Origin of Basalts in a Hot Subduction Setting ...

JOURNAL OF PETROLOGY

VOLUME 55

NUMBER 2

PAGES 241^281

2014

doi:10.1093/petrology/egt064

Origin of Basalts in a Hot Subduction Setting: Petrological and Geochemical Insights from Mt. Baker, Northern Cascade Arc EMILY K. MULLEN* AND I. STEWART McCALLUM DEPARTMENT OF EARTH AND SPACE SCIENCES, UNIVERSITY OF WASHINGTON, BOX 351310, SEATTLE, WA 98195, USA

The northern Cascade arc is an end-member ‘hot’ subduction zone where slab dehydration may be essentially complete prior to reaching sub-arc depths, presenting a potential problem for a flux melting origin for arc basalts. Nevertheless, mafic lavas from the Mt. Baker volcanic field, the most magmatically productive volcano in the northern Cascade arc, record subduction input and compositions typical of calc-alkaline magmas. Relative to normal mid-ocean ridge basalt, the most primitive lavas at Mt. Baker show elevated abundances of large ion lithophile elements (LILE), Pb and light rare earth elements (LREE). High field strength elements (HFSE) and heavy REE (HREE) are depleted relative to LILE. Reconstructed primary magmas define three groups: Group I calcalkaline basalts (Sulphur Creek, Lake Shannon, Hogback) have lower SiO2, (La/Yb)N and LILE/HFSE, and are olivine- and plagioclase-phyric; Group II high-Mg basaltic andesites (Tarn Plateau, Cathedral Crag) also contain clinopyroxene phenocrysts and have higher H2O, SiO2, (La/Yb)N and LILE/HFSE; Group III low-K olivine tholeiite (Park Butte) has the lowest (La/Yb)N but highest LILE/HFSE. The redox states of all groups lie between QFM (quartz^fayalite^magnetite) and QFM þ1·4. Pb, Sr and Nd isotope compositions are similar among all the analysed samples and are consistent with a depleted mantle source modified by input from the subducting slab. Trace elementîsotope modeling indicates a subduction component composed predominantly of metabasaltderived fluid with lesser amounts of sediment melt and metabasalt melt. Group I basalts record the smallest melt fractions (5^7%), lowest water contents (1·5^2·1wt %), and highest temperatures and pressures of mantle segregation (up to 13548C, 1·5 GPa) from lherzolitic ( spinel) residues. Group II basaltic andesites show the greatest extents of mantle metasomatism, the highest water contents (2·7^3·7 wt %) and partial melt fractions (10^12%), and segregated from harzburgite at 12708C, 1 GPa, consistent with

*Corresponding author. Present address: Pacific Centre for Isotopic and Geochemical Research, Department of Earth, Ocean and Atmospheric Sciences, 2020-2207 Main Mall, University of British Columbia, Vancouver, BC, V6T 1Z4, Canada. Telephone: (604) 822-3764. Fax: (604) 822-6088. E-mail: [email protected]

pooling of melts at the Moho. Group III records P^Tconditions similar to Group I (1·4 GPa, 13268C) but melt fractions (12%) and mantle residues (harzburgite) are more similar to Group II, and H2O contents (2·1wt %) are intermediate. Melting beneath Mt. Baker was initiated by dehydration melting of amphibole peridotite at 95 km and 10208C, within the stability field of garnet lherzolite. Initial melt fractions were small (1^2%) and near water-saturation. Phase equilibria and trace element modeling show no evidence for garnet-bearing mantle residues, indicating that progressive melting during ascent of diapirs through the hot core of the convecting mantle wedge reduced H2O contents and erased any residual garnet signature. Because fluid release from the slab is restricted to the forearc, mantle hydrated at shallow depths in the serpentine and/or chlorite stability fields must be down-dragged to the region of amphibole stability to initiate dehydration melting.

thermobarometry; mineral chemistry; trace elements; isotopes; Garibaldi Belt

KEY WORDS:

I N T RO D U C T I O N The compositional diversity displayed by primary arc basalts reflects variability in inputs from two major source reservoirs, the subducting slab and mantle wedge (e.g. Gill, 1981; Hawkesworth et al., 1993; Pearce & Peate, 1995). In generally accepted models for arc basalt generation, melting is initiated in the mantle wedge, at or near the interface with the subducting slab, in response to the addition of a ‘subduction component’ (SC) that consists of fluids and/or melts derived from subducting sediment, altered oceanic crust, and, in some cases, hydrated subducting mantle

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RECEIVED NOVEMBER 22, 2011; ACCEPTED OCTOBER 3, 2013


VOLUME 55

FEBRUARY 2014

GEOLOGIC A L S ET T I NG OF MOU NT BA KER Mt. Baker is an active andesitic stratocone surrounded by an extensive, multi-vent volcanic field (Fig. 2) that has been almost continuously active over the past 1·3 Myr, punctuated by several major eruptive episodes (Hildreth et al., 2003, 2004). Eruptive products consist mainly of andesite and rhyodacite, whereas less than 1% by volume is basalt. The most voluminous eruption generated 50 km3 of rhyodacite ignimbrite 1·15 Myr ago, forming the Kulshan caldera. The modern stratocone was constructed between 40 and 12 ka on the eroded flank of the older and more voluminous Black Buttes stratovolcano (Hildreth et al., 2003). The youngest lava flow is the composite Sulphur Creek flow that erupted from the Schreibers Meadow cinder cone between 8800 and 8500 14C years BP (Tucker & Scott, 2009). Numerous phreatic eruptions have occurred at the stratocone in the time since, including the AD 1843 event that formed Sherman Crater, the currently active vent; the most recent tephra known to contain a juvenile magmatic component is a basaltic andesite erupted at 5740 14C years BP (Tucker & Scott, 2006). Increased gas emissions marked a renewal of activity at Sherman Crater in 1975 but no eruption occurred (Frank et al., 1977). Since 1975 gas emissions and heat flow have declined (Werner et al., 2009; Crider et al., 2011), but deep long-period earthquakes between 1980 and 2007, centered beneath Mt. Baker, provide evidence for continued movement of magma and/or fluids at mid- to lower-crustal levels (Nichols et al., 2011). Mt. Baker is part of the northernmost segment of the Cascade arc known as the Garibaldi Volcanic Belt (GVB), which consists of seven major volcanic centers extending from Glacier Peak in the south to Silverthrone in the north (Fig. 1) (Green et al., 1988). The age of the subducting plate at the trench decreases from 10 Ma in the southern GVB (Glacier Peak and Mt. Baker) to 6 Ma at the northern end (Wilson, 2002). The GVB is the least magmatically productive segment of the Cascade arc (Hildreth, 2007). Extrapolation of slab depth contours gives a depthto-slab beneath Mt. Baker of about 85 km (McCrory et al., 2004), whereas teleseismic imaging of the slab places the slab^wedge interface near the 100 km depth contour (Bostock & VanDecar, 1995). The slab age beneath Mt. Baker is 19 Ma (Green & Harry, 1999). The total crustal thickness in the Mt. Baker region is 40 2 km (Miller et al., 1997). The upper crust is an 10 km thick stack of imbricate thrust sheets known as the Northwest Cascade Thrust System, assembled between 114 and 84 Ma (Misch, 1966; Brandon et al., 1988; Brown & Gehrels, 2007). The volcano is located in the Mt. Baker ‘window’ where the thrust sheets have been eroded to expose the footwall unit (Nooksack Formation) (Tabor et al., 2003). The mid- to lower crust is most probably the

