I that of the stable, non-radiogenic isotope. We write the variation of the isotopic ratio in the upper mantle due to the radioactive decay and reinjection as: d
~ 0 ~ m = S~k~m + (o/r -- O~m)L(t )
(1)
where subscript "m" stands for upper mantle and " r " for reinjected material. S takes into account the branching for the radioactive decay or spontaneous fission * 4 ° K ~ 4°Ar: S -
X-----L--~ 0.1048 XB + Xe
238U-~4136Xe: S = _ _ h s f
. 136y = 3.70 × 10 - s
h~ + Xsf
L(t) is the transfer function for reinjected noble gases:
* X~=4.962×10 -1° a 1; X =1.54×10 10 a - l ; Xe = 0.581 :<10 - l ° a - l ; ~-sf= 8-47× 10-17 a - l ; 136y= 0.0673
For our approximation, we will use a (]~m) * * which is independent of continental crust formation over the time period that we consider. In fact, continental growth within the last 109 years is small [26]. Secondly, we will use L(t)= X. L to be constant, which means that as much oceanic crust is formed as oceanic crust is reinjected through subduction zones. X is a subduction barrier factor and stands for the fact that only a part of the noble gases in oceanic crust or sediments will be finally mixed into the upper mantle, while the rest ( 1 - X) will be lost before and during compaction of the sediments a n d / o r during subduction-related m a g m a t i c / m e t a s o m a t i c activity. We will define: L = m r-
[Ilo,:+s (i)m
(2)
where m r is the amount of subducted material per unit time, [I]oc+ s is the weighted concentration of the reference noble gas isotope in oceanic crust and sediments and ( I ) m Js the total amount of the corresponding isotope in the upper mantle. Integration of equation (1) gives us: (o~m - 0tr) t = (0~m -- Olr)0 • e - x L ' SX [e_X/ XL,] +/~0 XE--_ )~ --e-
(3)
where subscript t indicates the present-day ratio, and 0 the ratio t years ago. For a first estimate let us use: d
(O~m) = 0
(41
Such an assumption implies that otm is constant or that the variation in a due to subducted noble gases is equivalent and inverse to the variation due to radioactive decay. Of course such an assumption is crude for the following reasons: (1) When the main part of the present atmosphere was formed from the mantle, the isotopic ratios of 4°Ar/36Ar a n d 136Xe/13°Xe were ~< 296 and 2.17, respectively. Today, however, ratios of 4°Ar/36Ar as high as 28,000 [4,5,7,8] have been measured in mantle-derived rocks. Obviously, the radioactive part of equation (1) must be more important than the subduction part or such
** /.tm =~K/36Ar =10,500; ]~m=238U/130xe =1-4×106 [8].
180 TABLE 6 N o b l e g a s s i g n a t u r e a n d c o n c e n t r a t i o n s in the u p p e r m a n t l e [8], o c e a n i c s e d i m e n t s , a n d o c e a n i c crust. A t m o s p h e r e is g i v e n for reference a Reservoir
4 He
4He
Upper mantle
1.6×10 -5
82,000
Oceanic sediments Oceanic crust
1.0 × 1 0 - 6 1.0)<10 -6
---107
Atmosphere
0.153
20 N e
36Ar
40Ar
3He
722,500
s4 K r
130Xe
129Xe 130Xe
130Xe
2.7×10-13
6.6-7.6
2.2-2.55
1.6 × 1 0 - 9 1.6×10 -1°
4.3 × 1 0 - 1 1 7.4×10 -12
6.48 6.48
2.17 2.17
0.0207
0.000112
6.48
2.17
36Ar 7.9×10
10
7.0 × 1 0 - 1 o =1.0×10 -1° 0.525
3.5×10
10
2.7 × 10 - s 3.2)<10 - 9 =- 1
10,000 28,000 375 350 296
3.2×10
11
136 Xe
a C o n c e n t r a t i o n s a r e given in c m 3 S T P / g .
