Multiple sources of Quaternary tephra layers in the Mariana Trough

Journal of volcanology and geothennal research

ELSEVIER

Journal of Volcanology and Geothennal Research 76 (1997) 251-276

Multiple sources of Quaternary tephra layers in the Mariana Trough S.M. Straub

*

Department of Volcanology and Petrology, GEOMAR Research Center, Wischhofstrasse 1-3,24148 Kiel, FRG

Received 21 February 1996; accepted 12 September 1996

Abstract Tephra layers in Quaternary sediments ( < 1 Ma) in the central Mariana Trough contain three different series of glass shards: (1) low-K basaltic to basaltic-andesitic series (LKB: Sial = 47.3-53.6 wt.%; KlO = 0.2-0.8 wt.%); (2) mediumto high-K basaltic to dacitic series (HKS: Sial = 47.6-70.9 wt.%; K 2 0 = 0.8-3.2 wt.%); and (3) low-K dacitic to rhyolitic series (LKDR: Sial = 63.9-77.9 wt.%; KlO = 0.8-1.5 wt.%). LKB glasses are sideromelane shards which occur as minor component in two thirds of the tephra layers investigated. HKS glasses, mostly pyroclastic ally fragmented, dominate the discrete millimeter- to centimeter-thick fallout layers and are ubitiquous as dispersed shards. LKDR series comprise silicic pumice shards that prevail in decimeter-thick ash turbidites and crystal-poor, lapilli- to bomb-sized pumices. The composition of the LKB glass shards match those of coeval basalts and basaltic andesites of the Mariana Trough. LKB glasses are interpreted to have formed by the spalling of glassy rims of deep-submarine pillow and sheet lavas and subsequently became incorporated in the arc-derived tephra fallout and flow deposits. HKS and LKDR series both compositionally differ from coeval CIP lavas. Characteristic trace-element ratios, however, such as high large-ion-lithopileelements (LILE)/high-field-strength elements (HFSE) ratios (BalLa = 33-81; Cs/Nb = 0.3-0.8) and Nb depletion (La/Nb = 2-22) constrain their origin from explosive eruptions of the Central Island Province (CIP) of the Mariana arc. The relative enrichments of semi-incompatible elements preclude HKS and LKDR series from being linked by fractional crystallization and they are interpreted as derivative liquids from different magmatic series of the CIP. HKS glasses are suggested to be counterparts of CIP volcanos with medium-K trends (:::;; 1.5-2 wt.% K 2 0 at 60 wt.% Si0 2 ), whereas the relatively LILE-depleted LKDR series are derived from CIP volcanic islands and seamounts with lower-K trends ( :::;; 0.5-1.5 wt.% KzO at 60 wt.% Sial)' The magmatic diversity of the CIP - and thus the variability of the distal tephra deposits can be explained neither by crustal-level differentiation nor by variable degrees of partial melting of a homogenous mantle source. In view of the isotopic homogeneity of the CIP, it can best be explained by the Stolper and Newman (1994) model of binary mixing of source components, with one endmember being the NMORB source of Stolper and Newman (1994) and the other one being a compositionally variable slab-derived H 2 0-rich component similar to the H 2 0-component of Stolper and Newman (1994). This interpretation implies that the variation in K 2 0-content in the arc-derived glasses (about a factor

, Present address: Lamont Doherty Earth Observatory of Columbia UniverSIty, Palisades, NY 10964, USA. TeL: +914-365-8662; fax: + 914-365-8155; e-mail: [email protected]. 0377-0273/97/$17.00 Copyright © 1997 Elsevier Science B.V. All rights reserved. PII S0377-0273(96)00075-3

252

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (J997) 251-276

of 3 at 60 wt.% Si0 2 ) reflects a magmatic diversity that is inherent to the central Mariana arc and is not necessarily reflective of temporal changes of the composition of arc magma sources. Keywords: Mariana arc/backarc system; Quaternary submarine tephra; arc magma diversity

1. Introduction

2. Geologic setting

Tephra layers embedded in marine sediments play a key role in deciphering the long-term geochemical evolution of oceanic arc /backarc systems, because they may be the only tool for reconstructing the petrogenetic evolution of volcanic terrains that have been eroded or are buried under younger series (e.g., Fisher and Schmincke, 1984; Arculus et aI., 1995). The petrogenetic significance of submarine tephra layers, however, is not always easily interpreted. Marine tephra layers - or their vitric fraction by which they are commonly characterized - are known to be compositionally heterogenous. This may reflect pre-eruptive zoning of the magma chambers, syn- or postdepositional mixing, or mixing of products from different contemporaneous eruptions. Moreover, glass shard compositions likely represent derivative liquids of source area magmas, the evolved compositions of which obscure essential information about the magma sources and their evolution. As ancient settings are irrevocably destroyed, case studies in modem arc /backarc settings are needed to establish the link between proximal lava and distal tephra compositions. In this study from the modem Mariana arc /backarc system, tephra fallout layers, an ash turbidite and three lapillit- to bomb-sized pumice fragments from sediment cores in the backarc basin are described. The samples are less than 1 Ma old and thus coeval with the most recent period of volcanism in the Mariana arc /backarc volcanism. Built on the synthesis of major- and trace-element compositions of glasses and source area lavas, fragmentation mechanisms, depositional structures and spatial distribution of the tephra deposits, it can be shown that both the backarc spreading center and Mariana arc volcanic front contribute to the tephra formation. The arc-derived tephra series, though evolved, are inferred to reflect the diversity of the central Mariana arc, which in tum is attributed to heterogeneity of the arc magma sources.

Volcanism in the intra-oceanic Mariana arc /backarc system is associated with the subduction of the Pacific Plate beneath the Philippine Sea Plate in a westerly direction. Since the initiation of subduction in the Eocene (:::; 49 Ma), the volcanic arc was twice split (at :::; 30 Ma and :::; 8-5 Ma) followed by the formation of back arc basins (e.g., Taylor, 1992). The present arc/backarc setting comprises from east to west (Karig, 1970: (1) the Mariana Trench; the (2) Mariana Forearc; (3) the Mariana volcanic arc; (4) the Mariana Trough; and (5) the West Mariana Ridge, an extinct remnant arc (Fig. O. The crescent-shaped Mariana volcanic arc consists of nine subaerial and more than 30 submarine volcanic edifices (Bloomer et aI., 1989a). The arc basement (:::; 20-25 km thick; Sager, 19810) formed by young backarc crust (:::; 5-3 Ma) and older forearc crust ( < 45 Ma) is entirely oceanic. It is composed of tholeiitic and calc-alkaline basalts to andesites and lavas similar to boninites interspersed with subordinate dacites and rhyolites as well as calcareous and volcanigenic sediments (e.g., Schmidt, 1957; Meijer, 1980; Meijer et aI., 1981; Wood et aI., 1981; Reagan and Meijer, 1984). Based on geographic distribution and lava compositions, the Mariana arc is subdivided into four provinces (e.g., Fig. 1; Bloomer et aI., 1989a). The Central Island Province (CIP) comprises the largest and possibly oldest volcanos, with about half being subaerial. The CIP extends towards the north and south in chains of active seamount volcanos, named Northern Seamount Province (NSP) and Southern Seamount Province (SSP). The northernmost NSP (N-NSP), which merges with the southernmost part of the lzu-Bonin arc at 24°N, is also named the Alkaline Volcanic Province (AVP), owing to the occurrence of high-K shoshonitic lavas (e.g., Stem et aI., 1989; Bloomer et aI., 1989a,b). The Mariana arc is bordered in the west by the Mariana Trough, an actively opening

253

S.M. Straub / Journal o/Volcanology and Geothermal Research 76 (1997) 251-276

a N Izu-Bonin Is:, '.

Mariana Is. ;

..

'

b

1440

E

'7 ::G

r, I '.

1·.'1

c::) \

Fig. 1. (a) Regional setting of Marina arc/backarc system with subdivisions into provinces from Bloomer et aI. (1989a). '" denote seamount volcanoes, b. subaerial volcanoes. Islands of frontal arc are indicated by capital letters. Solid black line marks the position of Mariana Trough spreading center, and stippled line the axis of neovolcanic zone at '" 22-24°N. Numbers denote CIP seamounts: 1 = Northwest Uracas; 2 = Makhahnas; 3 = Ahyi; 4 = Supply Reef; 5 = Cheref; 6 = Poyo; 7 = Daon; 8 = West Guguan; 9 = Zealandia Bank; 10 = West Sarigan. (b) Detail of the sample area in the axial rift zone of the Mariana Trough (modified from Hawkins et a!., 1990). Solid black hne marks active rift axis; thin stippled line shows the extent of arc-derived volcaniclastic mass flows. Black bars give relative length of sediment cores (longest core 61KL is 4 m). PFZ = Pagan Fracture Zone; SR = Seatar Rift Valley; RR = Rama Rift Valley.

254

S.M. Straub / Journal o/Volcanology and Geothermal Research 76 (1997) 251-276

backarc basin with an average water depth of 4000 m. Between 14 and 20 0 N, a mature spreading center exists with a central ridge rising up to 3000 m water depth. The ridge is strongly segmented by alongstrike normal faulting and perpendicular, E-Wtrending, transform-like offsets, and has a well-developed axial graben which descends locally to more than 5000 m water depth (e.g., Hawkins et aI., 1990). Since the Early Pliocene, active volcanism in the Mariana arc /backarc system has been confined to the Mariana Trough and the arc volcanic front. The modem arc erupts low- to medium-K basaltic through dacitic compositions; basaltic andesites are the dominant lithology. Alkaline rocks arc confined to the AVP and the Kasuga seamounts, a chain of seamounts which are located on the troughward-side of the S-NSP at 21°30'N (Fig. 1; Stem et aI., 1988; Jackson, 1989). Mariana Trough lavas are sub alkalic basalt and basaltic andesites with trace-element characterisitics that are transitional between MORB and Mariana arc lavas (e.g., Hart et aI., 1972; Fryer et aI., 1981; Volpe et al., 1987; Hawkins et aI., 1990; Stem et aI., 1990; Gribble et al., 1996).

central Mariana Trough at 18°N at a distance of 100 km west from the Mariana islands (Table 1; Fig. I). The sample area extends about 100 km along the axis of the backarc spreading center. It can be subdivided into two parts: (I) Seatar Rift Valley (18°20' to 17°50'N): Six sediment cores and the rock dredge are located within a 60-km-Iong elongate depression of the axial graben. The Seatar Rift Valley, surrounded by :s; 1000 m rift mountains, is effectively shielded from arc-derived volcaniclastic mass-flows. Any arc-derived tephra must have been subaerially transported or floated. The rift floor is covered by a thin veneer of brown to gray-green, siliceous, carbonate-free clay-rich silts that are interspersed with discrete layers of dark brown volcanic ash a few millimeters to centimeters thick. X-ray images confirmed the presence of tephra fallout layers in cores 36 KG, 46 KL and 58 KL by diagnostic sedimentary structures such as sharp bases, grain size grading and either sharp or transitional tops. Primary sedimentary structures in ash-rich core 57 KL have been destroyed by bioturbation. (2) Pagan Fracture Zone /Rama Rift VaHey (south of 17°30'N). Core 59 KL is located in the Pagan Fracture Zone and cores 60 KL arld 61 KL in the adjacent Rama Rift Valley. The K-W-trending Pagan Fracture zone is a transform-like fault offsetting the southern ridge axis to the west. It forms a depression through which arc-derived mass flows are channeled into the Rama Rift Valley which is not shielded by rift mountains. The direction of mass

3. Sample locations

During RV SONNE cruise SO 57 in 1988, tephra was recovered by means of box cores (n = 7 stations), spade core (I) and rock dredge (1) in the

Table 1 Sample locations and number of tephra samples investigated Station

Location

Water depth

Core length

Number of samples

(No.)