242


lithosphere (e.g. Tatsumi & Eggins, 1995; Ulmer & Trommsdorff, 1995). The mantle wedge may melt in response to the influx of fluids or melts (flux melting), or to pressure release induced by convection (decompression melting), or both (e.g. Pearce & Parkinson, 1993; Sisson & Bronto, 1998). The nature, source, and chemical composition of the subduction component, and the quantity in which it is added to the mantle wedge, are questions that we address here for the northern Cascade arc using petrological and geochemical data from the most primitive basalts at the Mt. Baker volcanic field. The Cascade magmatic arc (Fig. 1) is an example of a ‘hot’ subduction zone, a consequence of a young subducting plate with a thick sediment cover and a low convergence rate with respect to North America (Hyndman & Wang, 1993). In ‘hot’ subduction settings, the bulk of the fluid from the subducting slab is released to the accretionary wedge (Kirby et al., 2002; Peacock et al., 2002; van Keken et al., 2002; Hacker et al., 2003b; Hyndman & Peacock, 2003) and the capacity for flux melting of the mantle wedge is reduced. However, the occurrence of basalts with slab-derived contributions and relatively high water contents at many Cascade volcanic centers (e.g. Grove et al., 2002; Ruscitto et al., 2012) contradict the conventional wisdom that ‘hot’ subduction zones are dry. Consensus has not been reached on the source of water in Cascade arc primary basalts, whether from subducting sediment, altered oceanic crust, and/or hydrated oceanic mantle, and it is not clear whether water is released as a component of fluid, melt, supercritical fluid, or combination thereof. In the Mt. Baker volcanic field, basaltic lavas are rare and largely peripheral to the main magmatic focus (Fig. 2). The lavas have variable Mg/(Mg þ Fe) values, indicating that differentiation has modified the original mantle-derived primary magma compositions. In this study, we use petrological data, whole-rock major, minor, and trace element abundances, and isotopic compositions (Sr, Nd, Pb) for the most geochemically primitive lavas at Mt. Baker to (1) reconstruct the primary basaltic magma compositions and establish their compositional diversity, (2) constrain mantle source mineralogy and the temperatures, pressures, initial H2O content, and redox state associated with mantle melting, and (3) characterize the composition of the subduction component and determine its source (sediment and/or oceanic crust), identity (fluid and/or melt), abundance in the mantle wedge, and effect on mantle melting. The results are synthesized into a petrogenetic model for basalt generation beneath Mt. Baker within a framework of petrological and geophysical constraints. Our conclusions expand upon, and contrast with, previous work on the petrogenesis of mafic Mt. Baker lavas (Green, 1988, 2006; Moore & DeBari, 2012).

NUMBER 2

MULLEN & McCALLUM

52°N

MT. BAKER BASALTS, CASCADES

1 3 1°W

119°W

1 2 5°W

Queen Charlotte Fault

Volcanic rocks Silverthrone Franklin Glacier

0.7 50°N

Exp Rid lorer ge

Bridge River

BELT

Salal Glacier Meager Cayley Garibaldi

VANCOUVER

Pacific Plate

Canada USA

MT. BAKER VOLCANIC Glacier FIELD

4.6 idg e

48°N

Major volcanic centers

46°N

Jua nd e Fu

ca R

SEATTLE Rainier

Juan de Fuca Plate

St. Helens

PORTLAND

44°N

HIGH CASCADES Bla

nc oF au

Hood

North America Jefferson Plate Three Sisters Broken Top

2.9

lt

Adams

Newberry Mazama

e

42°N

Gor da R

idg

McLoughlin

40°N

200 km

Gorda Plate

Mendocino Fault

Shasta

San Andreas Fault

Medicine Lake

Lassen

Fig. 1. Map of the Cascade Arc and its tectonic setting. The extent of the Garibaldi Volcanic Belt and High Cascades segments of the arc are indicated with black arrows. Volcanic and plutonic rocks are shown in light gray and dark gray shading, respectively. Area enclosed by small rectangle is enlarged in Fig. 2. Compiled from Green et al. (1988), Monger (1989), Wheeler & McFeely (1991) and Du Bray et al. (2006). Oceanic plate configurations are from Braunmiller & Nabelek (2002) and Audet et al. (2008). Juan de Fuca plate convergence vectors (cm a1) are from Riddihough & Hyndman (1991) and McCrory et al. (2004) for a reference frame fixed relative to North America; vectors are not corrected for clockwise rotation of the Oregon forearc (Wells et al., 1998). Explorer plate convergence vectors are from Braunmiller & Nabelek (2002).

243


lt Explorer Fau a Plate tk oo GARIBALDI N 2.0

Plutonic rocks


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Mt. Baker basalts 3000

Other volcanic rocks 50

Plutonic rocks

Hannegan caldera

00

Kulshan caldera

00 70

500

0

Black + Buttes

+

Mount Baker +

5000

Sherman Crater 0

00

900

Chilliwack composite batholith

30

+

30

Lake Ann stock

00

Cathedral Crag

1 82 Park Butte

Sulphur Creek

5

112

Tarn Plateau

N

5 km

114 Lake Shannon

122

Fig. 2. Geological map of the Mt. Baker volcanic field after Hildreth et al. (2003), Tabor et al. (2003) and Tucker (2006). Circles with sample numbers indicate basalt sample localities. Elevation contours (gray dotted lines) are given in feet.

subsurface extension of Wrangellia (McGroder, 1991; Mullen, 2011). On its eastern edge, the Mt. Baker volcanic field overlaps the 34^2 Ma Chilliwack composite batholith, a product of older Cascade arc magmatism (Tepper et al., 1993; Tepper, 1996).

SELECT ION OF T H E PR I M I TI V E SU ITE In previous studies of Cascade arc basalts, criteria that have been used to judge basalts as primitive include Mg/(Mg þ Fe2þ) 0·60, MgO 47 wt %, low phenocryst

contents (55%), olivine close to equilibrium with mantle compositions (i.e. Fo 86), and relatively high concentrations of compatible trace elements (e.g. Cr4200 ppm, Ni4100 ppm) (Bacon et al., 1997; Borg et al., 1997; Conrey et al., 1997; Grove et al., 2002; Schmidt et al., 2008). Although none of the lavas at Mt. Baker meets all of these criteria, six lavas meet several of the criteria. Assuming 53 wt % SiO2 as an upper limit for basalt, three of the samples would be more appropriately defined as basaltic andesites. From oldest to youngest (dates from Hildreth et al., 2003), the six most primitive magmas are the Park Butte basalt (716 ka), Cathedral Crag basaltic andesite

244


Hogback Camp 97

MULLEN & McCALLUM


(331ka), Hogback Camp basaltic andesite (306 ka), Tarn Plateau basaltic andesite (203 ka), Lake Shannon basalt (94 ka), and Sulphur Creek basalt (9·8 ka). Sample localities, field occurrences and petrographic descriptions are provided in Supplementary Data Electronic Appendix 1 (supplementary data are available for downloading at http://www.petrology.oxfordjournals.org) and, with the exception of the Hogback sample, have also been discussed by Moore & DeBari (2012). The basalts erupted from monogenetic vents peripheral to the main magmatic focus, except for the Hogback basaltic andesite, which is part of the sequence of lavas forming the extinct Black Buttes volcano (Hildreth et al., 2003).

Electron microprobe analyses were conducted using a four-spectrometer JEOL 733 Superprobe at the University of Washington using a fixed accelerating voltage of 15 kV. Natural and synthetic standards were used for calibration. Mineral compositions were measured using a beam diameter of 1^5 mm and a 15 nA current for Fe^Ti oxides, Cr-spinel, olivine and pyroxene, and a 10 nA current for feldspars. K and Na were analyzed first to minimize the effects of count loss during the analysis. Counts were collected at peak positions and at both background positions. Data reduction employed the ZAF correction procedures of Armstrong (1988). Analytical errors are listed in Supplementary Data Electronic Appendix 2.

Whole-rock major and trace elements Major elements and 19 trace elements were measured on whole-rock samples by X-ray fluorescence (XRF) and 27 trace and rare earth elements (REE) by inductively coupled plasma mass spectrometry (ICP-MS) at the Washington State University GeoAnalytical Laboratory. Samples were crushed in a hardened steel jaw crusher and ground in an agate ball mill. Fused beads for XRFanalysis were prepared after mixing powders with dilithium tetraborate flux, then ground and re-fused to ensure sample homogeneity. Data are presented as analyzed and all iron is reported as FeO (Table 1). Details on preparation techniques, analytical methods, and the determination of accuracy and precision have been described by Johnson et al. (1999) and are given on the laboratory website (http:// www.sees.wsu.edu/Geolab/index.html). Uncertainties for major element oxides are less than 2% relative (except K2O at 3^7%), less than 5% for the REE, and less than 10% for the other trace elements.