elevated ratios in the upper mantle could not be obtained today. (2) Today about 8 x 1 0 1 6 g of material are reinjected into the upper mantle every year. In the past, however, when mantle temperatures were higher and mantle convection rates correspondingly greater, the rates of formation of oceanic crust and subduction of material into the upper mantle were certainly quantitatively larger too. (3) On the other hand, formation of 4°Ar* from 4°K was more important in the past due to the higher abundance of the parent isotope (T1/2 = 1.25 x 1 0 9 a ) . However, the rate of formation of 1 3 6 X e by the spontaneous fission of 23SU was not proportionally as great due to the longer half-life of 238U ( = 4.5 X 10 9 a). Therefore we consider the subduction barrier factor X (equation (6)) which results from the estimate in the equation (4) as an upper limit. In fact X should be smaller than the value we will calculate below. Let us now try to estimate L. Presently, about 7.0 × 1016 g of oceanic crust are formed every year *, the same amount should be subducted. Unfortunately, we do not know exactly to what depth oceanic crust is charged with atmospheric gases. We assume here 40-80% is in fact contaminated by atmospheric noble gases. Together with oceanic crust, oceanic sediments are subducted too. As sediments contain about ten times the noble gas content of oceanic basaltic crust, they are particularly important in determining the
m=60,1300km length of ridges
x
5km
× 2×4cm/aX2.9g/cm
depth of oceanic crust
spreading rate
density
3
mass of reinjected noble gases. Oceanic sediments represent a mass of about 1.0 × 10 24 g [27]. The amount of oceanic crust rnoc is about 4.5 × 1 0 24 g * *. This means that on average 18% by weight of oceanic crustal material (crust and sediments) is sediments. We will not take into account here that the distribution of oceanic sediments is not uniform but depends on the age of the oceanic crust. Young ridge areas which are usually further away from subduction zones have had less time to accumulate significant amounts of sediments and vice versa. An important complication arises, because the amount of subducted sediments is not known. Estimates, based on geochemical studies of island arc basalts range from no sediment reinjection [28] up to subduction of hundreds of meter thick sediment layer [29]. As we cannot distinguish, based on noble gases, between the two theories, we will use both to bracked the amount of sediment reinjection. In case one we only allow reinjection of oceanic crust but no reinjection of sediments, in case two we will reinject oceanic crust and all of the oceanic sediments, as estimated above. We then get the following balance of subducted rocks: o c e a n i c c r u s t m: altered oceanic crust ma: s e d i m e n t s ms:
7.0 X 1016 g / a (4.2_+1.4)×1016 g/a b e t w e e n 0 a n d 1.3 x 1016 g / a
Using the mean noble gas concentrations in Table 6, we estimate the total amounts of noble gases that should be subducted according to the equa-
** mo~=3.1×10
s km 2 ×
surface of the oceanic crust
5kin
× 2.9g/cm 3
depth of oceanic crust
density
181 tion: 'I = m r [ i l ] o c + s = ma'Ioc + ms'I s
(5)
or 4He: (4.2-5.5) X 101° cm 3 S T P / a 2°Ne: (4.2-13.3) x 10 6 c m 3 S T P / a 36mr: (1.3-4.9) X 108 cm 3 S T P / a 84Kr: (6.7-27.5) x 10 6 c m 3 S T P / a 13°Xe: (3.1-8.7) x 105 cm 3 S T P / a All~gre et al. [8] estimated noble gas quantities ( I ) m in the upper mantle of 3.8 X 1017cm3 STP of 36mr and 3.8 x 1014 cm 3 STP of 13°Xe. The value L in equation (2) will then be (3.42-12.9)x 10 - 1 ° and (8.16-22.9)X 10 -1° for 36Ar and 13°Xe, respectively. 