(latitude/longitude)

(m)

(em)

fallout tephra ash turbidite pumice lapillijbomb surface samples

27DS 36 KG 46KL 57 KL 58 KL 59KL 60KL 61 KL

18°13.22'N/144°40.84'E 18°10.58'N/144°41.48'E 18°02.65'N/144°46.83'E 17°54. 13'N/144°44.05'E 17°56.l4'N/144°50.31'E lr37.14'N/144°49.05'E 17°25.56' N /144°46.44' E 17°28.23'N/144°48.45'E

3860 4090 4700 3860 4680 3980 4866 4690

37 105 340 225 110 335 430

3 12

a

13

In the text, core samples are identified by numbers: the first two digits giving the core number and the remaining the depth below seafloor (in em). a Volcaniclastic sediment from top of cores.

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (J997) 251-276

flow is corroborated by tephra grain sizes, which decrease from medium to coarse ash in core 59 KL to silt and fine sand sized ashes in cores 60 KL and 61 KL. The presence of foraminifera tests also indicate that the turbidites are derived from shallower water depths. Tephra occurs mainly as 20 to > 60 cm thick dark ash turbidites (diagnosed by Bouma cycles B-E) alternating with pelagic clay and silts, that contain millimeter-thick blackish tephra fallout layers as well.

4. Samples Thirty-one fallout layers, one turbidite and three lapilli- and bomb-sized pumices were sampled for analyses (Table 1). Fallout layers (including three surface samples) were sampled only from the Seatar Rift Valley. The layers are composed of 80 to 100 vol. % volcanigenic particles, the remaining fraction being silicic bioclasts. Volcanigenic particles comprise greenish brown, brown and colorless glass shards (50-80 vol.%), tachylites (up to 20%), volcanic lithic fragments (up to 10%) and intratelluric phenocrysts (10-15 vol.%). Phenocrysts consist of subequal portions of plagioclase and clinopyroxene, in addition to subordinate titanomagnetite and rare orthopyroxene. A single 5 mm pumice lapillus occurs in core 58 KL fallout layer in 34 cm core depth (sample 5834). The ash turbidites of the Pagan Fracture Zone and the Rama Rift Valley are composed of variable amounts of essential and accidental volcanigenic particles (estimated from particle freshness and rounding) and a subordinate biogenic fraction. Cores 60 KL and 61 KL have four 20-60 cm thick ash turbidites which are distinguishable from core 59 KL turbidites and Seatar Rift fallout tephra by colorless pumice glasses. The colorless pumice shards contain also sparse phenocrysts of green, euhedral intratelluric amphibole (Table 3) in addition to minor quantitites of plagioclase, pyroxene and oxide minerals. From the thickest turbidite (> 60 cm) in core 61 KL (bottom at 379 cm core depth) four samples were taken in 13-cm depth intervals. A pumice lapillus in 269 cm depth in core 57 KL (sample 57269) and a fist-sized reddish-beige pumice bomb from rock dredge 27 DS (sample 27 DS) were also sampled. The pumices have plagioclase, pyroxene and oxides

255

«

15 vol.%) set in a matrix of vesicular glass. No amphibole was observed. The background sedimentation rate was determined to be 0.19 ± 2 cm/ka by z30 Th in core 46KL (analyst J. Scholten). Extrapolation of this rate to all cores investigated, yields a maximum age of approximately 1 Ma for the background sediments. Thus, the deposition of the interbedded tephra layers is contemporaneous with the youngest period of volcanic activity of subaerial and submarine CIP volcanos, that started at about 5 Ma in the Late Pliocene (Packham and Williams, 1981). About 30 glass shards from each tephra layer and several fragments of the glassy matrix of crushed pumices were mounted and polished for electron microprobe analysis (EMPA). Glass shards from the compositionally least heterogenous basaltic-andesitic layers (samples 3622, 5888 and 5894 with less than a 6 wt. % range in Si0 2 ) and the dacitic to rhyolitic glasses from ash turbidite 61379 (with"" 10 wt.% range in SiO z) were separated by centrifugeassisted gravitational separation based on sodium poly tungstate solution. The density interval between 2.50 and 2.65 g/cm3 was used to separate the basaltic-andesitic glasses; dacitic to rhyolitic glasses were concentrated using a 2.3 g/ cm3 liquid. After repeated rinsing with distilled water, about 6-10 mg of the vitric fraction of the tephra layers and 20-50 mg of the glassy matrix of the pumices were taken for trace-element analyses by inductively coupled plasma mass spectrometry (ICP-MS). Details of the EMPA and ICP-MS analytical procedures are given in Straub (1995). Results are listed in Table 2. 4.1. The Mariana arc / backarc data base

Analyses of about 523 samples from the Marina arc and the Mariana Trough (AVP:n = 25; S-NSP:81; CIP:209; SSP:27; Mariana Trough Basalts: 191) were compiled from Schmidt (1957), Hart et al. (1972), Larson et al. (1974), Dixon and Batiza (1979), Steru (1979), Fryer et al. (1981), Wood et al. (1981), Meijer and Reagan (1981, 1983), Meijer (1982), Dixon and Steru (1983), Hole et al. (1984); Stem and Bibee (1984), Volpe et al. (1987), Woodhead et al. (1987), Woodhead (1988, 1989), Bloomer et al. (1989b), Jackson (1989, 1993), Lin et al. (1989), Stem et al. (1989, 1990, 1993), Hawkins et al.

S.M. Straub /loumal of Volcanology and Geothermal Research 76 (1997) 251-276

256

Table 2 Major- and trace-element data investigated LKB average n= 87 a

Major elements (wt. %) Si0 2 50.2 ± 1.1 Ti0 2 1.3 ± 0.3 Al 2 0 3 15.7 ± 0.5 FeO T 8.0 ± 0.6 MnO 0.16 ± 0.04 MgO 3.5 ± 0.5 CaO 10.5 ± 0.8 3.5 ± 0.5 0.4 ± 0.1

Total

96.70

Trace elements (ppm) Rb Sr Y Zr Nb Mo Cd Cs Ba Hf Tl Pb Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb

Lu

3622 5888 5894 basaltic-andesIte glasses

5834B dacite lapillus

57269 dacite lapillus

272 dacite bomb

61Tur rhyolitic glasses

53.8 1.2 14.4 10.9 0.23 2.8 7.0 3.4 1.9 0.42

54.6 1.2 14.4 8.4 0.24 2.6 6.7 3.3 1.9 0.43

55.6 1.2 14.6 7.3 0.23 2.6 6.5 3.3 2.0 0.39

65.7 0.8 14.6 5.7 0.16 1.2 3.6 3.6 2.8 0.28

64.9 0.5 15.6 5.7 0.20 0.8 3.7 2.9 1.8 0.18

71.6 0.6 12.9 4.2 0.16 0.7 3.5 2.3 0.16

73.1 0.3 12.4 2.3 0.13 0.3 2.1 1.3 1.3 0.15

96.04

93.73

93.66

98.53

96.40

97.49

93.35

36 292 33 117 2.8 2.4 0.6 1.3 444 3.20 0.07 8.7

34 292 33 120 2.4 4.2 0.7

34 299 33 112 3.4 2.6 0.5 1.4 428 3.28 0.07 7.2 1.4 0.7 12.00 26.24 4.06 18.18

51 222 39 208 7.0 3.2 0.3 1.7 451 5.53 0.23 6.6 2.6 1.4 15.48 34.53 4.90 22.24 5.95 1.48 6.50 1.07 6.92 1.44 4.31 0.64 4.31 0.64 17

34 185 43 156 5.7 2.1 0.3 1.5 440 4.33

20 174 40 125 4.1 2.1 0.2 1.2 428 3.76 0.12 4.2 0.9 0.4 8.04 19.58 3.11 15.43 4.82 1.17 5.67 1.02 6.81 1.45 4.44 0.67 4.55 0.69

20 143 36 70 0.4 2.3 0.3 0.3 667 2.18 0.02 3.9 0.8 0.6 8.23 19.36 2.97 13.25 4.02 0.92 4.71 0.87 5.85 1.29 3.99 0.62 4.33 0.65

21

11 4 3 5

1.5

BR

b

BR C

this work

reference values

46 1332 26 271 146 2 0.4 0.9 1044 5.7 0.1 4.4 10.1 2.6 80 145.0 16.3 65.5 11.61 3.59 10.2.4 1.30 6.39 1.02 2.64 0.29 1.77 0.23 22 358

45 1264 26 260 116 2.4

d

1.5 0.8 12.43 27.58 4.04 19.32 5.46 1.59 5.84 0.99 6.31 1.30 3.83 0.57 3.85 0.55 29

Sc Cr Co

26

Ni

8

Cu

240

9

1.2 493 3.41 0.05 7.8 1.5 0.8 12.27 27.22 4.15 19.62 5.56

1.61 5.99 0.99 6.36 1.31 3.86 0.55 3.76 0.56 28 13

25 12 234

5.22 1.53 5.64 0.94 6.04 1.24 3.65 0.55 3.68 0.54 29 10 27

3 9

12 237

0.17 5.0 1.8 0.8 13.26 29.64 4.37 20.48 5.67 1.43 6.41 1.08 7.21 1.57 4.69 0.70 4.83 0.71 19

23

10 4 70

2 9 5 89

52

61 260 76

0.82 998 5.8 4.5 9.8

2.5 77 144.0 16.2 61.9 11.60 3.47 9.56 1.24 6.05 1.00 2.50 0.29 1.77 0.24 21 345 59 255 71

257

S.M. Straub / Journal o/Volcanology and Geothennal Research 76 (1997) 251-276

4 high-K

•

3

j. / /"

... ..

,... .... ..... :•

~

/

<)

,/

'"

•

medium-K

~

"

1

v

•

0"·'. '. /<)

2

•

CIP

1

LKDR tephra

+

~/~ .• D&SS3 L&H85 • •. . +: + +~-+--:I•

. +if~e + + + ++++ + +

+.- •

+

S

50

SA

V

57269

A

D272

+

61379

(Bulk sample)

+

low-K

o

HKS glass shards HKS glasses analyzed for trace elements

1;\

",.~.,

"'

LKS glass shards

/

Individual glass shards of 61379 ash turbidite

A

60

70

80

Si02 wt% Fig. 2. K zO-silica variation diagram illustrating the compositional range of glasses investigated and lavas (shaded field) from the Central Mariana arc (eIP) and the Mariana Trough at 15 to 18°N (MTB 15-18°N). Circled fields at SiO z > 63 wt.% are compositions of bomb-sized pumices recovered from an unnamed seamount west of the forearc island Rota (D&S83 = Dixon and Stem, 1983) and the central Mariana Trough z 60 kIn northwest from the Seatar Rift Valley (L&H85 = Lonsdale and Hawkins, 1985). For other data sources see text. B = basalt; BA = basaltic andesite; A = andesite; D = dacite. Grid based on LeMaitre (1989).

(1990), Wolff (1990), Plischel (1992), Gribble (1996), Gribble et aL (1996) and Elliott et aL (1996). Most of the samples were analyzed for major elements and about half of them for trace elements and various isotope ratios. Although pyroclastic rocks are reported to occur on all islands (e.g., Dixon and Batiza, 1979; Stem, 1979; Meijer, 1982) and submarine arc volcanos (Bloomer et aI., 1989a; Jackson, 1993), few compositional data exist (e.g., Banks et aI., 1984; J.D. Woodhead, pers. commun., 1993), and to date no detailed study of pyroclastic material

recovered directly from the arc volcanic front has been carried out. Vitric tephra from piston cores and DSDP drill sites have been studied by Packham and Williams (1981), Schmincke (1981), Lee et aI. (1995) and Straub (1995). The compositional data were taken as published with the exception of the heavy rare earth element (HREE) analyses from Woodhead (1988, 1989). These data are omitted because they are systematically lower compared to other samples from the eIP [(Gd/Yb)" = 2.10 ± 0.28 (n = 19) compared to 1.26 ± 0.13 (n = 24) from Lin et aI.