Sr, Nd, and Pb isotopes Sr, Nd, and Pb isotope ratios were measured at the University of Washington (UW), in most cases using

245


A N A LY T I C A L M E T H O D S Electron microprobe

aliquots of the same powders as used to measure major and trace element abundances. Samples were pulverized in an alumina or tungsten carbide shatterbox or handcrushed in an agate mortar. Sr isotope measurements were made by thermal ionization mass spectrometry (TIMS) using a VG Sector 54 system after loading onto Ta filaments following the procedures described by Nelson (1995). 87Sr/86Sr ratios were corrected for instrumental mass fractionation using 86Sr/88Sr ¼ 0·1194. Over the course of study, the mean 87Sr/86Sr of the SRM 987 standard was 0·710235 28 (2s; n ¼ 7), within error of the accepted value (0·710248). Long-term reproducibility of the standard is better than 0·00002, or 30 ppm (2s), greater than within-run precisions (Table 1). A Nu Plasma multicollector (MC)-ICP-MS was used to measure Nd and Pb isotope ratios following analytical procedures described by Harkins et al. (2008) and Gaffney et al. (2007), respectively. Nd was analyzed over a 3 day period and the Ames standard was run every three samples, yielding a mean 143 Nd/144Nd of 0·512110 16 (2s; n ¼ 25), within error of the accepted value (0·512130). Reported isotopic ratios were corrected by referencing to the daily mean of standard analyses. In-run errors are smaller than the long-term reproducibility of the standard (25 ppm, 2s). Pb was measured over a 2 day period and SRM981 was analyzed every two samples. Reported isotope ratios were corrected for instrumental mass fractionation using the Tl spiking technique (using NIST 997 thallium) and assuming exponential fractionation behavior. Over the course of study, SRM981 yielded average values of 206 Pb/204Pb ¼16·930 3, 207Pb/204Pb ¼15·481 3, and 208 Pb/204Pb ¼ 36·668 7 (2s; n ¼ 43). Reported Pb isotopic ratios are normalized by sample^standard bracketing to 206Pb/204Pb ¼16·9405, 207Pb/204Pb ¼15·4963, and 208 Pb/204Pb ¼ 36·7210 for SRM981 (Galer & Abouchami, 1998). In-run errors are smaller than the long-term reproducibility of the Pb standard (150 ppm, 2s). Discrepancies among replicate and duplicate analyses of the Mt. Baker samples (Table 1) are within 2s reproducibilities in all cases. Several samples were reanalyzed at the University of British Columbia on a Nu Plasma MC-ICP-MS system (Pb, Nd) and a Thermo Finnigan Triton TIMS system (Sr) following procedures described by Mullen & Weis (2013), and resulting isotope ratios are within analytical error of UW measurements (Table 1). For direct comparison with isotope data obtained from the literature, all data have been normalized to standard values of 87Sr/86Sr ¼ 0·710248 (SRM987) or 0·708028 (Eimer & Amend); 143Nd/144Nd ¼ 0·511973 (Rennes), 0·511858 (La Jolla), 0·512633 (BCR-1), or 0·512130 (Ames) (Weis et al., 2006); and 208Pb/204Pb ¼ 36·7219, 207 Pb/204Pb ¼15·4963, 206Pb/204Pb ¼16·9405 for SRM981 (Galer & Abouchami, 1998).


VOLUME 55

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Table 1: Major and trace element compositions, isotopic ratios, and mineral modes for Mt. Baker basalts Sample name:

Cathedral Crag

Tarn Plateau

Park Butte

Hogback Camp

Sulphur Creek

Lake Shannon 114

Lake Shannon 122

Sample no.:

02-MB-1

02-MB-5

06-MB-82

06-MB-97

07-MB-112

08-MB-114

08-MB-122

Group:

II

II

III

I

I

I

I

Unnormalized major elements (wt %) SiO2 TiO2

53·30 1·142

53·69 0·919

50·57 1·044

54·45 1·597

52·56 1·649

52·06 1·471

52·62 1·280

20·23

16·40

17·25

17·88

17·63

17·95

17·64

FeO*

6·52

7·08

9·44

8·15

8·73

8·51

7·88

MnO

0·112

0·135

0·165

0·154

0·165

0·157

0·147

MgO

3·96

7·76

8·38

3·74

5·42

6·44

7·57

CaO

9·40

9·74

9·41

8·16

8·39

8·89

8·32

Na2O

3·88

2·87

3·20

4·24

4·44

4·09

3·78 0·70

K2O

0·97

0·84

0·43

0·94

0·90

0·62

P2O5

0·331

0·179

0·145

0·369

0·430

0·281

Sum

99·85

99·61

100·05

99·66

100·31

100·48

0·228 100·17

Unnormalized trace elements (ppm) Concentrations from XRF Ni

23

61

71

13

40

81

134

Cr

24

214

251

25

90

163

255

Sc

21

29

31

25

27

27

25

V

204

181

184

218

200

175

163

Ba

425

344

201

400

279

225

271

Rb

11

8

5

11

11

7

10

Sr

1222

861

508

709

557

476

481

Zr

103

109

73

190

205

152

141

Y

17

17

21

32

31

26

24

Nb

5·3

3·7

1·4

7·3

7·1

5·2

4·3

Ga

21

18

17

19

18

16

15

Cu

38

37

18

42

34

46

50

Zn

71

69

80

92

83

73

70

Pb

6

5

0

3

1

0

0

La

22

12

9

18

15

15

12

Ce

41

31

14

41

38

27

27

Th

1

1

1

1

1

0

0

Nd

22

16

8

23

22

16

18

1

0

U Concentrations from ICP-MS La

17·09

12·71

7·65

18·61

16·18

11·49

11·18

Ce

38·11

28·26

17·44

43·81

38·61

27·27

25·82

Pr

5·21

3·91

2·52

6·03

5·31

3·78

3·58

Nd

22·19

16·34

11·32

25·30

23·02

16·65

15·66

Sm

4·71

3·66

2·91

6·00

5·69

4·25

4·00

Eu

1·56

1·17

1·14

1·92

1·92

1·53

1·43

Gd

4·16

3·56

3·38

6·08

5·93

4·73

4·28

Tb

0·62

0·56

0·59

1·03

1·00

0·82

0·73

Dy

3·55

3·45

3·84

6·24

6·27

5·14

4·64

Ho

0·71

0·70

0·81

1·29

1·26

1·08

0·97

(continued)

246


Al2O3

MULLEN & McCALLUM


Table 1: Continued Sample name:

Cathedral Crag

Tarn Plateau

Park Butte

Hogback Camp

Sulphur Creek

Lake Shannon 114

Lake Shannon 122

Sample no.:

02-MB-1

02-MB-5

06-MB-82

06-MB-97

07-MB-112

08-MB-114

08-MB-122

Group:

II

II

III

I

I

I

I

Er

1·86

1·83

2·19

3·48

3·40

2·96

2·67

Tm

0·27

0·28

0·33

0·51

0·49

0·43

0·38

Yb

1·60

1·66

1·98

3·10

3·03

2·68

2·44

Lu

0·26

0·26

0·33

0·50

0·49

0·43

Ba

426

335

199

402

279

215

0·38 266

1·86

2·15

0·60

1·82

1·35

1·10

Nb

4·95

3·57

2·56

7·30

7·40

5·13

4·60

Y

17·95

17·80

20·39

32·22

31·96

27·19

24·29

Hf

2·63

2·89

1·89

4·59

4·27

3·29

3·19

Ta

0·28

0·23

0·16

0·46

0·48

0·36

0·33

U

0·58

0·67

0·22

0·67

0·56

0·47

0·57

Pb

3·69

3·46

1·54

3·69

3·26

2·74

3·14

8·8

5·0

7·4

9·7

0·19

0·05

Rb

11·2

Cs

0·10

Sr

1194

Sc

24·0

Zr

99

860 33·4 104

502 29·7 70

11·6 0·23 712 24·2 183

10·0 0·27 563 26·0 200

0·17 486 25·2 145

1·34

0·25 493 22·7 136

Isotopic ratios 87

Sr/86Sr

0·703513(8)

dup 144

Nd

eNd

0·703248(7)

0·512986

0·512993

0·5130341

0·513037

0·513028

5·1

7·0

6·8

6·9

7·7

7·8

7·6

0·5130011

0·512985

0·512988

Pb/204Pb

38·3598

dup

38·2490

38·2661

38·2597

38·3104

38·2753

38·30771

38·2432 38·24951

dup2 204

Pb

15·5645

dup

15·5510

15·5518

15·5529

15·5587

15·5560

15·55751

15·5491 15·55051

dup2 Pb/

0·703213(7)

0·512995

rep 208

206

0·703240(7)1

0·512899

dup

Pb/

0·703173(10)

0·703109(7)1

Nd/

207

0·703156(7)

0·703116(8)

dup2 143

0·703137(7)

204

Pb

18·8515

18·7949

dup

18·7936

dup2

18·79751

18·8286

18·8356

18·8375

18·8466

18·83851

Phenocryst modes (% of whole rock) olivine plagioclase clinopyroxene

7

4

22

1

1

10

2

15

8

4

10

5

15

5

8

10

0

0

0

0

0

Major groundmass minerals plag4cpx

plag4cpx4

plag4cpx4

plag4ol4

plag

plag4ol4

ol plag4

opx4pig

opx

ol

cpx

ol cpx

cpx

cpx

1

Measured at the University of British Columbia (other isotope data from the University of Washington). Within-run 87Sr/86Sr errors are given in parentheses and represent errors in the last significant digit. eNd values represent the deviation from chondritic 143Nd/144Nd ¼ 0·512638 in parts in 104. Replicates are repeat analyses of the same sample solution; duplicates are full procedural duplicates starting with a second aliquot of the same sample powder. ol, olivine; cpx, clinopyroxene; plag, plagioclase; opx, orthopyroxene; pig, pigeonite.