5.1. Volatile reinjection without subduction barrier Let us briefly describe what would be the effects if all such gases in the subducted matter were to be admixed with the upper mantle. Assuming a constant subduction rate during 109 years about 9.7 × 1015 cm 3 STP of 2°Ne, 3.4 × 1017cm3 STP of 36Ar, 1.9 × 1016 c m 3 STP of 84Kr, 6.4 × 1014 c m 3 STP of 13°Xe and 1.7 × 10 24 c m 3 o f H 2 0 would be subducted. Such quantities, however, correspond to 1%, 90%, 110% and 170% of the total 2°Ne, 3 6 A r , 8 4 K r and 13°Xe respectively, that are presently stored in the upper mantle [8] and about 1.5 times the water in the world's oceans. The corresponding gas circulation would never allow the maintenance of high argon or xenon isotopic ratios within the mantle. Accordingly, these values (especially that for H 2 0 ) are completely unrealistic. 5.2. The subduction barrier factor Let us estimate now the quantity of gases that are liberated from the slab before mixing with the upper mantle. From equation (4) we get: S~k/.t m =
-- (o/r - -
Otm)LX
or:
X-
SA~rn
(~m--~r)L
(6)
Assuming that there was no significant change in the noble gas isotopic ratios of the atmosphere and using a r = 370 and 2.17 for 4°Ar/36Ar and 136Xe/13°Xe, respectively, we get a subduction barrier factor X of 0.04 + 0.02 for argon and 0.017 + 0.008 for xenon. This means that the sub-
ducted slab loses at least 98% of its gases before the material is admixed to the upper mantle. Of course, such an estimate is quite crude due to our assumption that d a / d t = 0. Isotopic ratios of Ar and Xe in the upper mantle are definitely high [4,5,8,30] so the subduction component in equation (1) must be much smaller than the radiogenic one. Consequently, the subduction barrier factor X must be even much smaller than our estimated value of - 0.017. We should comment here that our oceanic sediments are in fact dry samples and do not necessarily represent water-laden deep-sea sediments. Published noble gas data for seafloor sediments show Xe concentrations up to 30 times higher than our highest value for oceanic sediments [18,19]. Furthermore, shales may contain even more noble gases than oceanic sediments, especially in their wet state [31-33]. However, the only consequence of higher noble gas concentrations in the subducted matter would be an even smaller subduction barrier factor of 0.002 or less. The gas loss from the subduction layer occurs in two stages. First, sediments undergo compaction which results in the release of water and, certainly, noble gases. Vents, resembling the black and white smokers of the East Pacific Rise have been discovered along the subduction zone off Oregon [34] and indications for interstitial water venting from the down going oceanic plate off Japan [35] are probably the results of this process. Secondly, partial melting takes place in the diving slab, creating subduction volcanism through which large amounts of gases are released back into the atmosphere. It is worth to point out here a large difference in the geodynamic behaviour of different volatiles. While noble gases and water, once outgassed from the mantle stay nearly quantitatively in the atmosphere-hydrosphere, Javoy et al. [36] estimated that carbon, outgassed from the mantle essentially as CO 2 has only a residence time in the atmosphere-hydrosphere of 200 M a before it is reinjected into the mantle in the form of carbonates. 5.3. Xenon isotopic ratios and the subduction barrier factor We will finally demonstrate how the 129Xe// 13°Xe and 136Xe/a3°Xe ratios depend on the subduction barrier factor X.
182 10
la°Xe
8 ~.L~
0
7
~
,
2
-
,
1 Time (Ga)
(a)
4.0
\
13'Xe
lS°Xe 3 5 "
,
'
~00.5
0
,\ "~
30
0 2.5
2
1
0
Time (Gal Fig. 2. Evolution of (a) 129Xe/13°Xe and (b) 136Xe/13°Xe calculated back in time, using the subduction barrier factor X as a parameter and imposing present-day values of 7.55 and 2.55, respectively. (b)
F r o m equation (3) we find:
O/me = (O:m __ Otr) t eXL, +0/%
X LSX _ XP,o[e