Notes to Table 2: Trace-element analyses were only available for two of three glass shards groups identified (HKS and LKDR glasses). The trace-element concentrations of the three basalt andesitic glass separates of the HKS glasses are similiar within analytical error. a n = 18 for MnO. b Concurrent analysis of USGS Standard BR (basalt). C From Garbe-Schonberg (1993). d Analyst: C.D. Garbe-SchOnberg.

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

258

(1989), White and Patchett (1984) and Hole et al. (1984»). Light rare earth elements (LREE) however, are indistinguishable: (CejSm)n = 1.13 ± 0.16 as given by Woodhead (1988, 1989) compared to (CejSm)n = 1.08 ± 0.19 from other workers.

5. Results 5.1. Tephra characterization

Compositions of glass shards of tephra layers and the glassy matrix of pumices investigated ('glasses'

2

21

(V')

oN

17

«

~~~____~____~__~____ .~

LKDR tephra

3

~

~

3: oN

'i1

57269

A

0272

14

10

+

2

0

61379 (bulk sample)

oj(

6

~

o(1) l.Ji-

"

2

61379 glasses

r---~----~----r---~-----~15

•<>

(V')

oN

~ oro

MTB1S-18°N LKB glasses

10

oN

0.6

, o

~

....

5

U

N

..-=

0.2 .j=.

o

2

4

6

8

10

o

~':-

2

4

6

8

10

MgO wt% Fig. 3. Major-element variation of the three glass series compared to lavas from the central Mariana Trough (MTB l5-l8°N) and the Mariana Central Island province (CIP). Shaded field indicates single glass shard compositions of ash turbidite 61379. MTB lavas and trough-derived LKB glasses overlap in compositIOn, whereas parts of HKS glasses (dash-enclosed field) and LKDR plot outside the elP compositional field. For data sources see text.

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

for short) range from basalt through rhyolite (47-78 wt.% SiOz)' The glasses follow low-K to high-K trends (Fig. 2). Based on compositions, colors, morphologies and distribution patterns, the glasses were grouped into three series: 1. LKB - low-K basalts and basaltic andesites (Si0 2 = 47.2-53.6 wt.%; K 2 0 = 0.3-0.8 wt.%); 2. HKS - high-K basalts through dacites (SiO z = 47.6-70.8 wt.%; K 2 0 = 0.8-3.2 wt.%); 3. LKDR - low-K dacites through rhyolites (SiO z = 63.9-77.9 wt.%; K 2 0 = 0.8-2.1 wt.%). Major-element concentrations and trends of basaltic and basaltic andesitic LKB glasses match those of the Mariana Trough basalts and basaltic

259

andesites at I5-I8°N (= MTB I5-I8°N in the following; Fig. 3). LKB glasses are typical green brown to brown sideromelane shards with dense, blocky or splinter-like forms that have rare spherical vesicles and euhedral feldspar microliths (Fig. 4a). LKB glasses occur only as a minor group with up to 30 vol. % in two thirds of the fallout layers and ash turbidites. They are interpreted to have formed by spalling of glassy rims of deep submarine lava flows (3000-5000 m water depth) and subsequently became incorporated with the tephra fallout layers and turbidites during their deposition. On two-element plots, LKB compositions overlap with those of the mafic HKS glasses (Fig. 3). However, low K 2 0

a

d

9

Fig. 4. Scanning electron microscope photographs of glasses investigated (note different scales). (a) LKB sideromelane. (b-g) Varieties of HKS glasses. Glass shards (b-f) are transparent, crystal-poor brown and green colored basalts through andesItes. Blocky crystal-poor shard with spherical isolated vesicles (b) and blocky shard with large irregular vesicles (c) indicate hydroclastic fragmentation. Note concentric pattern on shard in (c) that typically occurs on the surface of violently burst-apart solids, in this case perhaps as a consequence of thermal shock. Shard (e) is a bubble-wall shard which has an intact spherical vesicle in its center. The twisted droplet form (f) is typical for deformation of lOW-VIscosity fluid dunng subaerial transport. Shard in (g) is a slightly yellow tinted clear pumice shard characteristic for HKS dacites. (h) Typical assortment of high-silica LKDR shards. A single bubble-wall shard is indicated by an arrow.

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

260

« 1 wt.%; HKS glasses mostly> 1 wt.%), low FeO* (::::: 7-10 wt.% FeO*; HKS = 10-11 wt.%), high TiOdK20 ratios (3.7 ± 1.2; HKS = 0.7 ± 0.2), high S (> 1500 ppm; HKS < 1000 ppm) and the typical sideromelane morphology are diagnostic when compared with mafic HKS glasses. HKS glass shards are the dominant glass population in the fallout layers of the Seatar Rift Valley. They also occur in minor quantities in the turbidites of the Rama Rift Valley, where they have rounded forms, more altered appearance, and are interpreted as accidental clasts enveloped during turbidite flow. HKS glasses are 10w-MgO basalts (6 wt.% MgO maximum) through dacites with a dominant mode of basaltic andesitic and a subordinate mode of dacitic compositions. Basaltic to andesitic HKS glasses are brown to brown green and show a wide range of dense, blocky, vesicle-poor up to bubble-wall and highly vesicular morphologies which indicate both pyroclastic and hydroclastic fragmentation (Fig. 4bTable 3 Microprobe analyses (wt.%) of amphibole in ash turbidite 61379 A

B

SiOz TiO z Al z0 3 FeO' MnO MgO CaO NazO KzO PzOs CI

47.68 1.62 7.78 12.68 0.56 15.43 11.16 1.54 0.21 0.17 0.89

48.35 1.01 7.18 12.68 0.74 15.92 10.43 1.23 0.16 0.15 0.06

Total

99.72

97.85

No. of ions on basis of23 oxygens Si 6.7163 Ti 0.1716 AIOV) 1.2837 AI(V!) 0.0079 1.0448 FeOn) 0.4489 FeO!) 0.6681 Mn Mg 3.2396 1.6843 Ca Na 0.4206 0.0377 K P 0.0203

6.7835 0.1066 1.1872 0.0000 1.4803 0.0075 0.0879 3.3292 1.5678 0.3346 0.0286 0.0179

g). Dacitic HKS glasses are very light brown and yellow pumice shards (Fig. 4f). HKS glasses were interpreted by Straub (1995) as the distal counterparts of basaltic to dacitic lavas of the central Mariana arc (CIP) that follow higher-K trends (about 1.2-2.0 wt.% K 20 at 60 wt.% Si0 2). The distinctive compositional contrast with source area lavas, as given by (a) lower A1 203, CaO and MgO contents, (b) higher average concentrations of incompatible elements, and (c) differing trends of Ti0 2, K 20, FeO * and P20 S (see also fig. 4 in Straub, 1995) are considered to result from closed-system convective crystallization in crustal magma chambers; they do not reflect different, enriched source compositions (for detailed discussion see Straub, 1995). LKDR glasses comprise glass shards from Rama Rift Valley turbidite 61379 and the glassy matrix of lapilli- and bomb-sized pumices 27DS and 57269. The glasses in turbidite 61379 have mostly highly vesicular, pumiceous shards; bubble wall shards are rare (Fig. 4). In contrast to the slightly tinted dacitic HKS pumice glasses, 61379 glasses are colorless and also have rare euhedral green amphibole phenocrysts (Table 3) which permit distinction from dacitic HKS glasses despite similar shard morphologies. The major elements of the glasses investigated compare well with results from previous studies of Quaternary tephra layers from DSDP Leg 59 and 60 drill cores (Schmincke, 1981; Packham and Williams, 1981; Lee et aI., 1995). LKB glass compositions correspond with 'group II' glasses found by Lee et al. (1995) in backarc drill cores, thus supporting the Lee et al. (1995) view, that these glasses are hyaloclastite debris from Mariana Trough lavas. HKS and LKDR glasses compare to the bulk of the glasses reported by Packham and Williams (1981) and Lee et al. (1995) which are considered to originate at arc front volcanos. No shoshonitic compositions were found in this study like those that occur in Miocene (11-8 Ma) and Early Oligocene (33-30 Ma) sediments (Lee et aI., 1995). Both the trace-element patterns of HKS and LKDR glasses investigated are characteristic of subductionrelated magma series with high LILE (large ion lithophile) jHFSE (high field strength elements) ratios such as BajLa::::: 33-81 (N-MORB and OlB < 0; BajZr::::: 3-9 (N-MORB < 0.1; OlB::::: 1.25) (Table 4). LILE (K, Ba, Rb, Th and U) are enriched

S.M. Straub j Journal of Volcanology and Geothennal Research 76 (1997) 251-276

261

Table 4 Incompatible trace-element ratios of Mariana arc provinces CIP and AVP and central and northern Mariana Trough (MTB l5-l8°N and MTB 22-24°N) compared to HKS and LKDR glasses. For data sources see text

CsjNb BajLa BajZr BajSm KjZr (Ce/Yb)n (Ce/Sm)n Zr/Y Zr/Sm a b

AVP a

CIP

n,,;,7 20± 2 15 ± 11 102±9 385 ± 286 8.7 ± 0.6 2.8 ± 0.2

n,,;, 113

HKS glasses n=4

LKDR glasses n=3

Mariana T. 15-l8°N n";' 114

0.33 ± 0.11 38 ± 10 3.4 ± 0.8 77 ± 20 92± 23 1.5 ± .06 1.2 ± 0.3 3.0 ± 0.5 23 ±4

0.40 ± 0.11 35 ±5 3.5 ± 0.9 82±5 132 ± 15 2.1 ±0.1 1.3 ± 0.1 4.0 ± 0.9 25 ±7

0.45 ± 0.29 35 ± 23 5.3 ± 3.7 111 ± 48 118 ± 35 2.1 ± 0.3 1.3 ± 0.2 2.9 ± 0.8 24±6

0.04 ± 0.03 9±4 0.5 ± 0.2 15 ±7 40 ±22 1.4 ± 0.3 0.9 ± 0.4 3.4 ± 0.5 28 ±4

b

Mariana T. 22-24°N n,,;, 27

23 ± 13 1.26 56 ± 25 47 ± 28 1.9 ± 0.6 1.4 ± 0.30

AVP = Alkaline Volcanic Province. CIP = Central Island Province.

up to 100 times, as compared to N-MORB, while Nb is depleted relative to other elements with a similar degree of mantle incompatibility (e.g., La/Nb "'" 222; N-MORB and OIB < 1; Fig. 5). HFSE are either

III CC

slightly enriched (about 5 times maximum for HKS glasses) or equal to N-MORB (LKDR glasses) despite the high silica content of the LKDR glasses. The rare earth elements (REE) form smooth patterns

•

HKS basaltic-andesite HKS dacite

()

100

LKDR tephra 57269

o

D272

~

+

I

Z

61379 (Bulk sample)

........... (J)

a.

E CO

(f)

0.1

+

+

MTB15-18'N b

J

l J I

8r

Rb K

~

J i

Th Sa

-'

~

Nb U

L J

i

Ce

La

Nd

p

I

I

J J

J

J

J

I

I

J

Zr Eu Gd Y Yb 8m Hf Ti Dy Er Lu

Fig. 5. Trace-element concentrations of HKS and LKDR glasses compared to maximum range of lava compositions from Mariana Central Island Province (elP) and Mariana Trough at IS-18°N (MTB 15-18°N). Normalized by N-MORB compositions from Sun and McDonough (1989); elemental order based on Pearce (1983).

262

S.M. Straub /loumal o/Volcanology and Geothermal Research 76 (1997) 251-276

with negative Eu anomalies (Eu/Eu * z 0.7) that have 20-40 times chondritic concentrations. LKDR glasses are less enriched in REE [(Ce/Yb)n = 1.21.7] than HKS glasses [(Ce/Yb)n = 2.1 ± 0.1]. Consistent with the higher silica content of the LKDR glasses, transition metals Sc, Cr, Ni and Cu have lower concentrations than HKS glasses (Table 2). In the following, the term semi-incompatible is used for trace elements which are generally considered as incompatible in the upper mantle, but may become compatible with respect to the fractionating mineral assemblage considered. They include Sr, Eu

(plagioclase), Ti (Fe-Ti-oxides), P (apatite), Y (pyroxene, apatite) and the REE (apatite, amphibole) (the full range of the semi-incompatible elements is given in Fig. 6b). Potassium is the only incompatible major element which varies significantly within the LKDR series and between HKS and LKDR series (Figs. 2 and 3). Figs. 3 and 6 show that the K trend mirrors the general trend of semi-incompatible elements: within the LKDR series semi-incompatible trace elements decrease with decreasing K]O. However, silica increases with decreasing K 2 0, which is opposite to the expected trend if Ll(OR glasses are