247


Th


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GEOC H EM ICA L C H A R AC T E R I S T I C S O F M T. B A K E R B A S A LT S Major and trace elements Major and trace element abundances of the Mt. Baker basalts are reported in Table 1 and illustrated in Figs 3^5. With the exception of Hogback, additional sample analyses

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have been provided and discussed by Moore & DeBari (2012). The samples are classified as basalts and basaltic andesites with SiO2 contents of 50·5^54·6 wt %. Mg# (100 Mg/[Mg þ Fe2þ]) calculated on a molar basis ranges from 71 (Tarn Plateau) to 58 (Sulphur Creek), assuming a Fe3þ/Fe value of 0·20 for all samples [based on calculated oxygen fugacities and equation (7) of Kress & Carmichael

(a) Shoshonite

wt.% K2O

High K

2 Med K

1 Low K

0

43 12

47

51

55

59

63

(b)

71

75

Group I Sulphur Creek Lk Shannon 122 Lk Shannon 114 Hogback Camp Group II Tarn Plateau Cathedral Crag Group III Park Butte other Mt. Baker lavas primitive Cascade arc

10 8

wt.% MgO

67

6 4 2 0

43

47

51

55

59

63

67

71

75

wt.% SiO2 Fig. 3. Major element variation diagrams for the Mt. Baker basalts (large symbols). (a) K2O vs SiO2; (b) MgO vs SiO2. Other Mt. Baker lavas are shown with small white circles; data from Hildreth et al. (2003) and Mullen (2011). Other lavas from the Cascade arc considered as primitive are shown as small gray circles (data from Bacon et al., 1997; Borg et al., 1997, 2002; Conrey et al., 1997; Green & Harry, 1999; Grove et al., 2002; Green & Sinha, 2005; Leeman et al., 2005; Green, 2006; Schmidt et al., 2008, and additional references therein). Asterisks indicate Mt. Shasta sample 85-44, a primitive calc-alkaline basalt used as starting material in studies on the phase equilibria of hydrous arc basalts (Baker et al., 1994; Gaetani & Grove, 1998; Mu«ntener et al., 2001).

248


3

MULLEN & McCALLUM


(a) 100

10

1

0.1Cs RbBaTh U NbTa K LaCePbPr SrNdSmZr HfEu Ti GdTbDyY HoEr TmYb LuScZnCr Ni V

(b) 600

1

3G

HMBA

400

OIB

CAB

Ba 200

ALK

5

LKOT

0

0

10

7

1 1

3 GPa

1 GPa

10

1G

1

Pa

CAB: calc-alkaline basalt HMBA: high-Mg basaltic andesite LKOT: low-K olivine tholeiite OIB: intraplate-type basalt, tholeiitic ALK: intraplate-type basalt, alkalic

3

10 10 5 33

3

Pa

20

30

40

50

Nb Fig. 4. (a) Trace elements in Mt. Baker basalts, normalized to N-MORB (Sun & McDonough,1989). Light gray field is for Mt. Baker andesites, dacites and rhyolites (Mullen, 2011). (b) Ba vs Nb (ppm). Basalt groups in the Cascade arc (gray fields) are LKOT (low-K olivine tholeiite; small filled black circles), CAB (calc-alkaline basalt; asterisks), HMBA (high-Mg basaltic andesite and andesite; white diamonds), OIB (hypersthene-normative intraplate-type basalts), and ALK (nepheline-normative intraplate-type alkalic basalts). Data sources as in Fig. 3. Bold gray and black curves are anhydrous mantle melts at 1 and 3 GPa, respectively, of DMM (depleted mantle; Workman & Hart, 2005) (dashed curves) and PM (primitive mantle; Sun & McDonough, 1989) (continuous curves). Tick marks with numbers indicate per cent partial melt. Melt compositions calculated using a batch melting equation with mineral/melt partition coefficients from Table 2 and residual mantle mineral assemblages determined by BATCH modeling (Longhi, 2002).

(1991)]. The basalts are hypersthene-normative, low-K (Park Butte) to medium-K (Gill, 1981) (Fig. 3a), and calcalkaline (Miyashiro, 1974). Relative to normal mid-ocean ridge basalt (N-MORB), the samples have elevated abundances of large ion lithophile elements (LILE: Cs, Ba, Rb, K, Sr, U) and Pb, but similar contents of high field strength elements (HFSE: Ta, Nb), resulting in high LILE/HFSE values (Fig. 4). The samples are enriched in light REE (LREE) relative to

heavy REE (HREE) (Fig. 5). These geochemical characteristics are common in magmas generated in subduction settings worldwide (e.g. Pearce & Peate, 1995). The basalts define three groups with distinct trace and major element characteristics (Moore & DeBari, 2012). Group I (Sulphur Creek, Lake Shannon, and Hogback Camp) has the smallest negative Nb anomalies (as measured by chondrite-normalized Nb/La; Briqueu et al., 1984), lowest LILE/HFSE ratios, and highest Yb (Fig. 5).

249


Sample/N-MORB

Group I Sulphur Creek Lk Shannon 122 Lk Shannon 114 Hogback Camp Group II Tarn Plateau Cathedral Crag Group III Park Butte

intermediate–silicic Mt. Baker lavas


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other Mt. Baker lavas

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La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

(La/Yb)N

16 12

CAB: calc-alkaline basalt HMBA: high-Mg basaltic andesite LKOT: low-K olivine tholeiite Mt. Baker, other lavas

(b)

HMBA

Fractionation Paths: cpx opx oliv plag

CAB

8 4 LKOT

0 0.5

1.5

Yb

2.5

3.5

Fig. 5. (a) Chondrite-normalized rare earth element abundances. Chondrite values from McDonough & Sun (1995). Light gray field is for Mt. Baker andesites, dacites and rhyolites (Mullen, 2011). (b) (La/Yb)N vs Yb (ppm). Gray squares, Mt. Baker andesites, dacites and rhyolites (Mullen, 2011). HMBA, CAB and LKOT are end-member Cascade arc basalt groups (data sources as in Fig. 3). Group II basalts (Tarn Plateau and Cathedral Crag) cannot be derived from a Group I or Group III parent by fractionation, as shown by vectors (bold black arrows) for 15% fractional crystallization of olivine (ol), clinopyroxene (cpx), orthopyroxene (opx), and plagioclase (plag), calculated using the Rayleigh equation and partition coefficients fromTable 2. Group II also is not explained by magma mixing with a Group I or III parent because no intermediate or silicic Mt. Baker lava has sufficiently lowYb to serve as a mixing end-member. Groups I and III are not related by fractionation or accumulation because Mg and Ni contents are similar between the two groups, and pronounced Eu anomalies are absent.

Group II samples (Tarn Plateau and Cathedral Crag) are more enriched in LILE than the other basalts and have the most pronounced negative Nb anomalies, steepest REE patterns ([La/Yb]N ¼ 5·2^7·3), and lowest Yb contents (Fig. 5). Cathedral Crag has high Sr (1200 ppm) and a small positive Eu anomaly (defined as Eu/Eu*, where Eu* ¼ [Sm þ Nd]/2). The Group III basalt (Park Butte) has LILE/HFSE ratios similar to Group II, but low

(La/Yb)N (similar to Group I), the lowest LREE, and the lowest total abundances of trace elements and REE (Figs 4 and 5).

Isotopic compositions Sr, Nd, and Pb isotope compositions are given in Table 1. Additional Sr and Nd isotope data for five samples (exception: Hogback) and Pb isotopic data for three samples

250


10

MULLEN & McCALLUM


Olivine

Plagioclase

Olivine phenocrysts are euhedral, subhedral, embayed, or skeletal (Fig. 8a). The most magnesian olivines in each sample (Fig. 9) are more fayalitic than mantle olivine (Fo90 2), indicating olivine fractionation from the more primitive magmas. Normal zoning is prevalent, although Lake Shannon 122 contains a population of reversely zoned antecrysts. Melt inclusions (Fig. 8b) are microcrystalline. Olivine is present as a groundmass mineral in Sulphur Creek, Lake Shannon, and Hogback. In Cathedral Crag, the smallest olivines are mantled by clinopyroxene; inTarn Plateau they are resorbed and rimmed by symplectitic orthopyroxene þ magnetite.