where a m is the isotopic ratio t years ago. For 129Xe/13°Xe we can simplify (equation (7)) as S = 0. In Fig. 2a and b we have calculated the theoretical evolution of the isotopic ratios 129Xe/13°Xe and 136Xe/13°Xe within the earth's mantle using X as a parameter. For each value of X, the intersection of the curve with the Y-axis gives us the respective ratio that would be required in the upper mantle 2 × 10 9 years ago to, ultimately, produce the greater-than-atmospheric ratios that are observed today in mantle-derived rocks [5,6]. The ]29Xe/13°Xe ratio (Fig. 2a) is always decreasing with time, due to the fact that the parent isotope 129I with its small half-live (7"1/2 = 18 Ma) has completely decayed by the time of our calculation and reinjection of atmospheric xenon will always dilute ]29Xe excesses in the M O R B source. Fig. 2b portrays the evolution of 136Xe/13°Xe. For small values of the subduction barrier factor
X, the ratio is increasing, due to the production of ]36Xe by spontaneous fission of 23Su. F o r large values of X, however, the dilution by reinjected atmospheric xenon is more important than the production of fissiogenic 136Xe. Fig. 2a and b demonstrate that for large values of X it would be i m p o s s i b l e (a) to f o r m large excesses of 136Xe/13°Xe by spontaneous fission of 238U, and (b) to m a i n t a i n large 129xe/a3°xe and 136Xe/13°Xe ratios in the mantle, as measured in M O R B glasses, over a time span of the order of 10 9 years. Let us use an example. If X were 0.3, then we would need 129Xe/13°Xe and 136Xe/13°Xe ratios in the mantle 2 × 10 9 years ago of > 9.5 and > 3.3 respectively in order to be able to measure ratios of 7.5 and 2.6 in M O R B glasses today. Even if an extrapolation over 4.5 × 10 9 years is difficult because of the numerous abovementioned processes operating during mantle evolution, such estimates show us that we would need unreasonably high xenon excesses in the early earth, in order to reconcile them with the actual data. Consequently we propose that subduction zone volcanism forms a very efficient " s u b d u c t i o n barrier for volatiles" and keeps the earth's mantle almost inaccessible to atmospheric noble gases. Acknowledgements
D. Rousseau kindly measured the K contents of the oceanic crust samples. M. Javoy donated the volcanic gas ampules. We thank B.J. Pegram and S.H. Richardson for improving the English and C. Mercier and M. Sennegon for typing the manuscript. We thank the a n o n y m o u s reviewers for their critical remarks. This is I P G Contribution No. 1022. References
1 M. Ozima, F.A. Podosek and G. Igarashi, Terrestrial xenon isotope constraints on the early history of the Earth, Nature 315, 471-474, 1985. 2 O.K. Manuel and D.D. Sabu, The noble gas record of the terrestrial planets, Geochem. J. 15, 245-267, 1981. 3 T. Kirsten, H. Richter and D. Storzer, Abundance patterns of rare gases in submarine basalts and glasses (abstract), Meteoritics 16(4), 341, 1981. 4 Ph. Sarda, Th. Staudacher and C.J. All+gre, 4°Ar/36Ar in MORB glasses: constraints on atmosphere and mantle evolution, Earth Planet. Sci. Lett. 72, 357-375, 1985.
183 5 Th. Staudacher, Ph. Sarda, S. Richardson, C.J. All6gre, I. Sagna and L. Dimitriev, Noble gas isotopes from a topographic high on the MAR at 14 ° N, in preparation. 6 Th. Staudacher and C.J. Allrgre, Terrestrial xenology, Earth Planet. Sci. Lett. 60, 389-406, 1982. 7 C.J. Allrgre, Th. Staudacher, Ph. Sarda and M. Kurz, Constraints on evolution of earth's mantle from rare gas systematics, Nature 303, 762-766, 1983. 8 C.J. Allrgre, Th. Staudacher and Ph. Sarda, Rare gas systematics: formation of the atmosphere, evolution and structure of the Earth's mantle, Earth Planet. Sci. Lett. 81, 127-150, 1986/87. 9 Ch.C. Borch, J.G. Sclater, S. Gartner, Jr., R. Hekinian, D.A. Johnson, B. McGowran, A.C. Pimm, R.W. Thompson, J.J. Veevers and L.S. Waterman, Initial Reports of the Deep Sea Drilling Project, Vol. XXII, U.S. Government Printing Office, Washington, D.C., 1972. 10 T. Donnelly, J. Francheteau, W. Bryan, P. Robinson, M. Flower and M. Salisbury, Initial Reports of the Deep Sea Drilling Project, Vol. LI, LII, LIII, U.S. Government Printing Office, Washington, D.C., 1980. 11 C. Bollinger, Etude grochimique des phrnocristaux de plagioclases des roches ocraniques, Thesis, Universit6 Paris 7, 1979. 