CD

() HKS daciite 0.6

LKDR tephra

La Ce Pr Nd 8m Eu Gd Tb Oy Ho Er Tm Yb Lu

a. E

\l

57269

.A

0272

+

61379 (Bulk samplle)

1.4

CO

en 0.6

0.2 TI Ba U K Ce Pr P 8m HI Ti Tb Y Er Yb Cs Rb Th Nb La Pb 8r Nd Zr Eu Gd Oy Ho Tm Lu

Fig. 6. Trace-element concentrations of LKDR tephra compared to HKS dacite. Elemental order after Sun and McDonough (1989) indicates decreasing upper mantle mcompatibility from left to the right. (a) REE only. (b) Complete range of semi-incompatible trace elements. Stippled line in (a) and (b) presents averaged compositIOns of the three basaltic-andesitic HKS glasses.

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

ble elements from the system via diffusion or vapor transfer is also unlikely.

on a liquid line of descent. At about the same concentration of Si0 2 , the K 2 0-richer HKS glasses have elevated contents of semi-incompatible elements compared to LKDR glasses. However, the relative enrichment of HKS glasses is not uniform: highly incompatible elements are more enriched (1/2 to 1/3) (with the exception of Ba) compared to the least incompatible elements (HREE, Y) (Fig. 6). The tilted incompatible trace-element patterns preclude a liquid line of descent relationship between HKS and LKDR glasses. Apatite and amphibole, the only phases that can incorporate LILE or REE cannot be the cause of the observed fractionation, as elements which are incompatible in amphibole and apatite, are depleted (e.g., Th, Cs) in the LKDR and those which are compatible are enriched (e.g., Y, HREE). As incompatible elements which are both mobile and immobile in hydrous solutions (e.g., K, Rb in contrast to Ce, Zr) follow the same enrichment and depletion trends, selective loss of highly incompati-

E c.. c.. 40

6. Discussion With the Mariana Trough and the Mariana arc identified as sources of the LKB and HKS glasses, the remaining question is: where does the LKDR tephra come from? Based on compositional similarities with melts produced in melting experiments on oceanic basalts, Lee et al. (1995) discussed the possibility of the low-K siliceous compositions produced by crustal anatexis. Major-element evidence alone, however, is inconclusive, and contrasting isotope ratios and trace-element behaviour of associated mafic and silicic suites are needed to argue in favor of crustal anatexis (e.g., Beard, 1995). The arc basement is interspersed with more fusible material like sedimentary rock or dacites, and it may be possible

<> ~ D~S83

..0

c::::

2

20

<> (jE):t::/ <> !(} i <><> i

V

6 L&H85 .0--<>

100

CIP

•

MTB1S-18°N at 18°N

•

o

<>

(> "Arc-like" lavas

•

o

263

8 HKS glasses LKDR tephra V

57269

~

D272

+

61379 (Bulk sample)

200

Zr ppm Fig. 7. Rb vs. Zr variation of HKS and LKDR glasses compared to lavas from the CIP and Mariana Trough lavas. Tephra compositions lie within the range of Rb /Zr ratios defined by e1P lavas with no evidence of excess Rb enrichment, an element strongly enriched in hydrothermally altered rocks. 'Arc-like' Mariana Trough compositions from l8°N are as reported by Hawkins et al. (1990). Other symbols are as in Fig. 2.

264

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

that the melting point of mafic and intermediate lavas was lowered by hydration with seawater or hydrothermal solutions. However, studies of major and trace elements and the limited isotopic variation of mafic through evolved Mariana arc lavas [e.g., a18 0 = 6.1 ± 0.3%0 (n = 45); 87 Sri 86Sr = 0.70344 ± 0.00009 (n = 113), including a medium-K dacite pumice from an SSP seamount with 87 Sri 86 Sr = 0.70353; Dixon and Stem, 1983)] give no indication of crustal anatexis, assimilation or selective contamination. Macdonald et aI. (1987) suggested that crustal contamination may be recognized by excess enrichment of elements that are incompatible in the system but enriched in hydrothermal solutions (e.g., Rb). In a Rb-Zr plot, however, the LKDR glasses follow the same trends as the elP lavas (Fig. 7) and give no evidence for derivation from hydrothermally altered crust. Volcanic rocks with the low-K to medium-K dacitic compositions similar to LKDR samples D272 and 61379 are not known for the subaerial arc, but silicic low- and medium-K tephra and fallout deposits are common in the Mariana backarc and fore-

3

• Pagan .... Anatahan

IE Cheref Sm. • Supply Reef

L::. Sangan Guguan

(>

o

arc basins (Packham and Williams, 1981; Dixon and Stem, 1983; Lonsdale and Hawkins, 1985; Lee et aI., 1995). These observations point to a submarine volcanic source for the LKDR tephra, with potential eruptive sites being either submarine volcanos at the arc front or at the backarc spreading center. Strong evidence exists that submarine explosive volcanism occurs at the arc front located about '" 100 km east from the sample area. Pyroclastic material is widespread on elP seamounts (Bloomer et aI., 1989a) and dacitic pumices have been dredged from seamounts in the S-NSP and SSP (e.g., Dixon and Stem, 1983; Jackson, 1993). Simi12ir to their subaerial counterparts, submarine elP volcanos range from basalts to dacites. A closer look at the K 2 0trends of the elP lavas shows that individual volcanos follow different K 2 0-enrichment trends: at a given Si0 2 = 60 wt.%, their K 2 0 content varies'" a factor of 4 (Fig. 8). The LKDR glass compositions lie at the silica-rich end of lower-K elP andesites and dacites suggesting that LKDR glasses are derivative liquids. On the other hand, recent results from ocean

medium-K

Poyo Sm.

L&H85

f) Agngan

./

LKDR tEiphra

1

\l

57269

...

D2?2

+

61379

(Bulk sample)

o S

SO

SA

A

60

D+R

70

80

SiOz wt% Fig. 8. Range of K 2 0-enrichment trends for individual arc volcanoes (for clarity, a selection is presented). Parental compositions used in model calculations are indicated (GUN: Si0 2 = 59.5 wt%/K 2 0 = 0.87 wt%; SUR13-22 = 59.3/0.68; SAP = 60.5/1.25). Stippled line indicates the fractionation trend for Maug volcano given in Meijer and Reagan (1983). For other symbols see Fig. 2.

S.M. Straub / Journal o/Volcanology and Geothermal Research 76 (1997) 251-276

drilling in the Lau Basin showed that part of the basin's infill of low-K silicic tephra did not originate from arc front volcanos but from intrabasin submarine volcanos, which grow on the rifted arc crust at ::::: 1000-2000 m water depth (Bednarz and Schmincke, 1994; Clift and Dixon, 1994). Though the greater water depth of the central Mariana Trough (3000-4000 m) is more likely to suppress explosive volcanism, at first sight a Trough origin of the LKDR tephra seems possible. LKDR plot also towards the end of the low-K trend of the MTB I5-18°N (Fig. 2), and could be a product of extreme crystallization (up to 90%) of a MTB 15-18°N parental composition. At the northern end of the Mariana Trough at 22-23°N where the backarc rifting zone appears to merge with the arc front, dacitic lavas occur which have K 2 0 contents similar to the LKDR samples D272 and 61379 as well as arc-like trace-element patterns and isotope ratios (Gribble, 1996). The large distance (::::: 500 km) to the depositional sites of the tephra investigated, however, excludes an origin from these sources. The only known evolved composition in the Central Mariana Trough is a dacite pumice, that was dredged from the west flank of the axial high at 18°N and is believed to have formed in situ (Lonsdale and Hawkins, 1985). The pumice sample, however, is coated by a 4-6cm-thick ferromanganese crust, that consists of slowly accumulating manganese oxide a-MnOz (Lonsdale and Hawkins, 1985). Assuming an aver-

I


..c

en Z

()

0.1

~

• • •

••

age growth rate of a-Mn0 2 ferromanganese crust of ::::: 10 mm/Ma (e.g., Halbach et aI., 1983a), this pumice erupted at least 4 Ma ago. This age is roughly consistent with an extrapolated age of the ocean crust (::::: 3 Ma) at this site (Lonsdale and Hawkins, 1985). Thus, this pumice cannot be taken as a product of the modem Mariana system. It very likely erupted at a time when arc and backarc sources were closer together, perhaps even before the Pliocene Mariana rift evolved into a spreading center. Because of the geochemical affinity between Quaternary Mariana arc and Trough lavas, geochemical tracers that unambiguously separate CIP and MTB 15-18°N compositions are rare. 'Arc-like' basaltic andesites recovered from 18°N in the Mariana Trough have the same Sr- and Nd-isotope ratios as CIP lavas (Volpe et aI., 1990) and ratios of incompatible trace elements, such as LILE/LILE and HFSE/HFSE widely overlap (Table 4; Fig. 11). Only some LILE/HFSE ratios, such as Ba/Zr or Ba/Sm completely separate Mariana arc and backarc compositions, with the LKDR tephra plotting in the arc field. However, with respect to the origin of high silicic tephra, this evidence remains ambiguous as dacitic and rhyolitic liquids are likely to be saturated in apatite and possibly zircon. Small amounts of these trace phases ( < 1 wt. % in the cumulate) could drive Ba/Sm or Ba/Zr ratios out of the range of MTB 15-18° compositions and into the field of the arc

I

o

265

+

I

Y0

0

•

...

0

CIP MTB 15-18°N HKS glasses

-:-

~ 0.01 '----_-'---_-'---_-'--_--'-_---'-_----' 50 70

LKDR tephra \j

...

+

57269 D272 61379 (Bulk sample)

Fig. 9. CIP-like Cs/Nb ratios of the tephra investigated strongly indicate provenance from arc front volcanos. To date, only subaerial ClP volcanoes are analysed for Cs and Nb (ClP data: Elliott et aI., 1996; MTB 15-J8°N data: Piischel, 1992).

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (J997) 251-276

266

lavas. To date, the best indicator of an arc vs. backarc origin of the LKDR tephra are the ratios of incompatible, highly fluid-mobile elements to incompatible, fluid-immobile elements, such as B /Nb, B /Be or Cs /Nb, because: (a) these ratios are assumed to be considerably higher at the arc front as a consequence of the higher fluid input beneath the arc front compared to the backarc (Ishikawa and Nakamura, 1994; Ryan et aI., 1995); and (b) the element pairs are incompatible with respect to the fractionation mineral assemblage of olivine, pyroxene, plagioclase, oxides, amphibole, apatite and zircon. To date, only Cs /Nb data are available for CIP and MTB 15-18°N lavas, but all LKDR compositions plot in the arc field, corroborating the arc front origin (Fig. 9). If both HKS and LKDR glasses are CIP-derived derivative liquids, their characteristic compositional diversity - as given by their tilted semi-incompatible trace-element patterns - must also be discernible in their source area lavas. To test this hypothesis, I selected Mariana basalts and basaltic andesites from the CIP which have MgO > 4 wt. % and Si0 2 < 57 wt. % (n = 96). The low cut-off at 4 wt. % MgO was necessary in order to include a reasonably large number of samples, since only a few of the arc lavas have more than 6 wt. % MgO. These samples (= 'primitive' arc lavas) are from eight volcanic islands [Agrigan (n = 10), Alamagan (5), Anatahan (3), Ascuncion (2), Guguan (9), Pagan (29), Sarigan (21), Uracas (1)] and six seamounts [Ahyi (3), Cheref (4), NW Uracas (2), Poyo (2), Supply Reef (2),