Plagioclase is a phenocryst, microphenocryst, and groundmass mineral in all samples. Compositional ranges are shown in ternary diagrams in Supplementary Data Electronic Appendix 3A. Phenocrysts display a range of complex compositional patterns including oscillatory zoning, gradational zoning (normal or reverse), or stepped zoning (normal or reverse) (Fig. 10a^c). The most anorthitic plagioclase compositions occur in phenocryst cores and the most albitic compositions in microlites, crystal rims, and resorbed crystal interiors. Plagioclase compositions in Cathedral Crag, Tarn Plateau, and Park Butte extend to the most anorthitic compositions at An86, whereas Sulphur Creek has the lowest maxima at An79. In Tarn Plateau, Lake Shannon 114, and Hogback, plagioclase cores show evidence of partial dissolution in the form of patchy zoning and fine to coarse sieve textures. In Lake Shannon 122 and Sulphur Creek, many plagioclase cores are bounded by sharp dissolution surfaces and occur as two distinct compositional populations (Fig. 10a and c). Cathedral Crag and Park Butte have compositional gaps

S U M M A RY O F P E T RO G R A P H Y A ND MI NER A L COMPOSITIONS

Relationship between olivine and whole-rock compositions ol=wr If the liquid phase is not preserved, KDðFeMgÞ values (where wr refers to whole-rock) can be used to deduce details of olivine crystallization (Rhodes et al., 1979). Assuming that the whole-rock composition represents the initial liquid composition, the composition of olivine phenocryst cores represents the olivine in equilibrium with the

251


Detailed petrographic descriptions and selected mineral compositions are provided in Supplementary Data Electronic Appendices 1 and 2, respectively. All samples are porphyritic or micro-porphyritic with euhedral phenocrysts of olivine and plagioclase; Cathedral Crag and Tarn Plateau also contain clinopyroxene phenocrysts. Park Butte, Tarn Plateau, and Cathedral Crag are holocrystalline whereas the other basalts contain minor interstitial glass (510%). Groundmass textures are intersertal to sub-ophitic with plagioclase4clinopyroxene4 titanomagnetite ilmenite. Trace minerals include zircon, apatite allanite.

ol=wr value is coexisting initial melt if the KDðFeMgÞ ol=liquid 0·32 0·03; that is, the equilibrium KDðFeMgÞ value for olivine^melt pairs (Toplis, 2005). If the sample has a high ol=wr value greater modal abundance of olivine, a KDðFeMgÞ than 0·32 0·03 may be interpreted as evidence of olivine accumulation. However, as shown in the inset to Fig. 9, ol=wr 40·32 0·03 can be also achieved by closedKDðFeMgÞ system equilibrium crystallization. wr values are plotted against the range of In Fig. 9, XMg olivine compositions in each sample. The most magnesian olivine cores in Lake Shannon and Hogback Camp give ol=wr values of 0·31 and 0·33, respectively, indicating KDðFeMgÞ that the olivine phenocrysts most probably crystallized from liquids with the same compositions as the wholerocks. In Cathedral Crag and Tarn Plateau, the most ol=wr values of 0·39 Mg-rich cores of olivines give KDðFeMgÞ and 0·43, respectively; the low abundance of olivine phenocrysts is consistent with an initial stage of equilibrium crysol=wr values, tallization producing slightly elevated KDðFeMgÞ although some olivine accumulation cannot be ruled out. ol=wr value of 0·29 is consistent In Sulphur Creek, the KDðFeMgÞ with the presence of olivine antecrysts formed from a more magnesian precursor magma. ol=wr (0·9) with olivine Park Butte has the highest KDðFeMgÞ phenocrysts that are relatively fayalitic, Fo68 to Fo61. This sample has a relatively low olivine mode and a bulk composition that plots close to the low-pressure olivine^ plagioclase cotectic in the basalt system, indicating that it represents a liquid composition. Because the coarsegrained texture of this sample is consistent with slow coolol=wr in Park Butte to exing, we attribute the high KDðFeMgÞ tended equilibrium crystallization as shown in Fig. 9 (inset).

(Park Butte, Tarn Plateau, Lake Shannon) were reported by Moore & DeBari (2012). The eNd values of Mt. Baker basalts are among the highest measured in the Cascade arc, and Pb isotopic ratios are among the least radiogenic (Fig. 6). In Pb^Pb isotope plots, primitive Cascade arc basalts define a mixing array between Juan de Fuca MORB and northern Cascadia sediment (Fig. 6b and c). Isotope data for intermediate and felsic lavas from Mt. Baker (Mullen, 2011) are plotted together with the basalt data in Fig. 6. The volcanic field defines near-linear Sr^ Nd^Pb isotopic arrays that are correlated with whole-rock SiO2 contents (Fig. 7), reflecting assimilation of slightly more radiogenic, silicic crust or mixing with crustal melts (Mullen, 2011). Except for Cathedral Crag, which has higher Sr and Pb isotope ratios and lower eNd, the basalts plot in a cluster at the primitive end of the isotopic arrays defined by the volcanic field (Figs 6 and 7).


VOLUME 55

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12

(a)

Cascadia MORBs

10

1 3

8

εNd

2

6

primitive Cascade arc

Mt. Baker volcanic field

4

N. Cascadia sediment

Mt. Baker Volcanic Field

2 0.7033

(b) 39.10

0.7043

87

0.7053

86

Sr/ Sr


38.5

RL

RL

NH

NH

38.4

38.70 Mt. Baker basalts

primitive Cascade arc

38.3

38.30

18.75

18.85

Group III basalts Park Butte

18.95


37.90

primitive Cascade arc Juan de Fuca MORB Explorer MORB

Cascadia MORBs

(c)

Gorda MORB N. Cascadia sediment N. Cascadia sediment

15.59

15.65

Group I basalts Sulphur Creek Lk Shannon 122 Lk Shannon 114 Hogback Camp Group II basalts Tarn Plateau Cathedral Crag

38.2

208

Pb/204Pb

andesites dacites rhyodacites rhyolites

primitive

15.57 Cascade

arc

Pb/204Pb

Mt. Baker basalts 15.55 L R H

L

NHR

N

15.55

18.75

18.85

18.95

207


15.45 18.10

Cascadia MORBs

18.50

18.90

206

19.30

204

Pb/ Pb

Fig. 6. (a) eNd vs 87Sr/86Sr, (b) 208Pb/204Pb vs 206Pb/204Pb, and (c) 207Pb/204Pb vs 206Pb/204Pb for Mt. Baker basalts and other Mt. Baker lavas (Mullen, 2011). Also shown are northern Cascadia subducting sediment from ODP sites 1027 and 888 (Carpentier et al., 2010), northern Juan de Fuca MORB (Cousens et al., 1995), Explorer MORB (B. Cousens, unpublished data), southern Gorda MORB (Davis et al., 2008), and other primitive basalts from the Cascade arc (references as in Fig. 3, and Mullen & Weis, 2013). NHRL, Northern Hemisphere Reference Line values (discussed in the text): (Hart, 1984). The curves (marked 1, 2, 3) in (a) are model solutions for Tarn Plateau using three Dmetabasalt=fluid Sr Curve (1), 2·9 from Kessel et al. (2005a); Curve (2), 0·19 from Brenan et al. (1995); Curve (3), an intermediate value of 0·6. Tick marks are shown at intervals of 2% SC addition to the mantle wedge. The 2s error bars are shown in the corner of (a) and the insets to (b) and (c).

252


0 0.7023

MULLEN & McCALLUM


0.7040

Sr/86Sr 87

0.7035

0.7030 48 38.55

53

58

63

68

73

78

(b)

38.35

208

Pb/204Pb

38.45

Chromian spinel

38.25

38.15 48

53

58

63

wt.% SiO2

68

73

78

Mt. Baker volcanic field Group I basalts Sulphur Creek Lk Shannon 122 Lk Shannon 114 Hogback Camp

Group II basalts Tarn Plateau Cathedral Crag Group III basalts Park Butte

andesites dacites rhyodacites rhyolites

Fig. 7. (a) 87Sr/86Sr vs wt % SiO2 and (b) 208Pb/204Pb vs wt % SiO2 for the Mt. Baker volcanic field.