12 B. Hamelin and C.J. Allrgre, Large scale regional units in the depleted upper mantle revealed by an isotope study of the South-West Indian Ridge, Nature 315, 196-199, 1985. 13 J. Watkins, J.C. Moore, S.B. Bachman, F. Beghtel et al., Initial Reports of the Deep Sea Drilling Project, Vol. LXVI, U.S. Government Printing Office, Washington, D.C., 1979. 14 Th. Staudacher, M.D. Kurz and C.J. Allrgre, New noble gas data on glass samples from Loihi seamount and Hualalai and on dunite samples from Loihi and Rrunion Island, Chem. Geol. 56, 193-205, 1986. 15 K.P. Jochum, A.W. Hofmann, E. Ito, H.M. Seufert and W.M. White, K, U and Th in mid-ocean ridge basalt glasses and heat production, K / U and K / R b in the mantle, Nature 306, 431-436, 1983. 16 J.G. Sclater, B. Parson and C. Jaupart, Oceans and continents: similarities and differences in the mechanisms of heat loss, J. Geophys. Res. 86, 11,535-11,552, 1981. 17 J. Dymond and L. Hogan, Noble gas abundance patterns in deep-sea basalts--primordial gases from the mantle, Earth Planet. Sci. Lett. 20, 131-139, 1973. 18 F.A. Podosek, M. Honda and M. Ozima, Sedimentary noble gases, Geochim. Cosmochim. Acta 44, 1875-1884, 1980. 19 J.-I. Matsuda and K. Nagao, Noble gas abundance in a deep-sea sediment core from eastern equatorial Pacific, Geochem. J. 20, 71-80, 1986. 20 H. Craig, J.E. Lupton and Y. Horibe, A mantle helium component in circum pacific volcanic gases: Hakone, the Marianas and Mount Larsen, in: Terrestrial Rare Gases,
21
22
23
24
25
26
27
28
29
30
31
32 33
34
35
36
E.C. Alexander and M. Ozima, eds., pp. 3-16, Cent. Acad. Publ. Japan, Tokyo, 1978. Y. Sano, Y. Nakamura and H. Wakita, Areal distribution of 3He/4He ratios in the Tohoku district, Northeastern Japan, Chem. Geol. 52, 1-8, 1985. R. Poreda, Helium-3 and deuterium in back-arc basalts: Lau Basin and the Mariana Trough, Earth Planet. Sci. Lett. 73, 244-254, 1985. Y. Sano, Y. Nakamura, H. Wakita and T. Ishii, Light noble gases in basalt glasses from Mariana Trough, Geochim. Cosmochim. Acta 50, 2429-2432, 1986. K. Nagao, N.S. Takaoka and O. Matsubayashi, Isotopic anomalies of rare gases in the Nigorikawa geothermal area, Hokkaido, Japan, Earth Planet. Sci. Lett. 44, 82-90, 1979. C.J. Allrgre, S.R. Hart and J.F. Minster, Chemical structure and evolution of the mantle and continents determined by inversion of Nd and Sr isotopic data, II. Numerical experiments and discussion, Earth Planet. Sci. Lett. 66, 191-213, 1983. C.J. Allrgre and C. Jaupart, Continental tectonics and continental kinetics, Earth Planet. Sci. Lett. 74, 171-186, 1985. R.M. Garrels and F.T. MacKenzie, Sedimentary rocks, in: Evolution of Sedimentary Rocks, W.W. Norton&Company, New York, N.Y., 1971. R.J. Arculus, Island arc magmatism in relation to the evolution of the crust and mantle, Tectonophysics 75, 113-133, 1981. R.W. Kay, Volcanic arc magmas: implications of a melting-mixing model for element recycling in the crust-upper mantle system, J. Geol. 88, 497-522, 1980. C.J. Allrgre and Th. Staudacher, Terrestrial xenology, II, Japan-U.S. Seminar on Terrestrial Rare Gases, Yellowstone, Abstracts, pp. 1-4, 1986. R.A. Canalas, E.C. Alexander and O.K. Manuel, Terrestrial abundances of noble gases, J. Geophys. Res. 73, 3331-3334, 1968. D. Phinney, 36Ar, Kr and Xe in terrestrial materials, Earth Planet. Sci. Lett. 16, 413-420, 1972. T.J. Bernatowicz, F.A. Podosek, M. Honda and F.E. Kramer, The atmospheric inventory of xenon and noble gases in shales: the plastic bag experiment, J. Geophys. Res. 89, 4597-4611, 1984. E. Suess, B. Carson, S.D. Ritger, J.C. Moore, L.D. Kulm and G.R. Cochrane, Biological communities at vent sites along the subduction zone off Oregon, Bull. Biol. Soc. Washington 6, 475-484, 1985. J. Boulrgue, J.T. Iiyama, J.L. Charlou and J. Jedwab, Nankal Trough, Japan Trench and Kuril Trench: geochemistry of fluid samples by submersible "Nautile", Earth Planet. Sci. Lett. 83, 363-375, 1987. M. Javoy, F. Pineau and C.J. Allrgre, Carbon geodynamic cycle, Nature 300, 171-173, 1982.