Table 5 Correlation coefficients between (Ce/Sm)" and incompatible elements for CIP lavas with SiOz < 56 wt.% and MgO > 4 wt.%. For data sources see text

Th Ce K 20 Rb

Sm Ba Zr Yb P Y

(Ce/Sm)n

n

0.82 0.81 0.62 0.59 0.42 0.36 0.33

12 54 54 53 54 54 46 35 54 44

O.1S 0.04 0.01

----,---

30

<> •

E a.. a.. Q)

CIP MTB15-18°N

20

• a

•

10

0

CIP r=O.80 MTB15-IBON r=O.93

0

-L-

b

.

30

>-

+ • o.:>S<>

•

10 0

•

•

••

•• ~ ••• +«:> <> ol/f;

E a.. 20 a..

0

<»

<> <>

<>

CIP r=O 01 MTB 15- IBoN r=0.45

1

2

(Ce/Sm)n Fig. 10. (Ce/Sm)" vs. Ce and Y diagram illustrating larger range of highly incompatible elements compared to less mantle incompatible elements in the CIP lavas with < 57 wt.% Si0 2 and> 4 wt.% MgO ('primitive' arc lavas; n = 54). Data from Mariana Trough lavas from 15 to IS oN (MTB15-18°N; n = 19) are shown for comparison. For data sources see text.

Zealandia (3)]. About half of them were analysed for trace elements (not all trace-element data are available for all samples). Low maximum Ni « 100 ppm) and covariance of Sc, Cr, Ni and - weakly - Eu/Eu' with MgO (rsc = 0.91., n = 12; rer = 0.88, n = 55; r Ni = 0.88, n = 68; rEu/Eu' = 0.48; n = 48; Eu/Eu' is the ratio of the measured Eu concentration over the Eu' concentration calculated from interpolation of neighbouring Sm and Gd) indicate that the 'primitive' arc lavas experienced fractionation of olivine, pyroxene and plagioclase (all modally observed) and are not primary compositions despite a maximum Mg' = 73 (calculated assuming 14% of the iron as Fe H ). However, the ratios of elements which are incompatible to this assemblage can be used to characterize these melts. Within the 'primitive' arc lavas, these ratios vary widely between centers as well as within a single center (Fig. 11; Table 6). Therefore, despite the isotopic homogeneity, significant magmatic diversity exists in the CIP. (Ce/Sm)n ratios of the 'primitive' arc lavas

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

correlate positively with the LREE and other LILE (such as K, Th and Rb, but to lesser extent with Ba), but have no correlation with the least incompatible elements (e.g., Y, Yb or p) (Table 5; Fig. 10). Thus, similar to the glasses, the 'primitive' arc lavas have a larger range in highly incompatible elements compared to lesser incompatible elements. Standard modelling procedures for major and trace elements were applied in order to further test for parent-daughter relationship between LKDR glasses and low-K arc andesites that have similar slopes of REE (see Appendix A for details). Mass balance calculations show that the LKDR glass compositions can be generated by approximately ~ 30-70% crystallization from lower-K arc andesites that have (CejSm)n ratios similar to those ofthe LKDR glasses (Table 7). The fractionating mineral assemblage consists dominantly of plagioclase (56-73 wt. %), followed by subordinate amounts of clinopyroxene (720 wt.%), orthopyroxene 02-15 wt.%), titanomagnetite (8-12 wt.%) and trace amounts of apatite and amphibole. This paragenesis is consistent with that observed in the CIP and with that in dacitic derivative liquids, which was obtained in fractionation experiments of oceanic basalts (Juster and Grove, 1989). The in-situ equation yields results within 10% maximum deviation on average between REE calculated and REE analyses in glasses. The Rayleigh equation projects excess REE enrichment for samples D272 and 61379 with the mass crystallized being about 20% too large for acceptable results. For the SAP-57269 assumed parent-daughter pair, both equations yield approximately the same results in LREE. In general, the model calculations, also by necessity simplistic, are consistent with the derivation of LKDR compositions from lower-K CIP andesites. CIP seamounts cannot be excluded a priori as sources of pyroclastically fragmented material. Submarine eruptions can eject pumice fragments, when the magma has vesiculated prior to eruption e.g., in a sealed magma chamber or conduit. Vesiculation of a melt in a sealed magma chamber or conduit can occur, when volatiles (dominantly HzO in arc melts) become satured and exsolve to form a separate vapor phase similar to the model suggested by Wallace and Gerlach (1994). Assuming a content of 1.5 wt.% Hz for Mariana arc basalts based on an estimate of

°

267

Baker (1987) and in comparison with Mariana Trough basalts (Stolper and Newman, 1994), dacitic and rhyolitic derivate melts evolving in a closed system will have 8-10 wt.% HzO. At pressures in an upper crustal magma chamber (~2 kbar corresponding about 6-7 km crustal depths) this amount of water is not soluble (Silver et aI., 1990) and bubble formation will occur. The hydrostatic pressure would only add up to 140 bar as CIP seamount peaks are shallower than 1300 m water depth. If the vapor exsolution is rapid (perhaps supported by second boiling, e.g., Burnham, 1983) the ensuing increase in internal pressure could lead to the catastrophic breakup of the reservoir, causing the magma (liquid + crystals + vapor) to rise to the seafloor. During ascent, the magma will further vesiculate by decompression, until the foam finally disrupts or the process is halted by quench fragmentation when seawater is encountered. Such a scenario could also explain why the dacitic and rhyolitic endmembers of the lower-K series form tephra, whereas less evolved (= less volatile-rich) basaltic through andesitic compositions still erupt effusively. Once in the sea, tephra becomes distributed by transport mechanisms largely similar to those of subaerial eruptions (e.g., Fisher and Schrnincke, 1984; Cashman and Fiske, 1993). It can either rise to form clouds that spread laterally driven by water currents or travel as density flows along the bottom. Daon seamount located on the west of the island Pagan could be a potential eruption center for ash turbidite 61379, although a genetic relationship would have to be ascertained by compositional data. An alternative possibility could be an unnamed seamount west of Guguan volcano which is located ~ 80 km south of Pagan. LKDR pumice lapilli and bombs in the Seatar Rift Valley are likely to have floated to their depositional sites. 6.1. Tephra variability and the magmatic diversity of the central Mariana arc

A major conclusion drawn from the preceding discussion is that the compositional diversity of the arc-derived glass series HKS and LKDR series reflects different magmatic series that occur in the CIP. The variability of the glasses is expressed by the variation in KzO, up to a factor 3 at 60 wt.% SiO z,

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

268

Williams, 1981; Lee et aI., 1995) showed that glasses with K 2 0 variations similar to the HKS and LKDR glasses occur throughout the 40 Ma history of the Mariana arc /backarc system [note, that a subordi-

which in turn reflects the variations in the slope of the semi-incompatible element patterns. The results of tephra studies from DSDP drill cores in the Mariana arc /backarc system (Packham and

Mk

L-

..c

Ay

0Q)

tT 1

~n

I--

/'

1

c

E

f-

0

~

Q)

...

•

~

•... ....

'V f-

U

.......... f-

o ....

f-

I I

I...

.

\

e\

-,

I

n

...

..........

ro

CO

t

to • r,iTR15--8 N

,

/

\

/

\

\

1

\

I

I

I

I

I

•

i/---

/

-~//

-I

\

/

J::.

\

I

••

•

\ \

o

\ \ \

....

I

20

"t,'~

./....

\

I

t

-,

\ ¢ \ ,(!'; (), ,,,

I

I

I-C

+

30

~ CO

l-

\

> 0

;

e

~

u '0

I

C1!

>

0

.......

1000

"ijj

>

~(I)

..ern 0)(1)

'(jj

U)

..eo o~

e(l)

0

> u:;::;

C1!

'O~

>~

E

'00",/

fl

-1000

-

20

l0

® 0

•i t

l-

o 16

I

I

I

I

-

...

JI.

....

10

..,i 1000

E ::I '0

r.:TB15 -8 N

o

.0 0:::

(I)

I

"

40

-

4

'

.......

'"E

..>:: '00",/

,

\

t

2000

-

4~

/\

\

0

t

-- ,

/

~\//: Il ....

I~

f)

....

(]

I'

t, ... t:

\/

\

.... .... ....

,

\

I

,, ,,

NU

III

6

I

\

0 0

~fi

I

b

0

40

...

, ,,

~s

(;;,

,,

\_/

+10

j

,,

,

-•-

, \

60

,,

1 1 1

.......... U) ..........

1

I

f-

?I

~g

j

a 2

u

C

~I

(J

i

6

••

I ...

0

I

I

~

...

r·HI:!15·' ii N

, •

f)

t -

~ JI.

...

•

l.

I

I

I

I

17

18

19

20

Latitude ON

t

- o HKS glasses -

21

LKDR t4:lphra V

57:269

~

0272

+

61a79 (Bulk sample)

S.M. Straub j Journal of Volcanology and Geothermal Research 76 (1997) 251-276

nate population (;:::;< 5%) of the glasses analyzed by Lee et al. (1995) have shoshonitic compositions which also differ by lower 143 Ndj 144Nd isotope ratios (Arculus et al., 1995); this group is excluded from the following discussion]. This KzO variation has been interpreted to reflect significant long- and short-term variations in the mantle sources of the Mariana arc. However, the results of this study indicate that considerable KzO variation occurs within a time span of ;:::;< 1 Ma. This suggests that K z0 variation is not only related to temporal change, but also reflects a compositional diversity which is inherent to the modem Mariana arc and might have been likewise during the arc's evolution. The magmatic diversity of the CIP is shown by the large range of incompatible element ratios in the 'primitive' arc magmas (Table 6, Fig. 10). These ratios do not converge at basalts with > 6.5 wt. % MgO which show almost the same range as the bulk of the 'primitive' arc lavas with MgO contents > 4 wt. % (Table 6). In the Aleutians, early precipitation of amphibole was suggested to cause significant fractionation of the LREEjHREE in calc-alkaline lavas, increasing (CejSm)n from about 1.5 to 2.7 with increasing silica (e.g., Kay and Kay, 1994; Miller, 1995). Similar to the Aleutians, Mariana arc volcanos have tholeiitic and calc-alkaline compositions (e.g., Meijer, 1982; Woodhead, 1988). At crustal thicknesses of 20-25 km (z 7.5 kbar) and the increased melt water content of the Mariana arc lavas (Garcia et al., 1979; Baker, 1987), amphibole could be stable in lower crustal magma chambers. Calc-alkaline CIP lavas have higher average (CejSm)" ratios (= 1.4 ± 0.3; range = 1.2-2.0; n = 10) than tholeiitic compositions (= 1.1 ± 0.2; range = 0.6-1.5; n = 49) (Fig. lOa). However, model

269