Chromian spinel occurs as small (typically 550 mm) inclusions in olivine phenocrysts and as isolated microphenocrysts in Lake Shannon and Sulphur Creek (Fig. 8c and d). Spinel compositions (Electronic Appendix 3C) lie close to the spinel (MgAl2O4)^chromite (FeCr2O4)^ magnetite (Fe3O4) plane in the spinel prism. Cr/ (Cr þAl) and Mg/(Mg þ Fe) are inversely correlated and highly variable, indicating trapping of spinel grains at different stages of crystallization. Spinels enclosed in olivine and spinel microphenocrysts show compositional gradations of increasing Cr and Ti, and decreasing Al, that are related to progressive fractionation. Groundmass spinels lie close to the magnetite (Fe3O4)ûlvo«spinel (Fe2TiO4) join.

of 10 mol % An separating anorthitic cores from all other plagioclase compositions in each sample (Fig. 10b). Cathedral Crag also displays a second compositional gap of 17 mol % that separates the thin albitic rims from interior zones. In Tarn Plateau, plagioclase antecrysts have cores with pervasive coarse sieving, and in Lake Shannon, antecrysts are rounded and anhedral, compositionally homogeneous, and typically occur in annealed troctolitic clots. In both samples, however, the ranges of core and rim compositions exhibited by possible antecrysts are similar to those of phenocrysts in the same sample, implying a common liquid line of descent.

Fe^Ti oxides

Pyroxene

R E C O N S T RU C T I O N O F P R I M A RY M AG M A C OM P O S I T I ON S

Clinopyroxene is a constituent of the groundmass of all the basalts and is a phenocryst phase in Tarn Plateau and Cathedral Crag; these two samples also contain groundmass orthopyroxene and (in Cathedral Crag only) pigeonite. Compositional ranges are shown in pyroxene

Titanomagnetite occurs as microphenocrysts and a groundmass mineral in all samples. Oxy-exsolution of ilmenite on {111} and/or sandwich lamellae of ilmenite (Fig. 8e and f) are restricted to basalts that also have primary ilmenite. Sandwich lamellae (Fig. 8f) usually occur as relatively wide intergrowths on {111} planes of the host spinel and we interpret these as primary intergrowths because they record higher temperatures than oxy-exsolution lamellae. Primary ilmenite is present only in Tarn Plateau, Cathedral Crag and Park Butte.

It is necessary to constrain primary basaltic magma compositions in order to decipher mantle source compositions. However, primary magmas are rare or absent at arc volcanic centers, as most of the basalts have undergone

253


quadrilaterals in Supplementary Data Electronic Appendix 3B together with olivine compositional ranges. In Cathedral Crag, many clinopyroxene phenocrysts have magnesian cores (Wo47En48) with subtle oscillatory or sector zoning and sharply defined Fe-rich rims (Wo38En48) (Fig. 10d). Tarn Plateau clinopyroxene phenocrysts (Fig. 10e and f) display a variety of zoning patterns (normal, reverse, oscillatory and sector). The most magnesian cores are Wo42^43En50^51. Approximately 30% of the pyroxene phenocrysts in Tarn Plateau have distinct Fe-rich cores rimmed by wide magnesian zones (Fig. 10f). The Fe-rich cores crystallized from a melt composition farther along the liquid line of descent, and were partially resorbed prior to crystallization of more magnesian rims. An anomalous magnesian and aluminous orthopyroxene (Wo2En92, 3·2 wt % Al2O3) occurs as an inclusion in a Tarn Plateau clinopyroxene phenocryst and appears to be a relict phase from the sub-arc mantle.

(a)


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Fig. 8. Backscattered electron images: (a) skeletal olivine phenocrysts in Sulphur Creek; (b) olivine with microcrystalline ‘melt’ inclusion, Hogback Camp; (c) chromian spinel inclusions in olivine, Park Butte; (d) skeletal chromite, Sulphur Creek; (e) titanomagnetite with ilmenite ‘sandwich’ lamellae in lower left and fine oxy-exsolution lamellae of ilmenite on {111} in upper right, Park Butte; (f) titanomagnetite with sandwich lamellae of ilmenite, Park Butte. mt, titanomagnetite; ilm, ilmenite.

254

MULLEN & McCALLUM


Group II Group I Sulphur Creek (n=12/51) Tarn Plateau (n=14/45) Lk Shannon 122 (n=9/57) Cathedral Crag (n=10/26) Lk Shannon 114 (n=21/104) Group III Hogback Camp (n=14/23) Park Butte (n=9/25)

1.0 Mantle melts

0.9

0.7 0.6

0.5Qz

Silica

0.5

En

En60

Pxss

Fs L2

0.4

L1

0.3 0.4

Fo

0.5

Fo85

L3

Olss Fa

Fo62

0.6 0.7 0.8 XMg (liquid or whole rock)

0.9

1.0

ol=wr Fig. 9. Roozeboom diagram for XMg in olivine vs XMg in whole-rock (wr) or liquid (liq), illustrating the range of KDðFeMgÞ values for olivine in ol=liq the Mt. Baker basalts. Continuous curve with gray envelope indicates an equilibrium KDðFeMgÞ value of 0·32 0·03 (Toplis, 2005). The dashed ol=wr value of 1·0. The white star indicates a mantle melt in equilibrium with Fo90. n: first number is the number of olivine crysline is for a KDðFeMgÞ tals analyzed per sample; second number is the total number of points analyzed per sample. Inset shows phase equilibria in the Fo^Fa^SiO2 system (after Bowen & Schairer, 1935) recalculated in oxygen units at a pressure of 200 MPa. Initial liquid (L1; black circle) has XMg ¼ 0·64 and is in equilibrium with Fo85 olivine (dash^dot line). During equilibrium crystallization, the liquid follows the path from L1 to L2 to L3 at which point the last liquid [XMg ¼ 0·34 in equilibrium with Fo62 olivine and En60 pyroxene (dotted three-phase triangle)] is used up. At L3, ol=wr ¼ 1·09. This example illustrates how the olivine compositions observed in the Park Butte basalt (Fo68 to Fo61) could be generated, as KDðFeMgÞ ol=wr ol=wr 1. KDðFeMgÞ can also increase through phenocryst accuthis sample shows textural evidence of near-equilibrium crystallization and KDðFeMgÞ mulation as the whole-rock composition (initially at L1) evolves towards a higher XMg (large gray arrow showing effect of accumulation of early formed olivine). En, enstatite; Fo, forsterite; Fs, ferrosilite; Qz, quartz; Pxss, pyroxene solid solution; Olss, olivine solid solution.

fractionation, assimilation, and/or mixing during ascent through, and residence in, the upper mantle and crust. This is the case at Mt. Baker, where even the most primitive lavas show evidence for some differentiation since leaving their mantle source, as none of the whole-rock compositions is in equilibrium with residual mantle olivine. To reconstruct the compositions of the primary basaltic liquids, the major and trace element compositions of the basalts were incrementally adjusted, as necessary, for plagioclase accumulation, olivine fractionation, and oxidation^reduction of iron. For the purpose of this study, we assumed that crustal contamination of the Mt. Baker basalts was sufficiently limited to permit the reconstruction of realistic primary compositions. An assessment of crustal

interactions has been presented by Moore & DeBari (2012), who concluded that the basalts retain their mantlederived geochemical features despite disequilibrium textures. Evidence for mixing and/or assimilation is strongest for Cathedral Crag, as it has the highest phenocryst content and higher 87Sr/86Sr and lower eNd values than the other basalts, similar to the more silicic Mt. Baker lavas (Figs 6a and 7a). Plagioclase is present in super-cotectic proportions; the whole-rock composition plots well within the plagioclase field in the low-pressure basalt system (Longhi, 1991). This is reflected in high Al2O3 (420 wt %) and Sr (41200 ppm) abundances and a positive Eu anomaly. Hand samples also contain crustal xenoliths.

255


XMg (olivine)

0.8


VOLUME 55

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Therefore we assess the petrogenesis of the high-Mg basaltic andesites (HMBA) mainly using Tarn Plateau, but also examine Cathedral Crag because it is the only other basalt in this group. The three Mt. Baker basalt groups are not related by the addition or removal of any realistic mineral phase assemblages (Fig. 5b; Moore & DeBari, 2012), indicating that the three groups inherited their distinct REE patterns from primary mantle-derived liquids.

Reconstruction procedures Adjustments were restricted to the addition of olivine (þ chromian spinel) to all samples and the removal of plagioclase from Cathedral Crag. The super-cotectic abundance of plagioclase coupled with high Al2O3 and Sr abundances and a positive Eu anomaly justify a correction for excess plagioclase in Cathedral Crag. Whole-rock compositions were adjusted through incremental addition of an equilibrium olivine composition, 0·5^1% per step, until reaching equilibrium with Fo90. After each olivine increment, an updated equilibrium olivine composition was ol=liq recalculated assuming a KDðFeMgÞ value of 0·32 (Toplis,

2005) and an average bulk-rock Fe3þ/Fe value of 0·20. Compositions were adjusted further by assuming 2 g chromian spinel was lost for every 100 g olivine lost (based on modal abundances). The quantities of olivine added range from a low of 7 wt % in Lake Shannon 122 to a high of 20 wt % in Hogback. Tarn Plateau was adjusted into equilibrium with Fo92 because its whole-rock composition is similar to that of high-Mg basaltic andesites of the Cascades that commonly contain particularly magnesian olivine phenocrysts (up to Fo94, Baker et al., 1994). The high-Mg orthopyroxene inclusion in Tarn Plateau also opx=ol implies equilibrium with Fo92, assuming that KDðFeMgÞ 1 (e.g. Gaetani & Grove, 1998). Because trace element abundances were not measured in single minerals, whole-rock trace element abundances were adjusted using mineral/melt partition coefficients (D) (Table 2). Addition of olivine, which has low D values for all trace elements (except Ni), results in lower absolute whole-rock abundances, but there is essentially no change in relative abundances. The final reconstructed primary basalt compositions are reported in Table 3 and plotted in Fig. 11.

256


Fig. 10. Backscattered electron images of plagioclase (a^c) and clinopyroxene (d^f) phenocrysts. (a) Sulphur Creek: resorbed albitic core (An45) surrounded by a zone of normally zoned plagioclase (An62^56), a penultimate zone of relatively uniform composition (An65), and a thin albitic rim (An49). (b) Cathedral Crag: phenocryst with normal step-zoning showing compositional gaps: An86 (core), An65 (inner rim), and An40 (outer rim). (c) Sulphur Creek: resorbed anorthitic core (An78) surrounded by a broad uniform zone (An50^53), a penultimate zone (An65), and a thin albitic outer rim (An48). (d) Cathedral Crag: reversely zoned clinopyroxene core (Wo46En47 to Wo44En50) is resorbed and overgrown by Fe-rich rim (Wo38En47) with titanomagnetite inclusions. The small orthopyroxene (Wo4En68) crystal enclosed in the rim should be noted. (e) Tarn Plateau: euhedral crystal with a uniform magnesian core (Wo43En50) and a broad, Fe-rich rim (Wo42En46) with subtle oscillatory zoning. (f) Tarn Plateau: clinopyroxene with a partially resorbed, more Fe-rich core (Wo41En45) and a broad magnesian rim (Wo43En49) with oscillatory zoning. The core has been compositionally modified around resorption channels.

MULLEN & McCALLUM


Table 2: Partition coefficients Dmineral=melt i

Table 3: Mt. Baker basalt whole-rock compositions adjusted into equilibrium with the mantle

cpx

opx

oliv

gar

plag

Cs

0·00021

0·00001*

0·000001*

0·00011

0·0068

Rb

0·00063

0·000453

0·0000453

0·00022

0·0188

Park

Cathedral Hogback Sulphur Tarn

Ba

0·0006811

0·000043

0·0000433

0·000072

0·328

Butte

Crag

Th

0·012010

0·000053

0·000053

0·00212

0·198

U

0·012710

0·000053

0·0000413

0·00944

0·348

Nb

0·00811

0·00103

0·00023

0·003113

wt % Adjusted compositions

1

1

Lake

Lake

Plateau Shannon Shannon 122

114

0·0088

SiO2

49·34

51·80

51·18

50·14

52·40

51·84

6

0·0278

TiO2

0·93

1·06

1·21

1·36

0·83

1·27

1·28

50·18

0·021

0·002*

0·00002

K

0·007211

0·00012

0·000022

0·0132

0·0979

Al2O3

15·41

15·17

15·02

14·55

14·71

15·11

15·55

La

0·05412

0·00061

0·000053

0·001613

0·118

Cr2O3

0·12

0·15

0·17

0·15

0·11

0·10

0·13

Ce

11

0·086

0·0017

1

3

0·085*

Fe2O3

2·14

1·83

2·02

2·09

1·63

1·92

1·99

Pb

0·07211

0·00142

0·0000713

0·00032

0·10858

FeO

7·70

6·60

7·28

7·55

5·87

6·90

7·15

Pr

0·137*

0·0026*

0·00013*

0·029*

0·065*

MnO

0·15

0·10

0·12

0·14

0·12

0·15

0·14

Sr

0·12811

0·0093

0·0002513

0·002513

1·948

MgO

12·42

11·95

11·75

12·27

12·07

11·09

11·53

Nd

0·18711

0·0041

0·000203

0·05213

0·0528

NiO

0·03

0·04

0·05

0·05

0·03

0·04

0·04

12

1

0·25

0·0418

CaO

8·42

6·88

6·92

6·92

8·75

7·06

7·70

2·86

3·46

3·29

3·67

2·58

3·60

3·54 0·54

3

0·005

13

13

Sm

0·380

0·011

0·0006

Zr

0·061

0·0113

0·00077

0·6614

0·00398

Na2O

Hf

0·121

0·0133

0·00117

0·6814

0·00158

K2O

0·38

0·90

0·72

0·74

0·75

0·70

Eu

5

0·45

0·016*

0·0008*

13

0·40

1·428

P2O5

0·13

0·31

0·28

0·36

0·16

0·23

0·24

Ti

0·305

0·0617

0·00217

0·2913

0·0478

Total

100·00 100·00

100·00

100·00

100·00

100·00

100·00

Gd

0·50*

0·022*

0·0013

0·90*

0·0359

0·56*

1

3

Tb

0·030

0·002

0·0319

13

1·4

Dy

0·61*

0·038*

0·004

2·2

0·0269

Y

0·65*

0·0461

0·0073

3·113

0·0269

1

3

13

0·0189

13

Ho Er Tm

12

0·65

0·048

3

13

0·006

2·8

0·69*

0·058*

0·009

3·6

0·01459

0·72*

1

0·013*

3·7*

0·0129

1

3

0·071

3

The basalts are adjusted into equilibrium with Fo90 mantle (except Fo92 for Tarn Plateau and Fo91 for Cathedral Crag; see text) and Fe3þ/Fe ¼ 0·20.

10

Yb

0·74*

0·077

0·017

3·9

0·0097

Lu

0·7512

0·091

0·023

3·810

0·0089

9

Data sources: 1Adam & Green (2006); 2Halliday et al. (1995) compilation; 3Donnelly et al. (2004) compilation; 4 Elkins et al. (2008); 5McDade et al. (2003); 6Green et al. (2000); 7Kennedy et al. (1993); 8Tepley et al. (2010); 9 Claeson & Meurer (2004) compilation; 10Hauri et al. (1994); 11Hart & Dunn (1993); 12Gaetani (2004); 13 Johnson (1998); 14Salters & Longhi (1999). *Interpolated. cpx, clinopyroxene; opx, orthopyroxene; oliv, olivine; gar, garnet; plag, plagioclase.

Comparison with primitive basalt suites of the Cascade arc Several primitive basalt end-members have been recognized in the Cascade arc (Bacon et al., 1997; Conrey et al., 1997; Hildreth, 2007; Schmidt et al., 2008), and the Mt. Baker basalts have been classified within this scheme (Moore & DeBari, 2012). However, the reconstructed primary Mt. Baker liquids differ from their compositions ‘aserupted’, as do many other Cascade arc basalts. To clarify

similarities and differences among the basalt suites, we compared the reconstructed Mt. Baker basalt compositions with end-member Cascade arc basalt compositions that have also been adjusted into equilibrium with mantle olivine (where necessary) (Fig. 11). The two most abundant basalt types in the Cascades are calc-alkaline basalt (CAB), which occurs throughout the arc, and low-K olivine tholeiite (LKOT), most predominant in the central and southern Cascades (Hildreth, 2007; Schmidt et al., 2008). CAB have higher SiO2 and K2O than LKOT, as well as lower Al2O3, CaO, and TiO2. LILE enrichments and HFSE depletions are prominent in CAB but muted or absent in LKOT, which are nearly indistinguishable from N-MORB in some cases (Bacon et al., 1997; Conrey et al., 1997; Leeman et al., 2005). Higher SiO2, yet still primitive variants of the CAB group, known as high-Mg basaltic andesite (HMBA) and highMg andesite (HMA), are most prevalent in the southern part of the arc (Lassen and Shasta) (Hildreth, 2007). CAB have initial water contents of up to 3 wt % (and even higher in the HMBA^HMA variant, up to 5·6 wt %) whereas LKOT are nearly anhydrous (Sisson & Layne, 1993; Ruscitto et al., 2010, 2011, 2012). A third, less common, suite of basalts has compositional affinities to

257


Ta

0·00006

0·019

Creek


20

VOLUME 55

NUMBER 2

(a)


wt% MgO at Fo90

18 16 ALK

14

OIB

LKOT

FEBRUARY 2014

other Garibaldi belt basalts

12

CAB

8 44 20

46

48

50

52

54

56

(b)

wt% Al2O3 at Fo90

18

LKOT

CAB

16 ALK

HMBA

14 OIB

12 calc-alkaline basalt (CAB) high-Mg basaltic andesite (HMBA) low-K olivine tholeiite (LKOT) intraplate-type basalt: tholeiitic (OIB) intraplate-type basalt: alkalic (ALK)