calculations (using Eq. (A.l) and Eq. (A.2) in Appendix A) show that even with 25% amphibole in the cumulate, the required (CejSm)" fractionation in basalts with > 6.5 MgO wt. % could only be achieved at a minimum degree of crystallization of ;:::;< 50%. This is inconsistent with the limited range of SiOz (=48.2-53.5 wt.%) and MgO (=6.9-10.7 wt.%). Unlike the Aleutians, amphibole is never observed early in the crystallization sequence, but only as a rare late stage phase in andesites from Sarigan, Cheref and Makhanas volcanos (Meijer and Reagan, 1981; Bloomer et al., 1989b). Moreover, both intercenter and intracenter variations in incompatible element ratios other than REE appear to be independent from the criteria that define tholeiitic and calc-alkaline series (Fig. 10). For example, note the lower average range in BajLa of Agrigan compared to other tholeiitic volcanos Pagan, Guguan and Ascuncion, as well as the large range in BajLa of transitional Ahyi lavas, that covers almost the entire range observed. Incompatible element pairs with similar mantle incompatibilities, such as BajRb, KjLa and ZrjSm are also fractionated, though the intercenter differences are less pronouced in comparison to the intracenter range (e.g., Fig. lOc). None of the incompatible element ratios appears to be related to the volcanos' relative position along the arc front nor to the volcanos' volume or height, implying that larger volumes of central volcanos Pagan and Agrigan simply reflect their longer activity, as suggested by Bloomer et al. (1989a). It seems that the fractionation of incompatible element pairs does not originate in the crust, but appears to have been inherited from the CIP mantle sources. The limited isotopic variation and the flat HREE pattern [(Dy jYb)n = 1.12 ± 0.06 (n = 30) for

Fig. 11. Compositonal variation within the ClP segment of the Mariana arc exemplified by (a) (CejSm)n (HFSEjHFSE), (b) BajLa (LILEjHFSE) and (c) BajRb (LILEjLILE) ratios, compared to ratios of HKS and LKDR tephra (left colunm). Shaded bands denote mean (±2u) of MTBI5-18°N compositions. Circles represent . primitive' arc lavas « 57 wt.% Si0 2 , > 4 wt.% MgO); triangles the more evolved compositions. Filled and open symbols denote tholeiitic and calc-alkaline compositions (according to the classification given by Miyashiro (1974). Half-filled symbols indicate compositions for which a tholeiitic or calc-alkaline trend cannot be assigned (An = Anatallan; S = Sarigan; Z = Zealandia; G = Guguan; Al = Alamagan; P = Pagan; Ag = Agrigan; Py = Poyo; C = Cheref; As = Ascuncion; M = Maug; SR = Supply Reef; Ay = Ahyi; Mk = Makhallanas; U = Uracas; NU = Northwest Uracas). Stippled line in upper diagram gives relative volcano volumes (not to scale) based on Bloomer et ai. (1989a). Stippled line in lower diagram indicates volcanic peaks relative to sealevel [data from Meijer (1982) and Bloomer et al. (l989a)]. Makhallanas volcano, a seamount on the backarc side of Uracas has comparatively high (CejSm)n ('" 2) ratios and low BajLa ratios ('" 20), a feature typical for Marianas cross-chain volcanoes (e.g., Stem et aI., 1993).

270

S.M. Straub / Journal of Volcanology and Geothennal Research 76 (1997) 251-276

'primitive' arc lavas] require the CIP magmas to form in an isotopically homogeneous spinel-bearing mantle with only minor contributions ( < 1%) from the subducted plate. Building on the co-variance of K 2 0 and Zr with the REE slopes and the offset enrichment trends of KzO and Zr between different CIP volcanoe, Meijer and Reagan (1983) suggested variable degrees of melting (about the factor of 2 at 10% partial melting) in the sub arc mantle as the major cause for the magmatic diversity of the CIP segment. The extended data base of this study and in particular the larger suite of trace-element data preclude a variation in degree of partial melting as the

only cause of magmatic diversity. Using the batch melting equation (Appendix B; Shaw, 1970) it can be demonstrated, that the range jin the ratios of incompatible elements in basalts with > 6.5 wt. % MgO cannot be generated by mellting a homogeneous source. Although the physical model of batch melting is implausible (accumulated fractional melting is a more likely scenario; see for example Hochstaedter et al., 1996, for discussion), at the high degree of melting inferred for the arc front (5-20%; Stolper and Newman, 1994), the chemical consequences of batch melting and accumulated fractional melting are indistinguishable (Williams and Gill,

Table 6 Incompatible element and isotope ratios for Mariana arc lavas from the Central Island Province Incompatible element ratios Central Mariana arc (CIP)

Calculated ratios (2)

'primitive arc lavas' (Si0 2 < 56 wt.%; MgO > 4 wt.%) MgO> 6.5 wt.% (I) min (La/Sm)n Sm/Y Zr/Sm Ba/Rb K/Rb Ba/La Ba/Sm K/Sm Rb/Sm

0.2 0.10 16 8 341 19 30 1262 1.9

max 2 0.18 31 28 747 61 121 2752 7.49

average

1.2 0.13 23 5.37 496 39 74 2001 4.23

stdev

n

0.4 0.02 3 18.44 91 11 20 322 1.17

38 63 63 41 46 46 45

stdev

n

0.00009 0.00003 0.09 0.02 0.09 0.3

113 46 41 41 41 48

43

min 0.9 0.13 16

11 386 22 30 1262 1.9

max 1.5 0.18 27 27 664 46 89 2184 5.65

NMORBS

NMORBS

NMORBS

5

0.1 7.8

0.55 20 b

n

5 5 7

11 11 5 7 7 7

b

1.0 0.31 26 10.8 923 3 4.3 369 0.40

a b

l.l 0.26 26 9.8 612 6 10 626 1.02

a

1.3 0.17 28 9.4 512 10 22 1161 2.27

Isotope ratios (for all compositions) (3) Central Mariana arc (CIP) min Sr/ 86 Sr l43Nd/144Nd 206Pb/204Pb 207 Pb / 204 Pb ~~8Pb/ 204Pb

al8 0

0.70319 0.513 18.52 15.51 38.05 5.5

max 0.70376 0.513 18.9 15.6 38.49 6.6

average 0.70344 0.513 18.75 15.55 38.33 6.1

(I) CIP lavas with> 6.5 wt.% MgO are from Ahyi, Poyo, Pagan, Sarigan and Zealandia volcanos. (2) Calculated ratios are ratios of partial melts (Eq. (A.3» derived from NMORB source (NMORBS) of Stolper and Newman (1994) or of mixture of NMORBS and the H 20-rich component of Stolper and Newman (1994) with the amount of H 20-rich component being 0.1 and 0.55 wt.%, respectively. The degree of melting and amount of H 20-rich component are related as in fig. 6 of Stolper and Newman (1994). (3) Data for CIP isotope ratios are from Meijer (1976), White and Patchett (1984), Ito and Stem (1980), Stem (1979), Dixon and Stem (1983), Woodhead and Fraser (1985), Woodhead (1989) and Lin et al. (1990). For other data sources see text. a Amount in wt.% of H 20-rich component of Stolper and Newman (1994) in source. b Degree partial melting (%).

S.M. Straub / Journal of Volcanology and Geothennal Research 76 (J997) 251-276

1989). The NMORB source composition which Stolper and Newman (1994) inferred from Mariana Trough basalts is also a possible arc front source (Stolper and Newman, 1994). It has an (La/Sm)" = 0.7 which falls at the lower end of (La/Sm)" of the 'primitive' arc lavas CIP lavas [only two 'primitive arc magmas' out of 43 have (La/Sm\ ratios < 0.7 indicating the existence of even more depleted source material]' At a degree of melting 5-20% of the NMORB sources, (La/Sm)n in the daughter product increases to z 0.8-1.0. A further increase in (La/Sm)" up to 2.0 as observed in the 'primitive' arc lavas requires a change in the composition of the Mariana arc source. Following the model of Stolper and Newman (1994) this could be achieved by adding the slab-derived HzO-rich component to the arc source. By adding 0.1 to 0.55 wt.% of the HzO-rich component (table 2 in Stolper and Newman, 1994) to the NMORB source, however, the (La/Sm)n only increases to 1.3 in the endproducts, with the amount of HzO-rich component and degree of melting being related as in Fig. 6 of Stolper and Newman (1994). Thus, simple binary mixing between these two endmembers is not able to explain the total (La/Sm)" range observed. The NMORB source has LILE/LILE and LILE/HFSE ratios which lie outside the range defined by the 'primitive' arc magmas (Table 6). Adding the HzO-rich component of Stolper and Newman (1994) to the NMORB source shifts only the LILE /LILE ratios of the partial melts into the field of the 'primitive' arc magmas, whereas the LILE/HFSE ratios are either too low or are at the lower end of the range of the 'primitive' arc magmas (Table 6). Also, adding a HzO-rich component with a constant composition does not fractionate the incompatible element ratios in the partial melt in the range observed. In order to match both the ratios and the range observed in the 'primitive' arc lavas, the source compositions need to be changed. Following the Stolper and Newman (1994) model and having in mind the intercenter and intracenter variability of these ratios along the arc front, it appears that changes in the source compositions are caused by a variable HzO-rich component, and not by significant compositional differences of the NMORB source endmember. In order to match the range in incompatible element ratios of the 'primitive' arc lavas, the HzO-

271

rich component infused in the CIP sources had to be enriched in the elements Ba, Rb, K by up to the factor of z 5, z 3 and z 2 relative to Sm, compared to the HzO-rich component of Stolper and Newman (1994) that reaches the backarc sources. According to Stolper and Newman (1994), the following factors can change the composition of the slab-derived component in the required way: (I) chromatographic effects in the subarc mantle column, which act to enrich LILE; (2) a higher fluid flux beneath the arc front source, that is coupled with a higher amount of recycled LILE; and (3) small variations in the composition of the subducted sediment transferred to the subarc mantle. Minor variability in the sediment component is supported by Pb-isotope data of Woodhead and Fraser (1985) who noted distinctly high Z06 Pb / Z04 Pb ratios in calc-alkaline Sarigan lavas, which might be attributable to a slighter larger sediment component in the Sarigan source. This could also be the cause for the variable Ba/HFSE ratios and fractionation of the Ba/LILE ratios (e.g., Ba/Rb ratios in Fig. lOc) making these ratios most variable along the arc front. Initial variation in Ba might also explain the excess Ba concentration observed in the LKDR glasses, if HKS glasses were derived from a volcano with low Ba/La and higher (Ce/Sm)n (e.g., Agrigan) and LKDR glasses from a source with high Ba/La and lower (Ce/Sm)n (e.g., Guguan or Supply Reef). This conclusion would be in agreement with a recent model of Elliott et aI. (1996) in which the CIP's diversity in incompatible, fluid-mobile elements is interpreted as a consequence of the compositional variability of the recycled, slab-derived component. The amount of fluid fluxing the mantle sources has been linked to the degree of partial melting (e.g., Fryer et aI., 1990; Miller et aI., 1992). If the HzO content in the H 2 0-rich component also varies as do the other elements, this should cause differences in the degree of partial melting of the CIP subarc mantle (Stolper and Newman, 1994). In addition to the variable content of slab-derived incompatible elements this could, contribute further to the compositional diversity observed in the CIP lavas. Important tracer elements, such as Cs /Nb and B /Nb, however, which could identify variable amounts of volatile transfer into arc sources are presently not available in sufficient amounts for the CIP.

272

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (]997) 251-276

7. Conclusions (I) Tephra in sediments of central Mariana Trough contain three different glass shard series, which are distinguished by major and trace elements, colors, morphologies and dispersal patterns. (2) The three glass shard series are derived from different volcanic sources: the subordinate LKB series is hydroclastic debris from Mariana Trough lavas, whereas HKS and LKDR series are both derived from explosive eruptions of volcanos of the central Mariana arc. Only LKB glasses match their sources lava compositions; arc-derived glasses are evolved derivative liquids which differ compositionally from source area lavas. (3) The variability of the arc-derived HKS and LKDR glass series reflects the magmatic diversity of the coeval Mariana source lavas. Using the quantitative approach of Stolper and Newman (1994), the arc diversity is suggested to reflect the metasomatism of arc sources by a compositionally variable, slab-derived H 2 0-rich component, possibly compounded by small differences in the degree of partial melting. Thus, the tephra variability, expressed by variation in KzO content up to a factor of 3 at 60 wt.% Si0 2 , reflects the compositional diversity inherent to the modem Mariana arc and not necessarily temporal variations in the composition of the sources.

Acknowledgements

C.W. Devey, D. Matthies, H.U. Schmincke and 1. Thiede contributed in various, but invaluable ways to this work. Tim Elliott and Terry Plank are thanked for allowing to use yet unpublished data. ICP-MS analyses were performed by C.D. Garbe-SchOnberg at the Geologisches Institut, Universitlit Kiel (GIK). Scanning electron microsopy was carried out at the SEM laboratory of C. Samtleben and U. Schuldt (GIK). 1. Scholten (GIK) determined the sedimentations rates of the core 46 KL by 230 Th methods. D. Ackermand and B. Mader (Mineralogisches Institut, Universitat Kiel) are thanked for assistance with microprobe analyses. Laboratory work was assisted by W. Rehder and F. Werner (both GIK). Financial support trough the DFG (German Science Founda-

Table 7 Results from major-element mass balance calculation given in wt.% Parent: Daughter:

GUN 61379

SUR13-22 61379

GUN D272

SUR13-22 D272

SAP 57269

R F OPX CPX PLAG TMG AP AMPH

0.465 0.457 11.53 8.69 66.56 12.92 0.31 (4)

0.393 0.344 11.76 6.97 70.93

0.201 0.483 11.66 7.18 69.32 11.56 0.60

0.24 0.332 11.99 6.86

0.027 0.727 15.067 20.337 55.796 8.801

1O.Q7

0.30 (5)

n.93 7.96 0.70

Na omitted from all calculations. R = sum of squared residuals. F = fraction of bquid remaining. Weight percentage for amphibole (observed in sample 61379) given in bracket was not derived from mass balance calculation, but introduced arbitrarily into the trace-element models in order to achieve a better fit for the REE. Mineral compositions used are phenocrysts from low-K ('" 1.3 K 2 0 at 70 wt.% Si0 2 ) dacitic to rhyolitic tephra layer 125-782A-2H-4-1l3-1l4 from the Izu-Bonin forearc basin (Straub and Schmincke, unpublished data).

tion) and the BMBF (German Federal Ministry for Research and Technology) is gratefully acknowledged. Constructive reviews by R.J. Arculus and R.J. Stem improved the manuscript as well as A. LaGatta's proof-reading and C.H. Langmuir's comments.

Appendix A

Calculations for fractional crystallization models Major-element trends were tested by mass balance using least-square methods of calculation. The results are tabulated in Table 7. The results were used for testing trace-element variations assuming either (a) perfect fractional crystallization:

c =c L

·F(D-l) 0

(A.I)

where c L = element concentration in the derivative melt, Co = element concentration in the parent melt, F = residual melt fraction and D = bulk distribution coefficient given by D = ED. Xa with D. = melt/mineral partition coefficient and X. = amount

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

of mineral fractionating, or (b) in-situ crystallization (Langmuir, 1989):

c

L

= C . F[f·(E- l)/(j-I)]

(A.2)

0

where F = residual melt fraction in the chamber interior after formation of a marginal crystal mush from an initially homogenous magma body, f = melt fraction returned from marginal mush into the chamber interior (an arbitrary f = 0.5 was chosen), cL trace-element concentration in the chamber after mixing of re-injected liquid and liquid in the interior, E = cfl c L where cf = element concentration of residual melt in the mush. Assuming constant partition coefficients, E can be approximated as E = [D ·(1-f)+f]-I. The choice of potential parentallow-K andesites with LKDR-similar (Ce/Sm)n ratios was limited by the availability of REE analyses on these samples. Arc andesites GUN [(Ce/Sm)n = 0.94 from Guguan island (Woodhead, 1989)] and SUR13-22 (1.08) from the CIP seamount Supply Reef (Bloomer et al., 1989b; Lin et aI., 1989) were the best candidates for glass samples 61379 (1.31) and D272 (1.02) glass compositions. Andesite SAP (1.35) from Sarigan island (Woodhead, 1989) was the best available parental composition for glass sample 57269 (1.20) (HREE Dy, Er and Yb were omitted for parental compositions GUN and SAP). Model calculations of the trace elements are confined to the REE, as mineral compositions and melt/mineral partition coefficients (Table 8) were taken from the literature. This implies rather large uncertainities, in addition to the fact that none of the compositions used are true

273

parent-daughter liquids. Thus, instead of rigorously testing individual element variation, I tested whether the observed elemental variations in general are consistent with the fractionation model proposed or not.

Appendix B

Model calculations for partial melting.

Incompatible element variations during partial melting were tested using the batch melting equation of Shaw (1970): CL =

co' 1/( D

+F -

(A.3)

FD)

where c L = element concentration in the daughter melt, Co = element concentration in the original solid, F = melt fraction formed, and D = bulk distribution coefficient given by D = '[,Da Xa with Da = melt/mineral partition coefficient and Xa = weight fraction of phase a in the source. Modal composition of mantle source used is OLlV:OPX:CPX:SP = 54:25:20:1. Partition coefficients are from Ewart and Hawkesworth (1987). The primary goal of the modelling was to evaluate the changes in the ratios of incompatible elements in dependence from degree of melting and source composition. The absolute enrichments of incompatible elements in the calculated daughter melts, however, are in general consistent with the those observed in the most primitive CIP lavas with > 6.5 wt. % MgO.

Table 8 Partition coefficients used in fractional crystallization models compiled from Henderson (1982). Sisson (1994) (amphIbole) and Nagasawa (1970) (apatite)

CPX OPX PLAG TIMG AMPH AP a b

La

Ce

Nd

Sm

0.085 0.D2 b 0.140 0.530 0.861 34.73

0.340 0.020 0.140 0.610 2.130 34.73

0.600 0.050 0.080 0.880 5.193 57.13

0.900 0.050 0.080 0.930 7.764 62.80

Linear interpolated. Assumed.

Eu 0.900 0.080 0.733 0.580 8.570

30045

Gd

Dy

Er

0.850 0.04 a 0.100 0.9 a 9.380 56.76 a

1.100 0.020 0.090 0.800 10.186 50.73

1.000 0.310 0.080 0.8 a 6.857 37.18

Yb

1.000 0.340 0.070

00400 40410 23.85

Source Henderson (1982) Henderson (1982) Henderson (1982) Henderson (1982) Sisson (1994) Nagasawa (1970)

274

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276

References Arculus, R.I., Gill, J.B., Cambray, H., Chen, W. and Stem, R.I., 1995. Geochemical evolution of arc systems in the Western PacifIc: the ash and turbidite record recovered by drilling. In: B. Taylor and J. Natland (Editors), Active Margins and Marginal Basins of the Western Pacific. Am. Geophys. Union, 88: 45-65. Baker, D.R., 1987. Depths and water content of magma chambers in the Aleutian and Mariana island arcs. Geology, 15: 496-499. Banks, N.G., Koyanagi, R.Y., Sinton, lM. and Honma, K.T., 1984. The eruptIOn of Mount Pagan volcano, Mariana islands, 15 May 1981. J. Volcanol. Geotherm. Res., 22: 225-269. Beard, J.S., 1995. Experimental, geological and geochemical constraints on the origins of low-K silicic magmas in oceanic arcs. J. Geophys. Res., 100(B8): 15,593-15,600. Bednarz, U. and Schmincke, H.U., 1994. Composition and origin of volcaniclastic sediments in the Lau Basin (SW -Pacific), ODP Leg 135 (Sites 834-893). Proc. ODP Sci Results, 135: 51-74. Bloomer, S.H., Stem, R.I. and Smoot, N.C., 1989a. Physical volcanology of the submarine Mariana and Volcano arcs. Bull. Volcanol., 51: 210-224. Bloomer, S.H., Stem, R.I., Fisk, E. and Geschwind, C.H., 1989b. Shoshonitic volcanism in the northern Mariana arc. J. Geophys. Res., 94: 4469-4496. Burnham, C.W., 1983. Deep submarine pyroclastic eruptions. In: H. Ohmoto and B.I. Skinner (Editors), The Kuroko and Related Massive Sulfide Deposit: Introduction and Summary of New Fmdings. Econ. Geol. Mon., 5: 142-148. Cashman, K.V. and Fiske, R.S., 1993. Fallout of pyroclastic debris from submarine volcanic eruptions. Science, 253: 275280. Clift, P.D. and Dixon, J.E., 1994. Variations in arc volcanism and sedimentation related to rifting of the Lau Basin (Southwest Pacific). Proc. ODP Sci. Results, 135: 23-45. Dixon, T.H. and Batiza, R., 1979. Petrology and chemistry of recent lavas in the northern Marianas: implications for the origin of island arc basalts. Contrib. Mineral. Petrol., 70: 167-181. Dixon, T.H. and Stem, R.J., 1983. Petrology, geochemistry and isotopic composition of submarine volcanoes in the southern Mariana arc. Geol. Soc. Am. Bull., 94: 1159-1172. Elliott, T., Plank, T., Zindler, A., White, W. and Bourdon, B., 1996. Element transport from subducted slab to juvenile crust at the Mariana arc. l Geophys. Res. (in review). Ewart, A. and Hawkesworth, C.I., 1987. The Pleistocene-Recent Tonga-Kermadec arc lavas: interpretation of new isotopic and rare earth data in terms of a depleted mantle source model. J. Petrol., 28: 495-530. Fisher, R.V. and Schmincke, H.U., 1984. Pyroclastic Rocks. Springer, Heidelberg, 472 pp. Fryer, P., Sinton, lM. and Philpotts, lA., 1981. Basaltic glasses from the Mariana Trough. Init. Rep. DSDP, 60: 71-74. Fryer, P., Taylor, B., Langmuir, C.H. and Hochstaedter, A.G., 1990. Petrology and geochemistry of lavas from the Sumisu

and Torshima backarc rifts. Earth Planet. Sci. Lett., 100: 161-178. Garbe-Schonberg, C.D., 1993. Simultaneous detenmination of thirty-seven trace elements in twenty-eight international rock standards by ICP-MS. Geostandards Newsletter, 17(1): 81-97. Garcia, M.O., Liu, N.w.K. and Muenow, D.W., 1979. Volatiles in submarine volcanic rocks from Mariana island arc and trough. Geochim. Cosmochim. Acta, 43: 305-312. Gribble, R.F., 1996. Variation in incompatible trace element and isotopic compositions in Mariana Trough lavas. Doct. Thesis, Univ. Texas, Dallas, TX (unpubl.). Gribble, R.F., Stem, R.I., Bloomer, S.H., Stiiben, D., O'Hearn, T. and Newman, S., 1996. MORB mantle and subduction components interact to generate basalts in the Southern Mariana Trough back-arc basin. Geochim. Cosmochim. Acta, 60(12): 2153-2166. Halbach, P., Segl, M., Puteanus, D. and Mangini, A., 1983a. Co-fluxes and growth rates in ferromanganese deposits from central PacifiC seamount areas. Nature, 304: 716-719. Hart, S.R., Glassley, W.E. and Karig, D.E., 1972. Basalts and sea floor spreading behind the Mariana island arc. Earth Planet. Sci. Lett., 15: 12-18. Hawkins, lW., Lonsdale, J.D., MacdougaJll, J.D. and Volpe, A.M., 1990. Petrology of the axial ndge of the Mariana trough backarc spreading center. Earth Planet. Sci. Lett., 100: 226250. Henderson, P., 1982. Inorganic Geochemisnry. Pergamon Press, 353 pp. Hochstaedter, A.G., Kepezhinskas, P., Defant., M., Drummond, M. and Koloskov, A., 1996. Insights into the volcanic arc mantle wedge from magnesian lavas from the Kamchatka arc. J. Geophys. Res., 101(Bl): 697-712. Hole, M.J., Saunders, A.D., Marriner, G.F. and Tarney, l, 1984. Subduction of pelagic sediments: implications for the origin of Ce-anomalous basalts from the Mariana islands. J. Geol. Soc. London, 141: 453-472. Ishikawa, T. and Nakamura, E., 1994. Origin of the slab component in arc lavas from across-arc variation of B and Pb isotopes. Nature, 370: 205-208. Ito, E. and Stem, R.J., 1980. Oxygen and strontium isotopic investigations on the origin of volcanism in the Izu-volcanoMariana island arc. Carnegie Inst. Washington Yearb., 80: 449-455. Jackson, M.C., 1989. Petrology and petrogenesis of recent submarine volcanics from the northern the Mariana arc and backarc basin. Doct. Thesis, Univ. Hawaii, 277 pp. (unpubl.). Jackson, M.C., 1993. Crystal accumulation and magma m~xing in the petrogenesis to tholeiitic andesites from Fukujin Seamount, Northern Mariana Island Arc. J. Petrol., 34(2): 259-289. Juster, T.C. and Grove, T.L., 1989. Experimental constraints on the generation of FeTi basalts, andesites, and rhyodacites at the Galapagos spreading center, 85°W and 95°W. J. Geophys. Res., 94: 9251-9274. Karig, D.E., 1971. Structural history of the Mariana island arc system. Geol. Soc. Am. Bull., 82: 323-344. Kay, S.M. and Kay, R.W., 1994. Aleutians magmas in space and

S.M. Straub / Journal of Volcanology and Geothermal Research 76 (1997) 251-276 time. In: G. Plafker and H.C. Berg (Editors), The Geology of Alaska. GeoL Soc. Am., Boulder, CO. pp. 687-722. Langmuir, C.H., 1989. Geochemical consequences of in-situ crystallisation. Nature, 340: 199-205. Larson, E.E., Reynolds, R.L., Merrill, R., Levi, S., Ozima, M., Aoki, Y., Kinoshita, H., Zasshu, S., Kawai, N., Nakajima, T. and Hirooka, K., 1974. Major-element petrochemistry of some extrusive rocks from the volcanically active Mariana islands. BulL VolcanoL, 38: 361-377. Lee, J.M., Stem, R.J. and Bloomer, S.H., 1995. Forty million years of magmatic evolution in the Mariana arc: the tephra record. J. Geophys. Res., 100(B9): 17,671-17,687. LeMaitre, R.W. (Editor), 1989. A Classification of Igneous Rocks. Blackwell, Oxford, 193 pp. Lin, P.N., Stem, R.J. and Bloomer, S.H" 1989. Shoshonitic volcanism in the northern Mariana arc 2. Large-ion lithophile and rare earth element abundances: evidence for the source of incompatible element enrichments in intraoceanic arcs. J. Geophys. Res., 94: 4497-4514. Lin, P.N., Stem, R.J., Morris, J. and Bloomer. S.H., 1990. Ndand Sr-Isotopic compositions of lavas from the northern Manana and southern Volcano arcs: implications for the origin of island arc melts. Contrib. MmeraL PetroL, 105: 381-392. Lonsdale, P. and Hawkins, J., 1985. Silicic volcanism at an off-axis geothermal field m the Mariana Trough back-arc basin. GeoL Soc. Am. BulL, 96: 940-951. Macdonald, R., Sparks, R.S.J., Sigurdsson, H .. Mattey, D.P., McGravie, D.W. and Smith, R.L., 1987. The 1875 eruption of Askja volcano, Iceland: combined fractional crystallization and selective contamination in the generation of rhyolitic magma. Mineral. Mag., 51(360): 183-202. Meijer, A., 1976. Pb and Sr isotopic data bearing on the origin of volcanic rocks from the Mariana island-arc system. GeoL Soc. Am. Bull., 87: 1358-1369. Meijer, A., 1980. Primitive arc volcanism and a boninite series: examples from western Pacific island arcs. Am. Geophys. Vnion, 23: 269-282. Meijer, A., 1982. Mariana - Volcano Islands. In: R. Thorpe (Editors), Andesites. Wiley, Chichester, pp. 293-306. Meijer, A. and Reagan, M., 1981. Petrology and geochemistry of the island of Sarigan in the Mariana arc; calc-alkaline volcanism in an oceanic setting. Contrib. MineraL PetroL, 77: 337354. Meijer, A. and Reagan, M., 1983. Origin of KzO-SiO z trends in volcanoes of the Mariana arc. Geology, 11: 67-71. Meijer, A., Anthony, E. and Reagan, M., 1981. Petrology of volcanic Rocks from the fore-arc sites. Init. Rep. DSDP, 60: 709-730. Miller, D.M., 1995. PetrogenesIs of adjacent calc-alkaline and tholeiitic volcanoes in Vnmak Island, Aleutian Islands, Alaska. Doct. Thesis, Columbia Vniv., 455 pp. (unpubL). Miller, D.M., Langmuir, C.H., Goldstein, S.L. and Franks, A.L., 1992. The importance of parental magma compOSItion to calc-alkaline and tholeitic evolution: evidence from Vmmak Island in the Aleutians. J. Geophys. Res., 97(Bl): 321-343. Miyashiro, A., 1974. Volcanic rock series in island arcs and active continental margins. Am. J. Sci., 274: 321-355.

275

Nagasawa, H., 1970. Rare earth concentrations in zircons and apatites and their host dacites and granites. Earth Planet. Sci. Lett., 9: 359-364. Packham, G.H. and Williams, K.L., 1981. Volcanic glasses from sediments from sites 453 and 454 in the Mariana Trough. Init. Rep. DSDP, 60: 483-497. Pearce, J.A., 1983. Role of the sub-continental lithosphere in magma genesis at actIve continental margins. In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths. Birkhauser, Boston, Nantwich Shiva, Cambndge, pp. 230-249. Piischel. V., 1992. Petrographische und geochemische Charakterisierung von magmatischen Differentiaten entlang des 'Backarc' -Riftsystems im siidlichen Marianen Trog. M.Sci. Thesis, Vniv. Kiel, 90 pp. (unpubL). Reagan, M.K. and Meijer, A., 1984. Geology and geochemistry of early arc-volcanic rocks from Guam. GeoL Soc. Am. BulL, 95: 701-713. Ryan, J.G., Morris, J., Tera, F., Leeman, W.P. and Tsvetkov, A" 1995. Cross-arc geochemical variations in the Kurile Arc as a function of slab depth. Science, 270: 625-627. Sager, W.W., 1980. Mariana arc structure inferred from gravity and seismic data. J. Geophys. Res., 85: 5382-5388. Schmidt, R.G., 1957. Petrology of the volcanic rocks. In: P.E. Cloud et aL (Editors), Geology of Saipan Mariana Islands. V.S. GeoL Surv., Prof. Pap., Washington, DC, Part 2, Ch. B. pp. 127-172. Schmincke, H.V., 1981. Ash from vitric muds in deep dea cores from the Manana trough and forearc region (South Philippine Sea), (sites 453, 454,455, 458, 459 and 460) Deep Sea Drilling Project Leg 60. Init. Rep. DSDP, 60: 473-482. Shaw, D.M., 1970. Trace element fractionation during anatexis. Geochim. Cosmochim. Acta, 34: 237-243. Silver, L.A., Ihinger, P.D. and Stolper, E., 1990. The influence on bulk composition on the speciation of water in silicate glasses. Contrib. MineraL PetroL, 104: 142-162. Sisson, T.W., 1994. Hornblende-melt trace-element partitioning measured by ion microprobe. Chem. GeoL, 117: 331-344. Stem, R.J., 1979. On the origin of andesite in the northern Mariana island arc: implications from Agrigan. Contrib. Mineral. PetroL, 68: 207-219. Stem, R.J. and Bibee, L.D., 1984. Esmeralda Bank: geochemistry of an active submarine volcano m the Mariana island arc. Contrib. MineraL PetroL, 86: 159-169. Stem, R.J., Bloomer, S.H., Lin, P.N., Ito, E. and Morris, J., 1988. Shoshonitic magmas in nascent arcs: new evidence from submarine volcanoes in the northern Marianas. Geology, 16: 426-430. Stem, R.J., Bloomer, S.H .. Lin, P.N. and Smoot, N.C., 1989. Submarine arc volcanism in the southern Mariana arc as an ophiolite analogue. Tectonophysics, 168: 151-170. Stem, R.J., Lin, P.N., Morris, J.D., Jackson, M.C., Fryer, P., Bloomer, S.H. and Ito, E., 1990. Enriched back-arc basin basalts from the northern Mariana Trough: implications for the magmatic evolution of back-arc basins. Earth Planet. SCI. Lett., 100: 210-225. Stem, R.J., Jackson, M.C., Fryer, P. and Ito, E., 1993. 0, Sr. Nd

276

S.M. Straub / Journal a/Volcanology and Geothennal Research 76 (1997) 251-276

and Pb Isotopic compositions of the Kasuga Cross-Chain in the Mariana Arc: A new perspective on the K-h-relationship. Earth Planet. Sci. Lett., 119: 459-475. Stolper, E. and Newman, S., 1994. The role of water in the petrogenesis of Mariana trough magmas. Earth Planet. Sci. Lett., 121: 293-325. Straub, S.M., 1995. Contrasting compositions of Mariana Trough fallout tephra and Mariana island arc volcanics: a fractional crystallization link. Bull. Volcano!.. 57(6): 403-421. Sun, S.S. and McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle compoSItion and processes. In: A.D. Saunders and MJ. Norry (Editors), Magmatism in the Ocean Basins. Geol. Soc. London, Spec Pap., Blackwell, Oxford, pp. 313-345. Taylor, B., 1992. Rifting and the volcanic-tectonic evolution of the Izu-Bonin-Mariana Arc. Proc. ODP Sci. Results, 126: 627-651. Volpe, A.M., Macdougall, J.D. and Hawkins, J.W., 1987. Mariana trough basalts (MTB): trace element and Sr-Nd isotopic evidence for mixing between MORB-like and arc-like melts. Earth Planet. Sci. Lett., 82: 241-254. Volpe, A.M., Macdougall, J.D., Lugmair, G.w., Hawkins, J.W. and Lonsdale, P., 1990. Fine-scale isotopic variation in Marrana trough basalts: evidence for heterogeneity and a recycled component in backarc baSIn mantle. Earth Planet. Sci. Lett., 100: 251-264. Wallace, PJ. and Gerlach, T.M., 1994. Magmatic vapor source for sulfur dioxide released during volcanic eruptions: evidence from Mount Pinatubo. SCIence, 265: 497-499.

White, W.M. and Patchett, l, 1984. Hf-Nd-Sr isotopes and incompatible element abundances in island arcs: implications for magma origins and crust-mantle evolutIOn. Earth Planet. Sci. Lett., 67: 167-185. Williams, R.W. and Gill, lB., 1989. Effects of partial melting on the uranium decay series. Geochim. Cosmochim. Acta, 53: 1607-1619. Wolff, G., 1990. Geochemische, petrograprusche und petrologische Untersuchungen an Vulkaniten des Marianen Beckens. M.Sci. Thesis, Univ. Kiel, 87 pp. (unpubl.). Wood, D.A., Marsh, N.G., Tarney, l, Joron, J.L., Fryer, P. and Treuil, M., 1981. Geochemistry of igneous rocks recovered from a transect across the Mariana trough, arc, fore-arc, and trench, sites 453 through 461, Deep Sea Drilling Project Leg 60. Init. Rep. DSDP, 60: 611-646. Woodhead, lD., 1988. The origin of geochemical variations in Mariana lavas: a general model for petrogenesis in intra-oceanic island arcs? l Petrol., 29: 805-830. Woodhead, lD., 1989. Geochemistry of the Mariana arc (western Pacific): Source composition and processes. Chem. Geol., 76: 1-24. Woodhead, J.D. and Fraser, D.G., 1985. Pb, Sr and lOBe isotopic studies of volcanic rocks from the northern Mariana islands. Implications for magma genesis and crustal recycling in the western Pacific. Geochim. Cosmochim. Acta, 49: 1925-1930. Woodhead, J.D., Harmon, R.S. and Fraser, D.G., 1987. 0, S, Sr and Pb isotope varratIOns in volcanic rocks from the northern Mariana islands: implications for crustal recycling in intra-oceanic arcs. Earth Planet. Sci. Lett., 83: 39-52.