10 8 44

46

48

50

52

54

56

wt% SiO2 at Fo90 Fig. 11. (a) Wt % MgO vs wt % SiO2 and (b) wt % Al2O3 vs wt % SiO2 for Mt. Baker basalts and other primitive basalts from the Cascade arc. All compositions are adjusted into equilibrium with Fo90 mantle (where necessary). Data sources for basalt groups in the Cascade arc are given in Fig. 3. Primitive lavas from the Garibaldi Belt are indicated with small gray circles (data from Green & Sinha, 2005; Green, 2006).

intraplate basalts (IPB), which have been subdivided into hypersthene-normative (or ‘ocean island basalt-like’; ‘OIBlike’) basalts and nepheline-normative alkalic basalts (Leeman et al., 2005). Alkalic basalts occur in the northernmost Garibaldi Belt (Green, 2006; Mullen & Weis, 2013), the Cascades^Columbia transect of northern Oregon and southwestern Washington (Leeman et al., 1990, 2005; Conrey et al., 1997), and north of Mt. Rainier (Reiners

et al., 2000). IPB have lower Al2O3 (Fig. 11b) and higher TiO2 than both LKOTand CAB, but elevated abundances of both LILE and HFSE (Fig. 4b). Based on the reconstructed major and trace element characteristics of the Mt. Baker basalts, Group I (Lake Shannon, Sulphur Creek, Hogback) is most similar to the CAB group, and the higher SiO2 contents of Group II (Tarn Plateau and Cathedral Crag) overlap with the

258


HMBA

10

MULLEN & McCALLUM


HMBA variant (Fig. 11), consistent with the classification of Moore & DeBari (2012). Na2O þ K2O abundances in Mt. Baker CAB are not as high as those in alkalic basalts of the northern Garibaldi Belt (Green, 2006), but Sulphur Creek and Hogback are ‘borderline’ alkalic and have the highest TiO2. HREE are also slightly higher in Mt. Baker CAB than in end-member Cascade CAB. With the lowest K2O at Mt. Baker (Fig. 3a), Park Butte (Group III) is similar to the LKOTgroup. This lava also has a small positive Eu anomaly that is consistent with other Cascades LKOT (Bacon et al., 1997). However, Park Butte has notably lower Al2O3, higher SiO2 (Fig. 11), and higher (La/ Yb)N and LILE than other LKOT (Figs 4b and 5b).

The intensive parameters of pressure (P), temperature (T), water fugacity (fH2O), and oxygen fugacity (fO2) are major variables controlling the compositions of primary mantle melts, and in determining the stability, compositions, and crystallization of mineral phases during transit of magmas from the mantle source region to the surface. Computed temperatures and pressures based on the measured compositions of coexisting minerals and glass are expected to reflect conditions prevailing during transit through, or residence in, the crust. On the other hand, the P and Tdetermined on reconstructed primary basalt compositions should reflect conditions of segregation of melts from their mantle sources.

Water contents Evidence that the Mt. Baker basalts were hydrous is provided by the high anorthite contents of plagioclase and the presence of vapor bubbles in melt inclusions. Shaw (2011) reported 2·3 wt % H2O in Sulphur Creek basaltic tephra by Fourier transform infra-red (FTIR) measurements on olivine melt inclusions, but direct measurement of H2O in the other samples is impractical owing to the microcrystalline nature of the melt inclusions (Fig. 8b) and absence of quenched tephra samples. In principle, plagioclase core compositions can be used as an indicator of pre-eruptive water contents because the presence of dissolved water affects Ca^Na exchange between plagioclase and coexisting liquid (Kudo & Weill, 1970). Lange et al. (2009) quantified this effect in a ‘plagioclase hygrometer’ in which the input parameters are T, P and liquid and plagioclase compositions. We also used the Lange et al. formulation to calculate water contents at plagioclase saturation using the most An-rich plagioclase core composition measured in each sample (excluding antecrysts). We used MELTS (Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998) to determine crystallization sequences for the primary basalt compositions at a wide range of input P and H2O. By trial and error we found P and H2O ranges for which

Sample

Plagioclase

P (MPa)

T (8C) at wt %

composition1 at plag sat plag sat

H2O at

wt % H2O at

(MELTS)

(MELTS) plag sat2 liquidus3

Tarn Plateau

An88

100

1100

3·4

2·7

Cathedral Crag

An88

100

1070

4·5

3·7

Sulphur Creek

An79

100

1155

1·8

1·5

Hogback Camp An82

100

1118

2·5

2·1

Lake Shannon

An82

100

1170

1·7

1·5

Park Butte

An86

100

1140

2·4

2·1

sat, saturation. 1 Maximum An content in plagioclase cores. 2 From Lange et al. (2009) hygrometer. 3 Calculated using % olivine þ chromite crystallized prior to plagioclase saturation (from MELTS).

predicted and observed crystallization sequences were in agreement. Although this procedure does provide a range of possible H2O contents for each sample, the Lange et al. hygrometer provides more precise values for H2O. We therefore used the T, P, and coexisting liquid composition at plagioclase saturation as given by MELTS as input into the Lange et al. hygrometer. At plagioclase saturation, water contents range from 1·7 to 4·5 wt % (Table 4). These values were adjusted for olivine and chromite that crystallized prior to plagioclase saturation to obtain liquidus H2O contents, which range from 1·5 to 3·7 wt % (Table 4), lowest in the CAB (Sulphur Creek), intermediate in the LKOT (Park Butte), and highest in HMBA (Cathedral Crag). Moore & DeBari (2012) calculated H2O contents using a method based on water-saturated experiments at 200 MPa (Sisson & Grove, 1993) and their values are systematically higher than ours (2^4·5 wt %). Our result for Sulphur Creek (1·5 wt %) is slightly lower than that determined by Shaw (2011). Overall, our values are within the range of water contents measured in primitive Cascade arc magmas (Sisson & Layne, 1993; Le Voyer et al., 2010; Ruscitto et al., 2010, 2011). However, LKOTelsewhere in the Cascades are associated with near-anhydrous melting conditions (Sisson & Layne, 1993; Le Voyer et al., 2010), whereas the Park Butte LKOT has a water content between Mt. Baker CAB and HMBA.

Temperatures and redox states Olivine^spinel thermometer In the Mt. Baker basalts, thermometric data are provided by coexisting olivine and Cr-spinel that represent equilibrium compositional pairs. Olivine phenocrysts with

259


CONST R A I N TS ON I N T E NSI V E PA R A M E T E R S

Table 4: H2O contents (wt %) calculated for the Mt. Baker basalts


-6.0

VOLUME 55

Group I Group II Tarn Plateau Ol-Px-Sp Tarn Plateau FeTi oxides Tarn Plateau Sym Cathedral Crag Ol-Px-Sp Cathedral Crag FeTi oxides

-10.0

log fO2

FEBRUARY 2014

QFM QFM+1

Sulphur Creek Ol-Px-Sp

-8.0

NUMBER 2

Group III

-12.0

Park Butte Ol-Px-Sp Park Butte FeTi oxides

-14.0

-18.0 650

Lk Shannon 122 Lk Shannon 114 Sulphur Cr

750

850

950

1050

1100

T (°C) Fig. 12. Variation of log fO2 vs temperature based on olivine^spinel^pyroxene oxybarometry (calculations and calibrations discussed in text) and Fe^Ti oxide oxybarometry [calibration of Ghiorso & Evans (2008)]. Horizontal lines are temperature ranges given by olivine^spinel pairs that have no associated pyroxene for calculating fO2. Continuous and dashed curves represent QFM and QFM þ1 buffer curves, respectively (Frost, 1991). Sym, symplectite.

inclusions of chromian spinel crystallized under near-liquidus conditions, most probably during transit of magmas through, or storage in, the crust. Textural evidence (zoning, skeletal crystals, melt inclusions) indicates that the olivine^chromite pairs in Sulphur Creek and Lake Shannon crystallized rapidly, whereas those in Tarn Plateau, Cathedral Crag, and Park Butte formed in progressively slower cooling regimes. Temperatures calculated using the Sack & Ghiorso (1991) formulation of the olivine^spinel geothermometer are listed in Supplementary Data Electronic Appendix 2 (Table B) and shown in Fig. 12. A pressure of 200 MPa was assumed for the calculations because the basalt compositions plot close to the 200 MPa water-saturated cotectic (liquid þ olivine þ plagioclase þ clinopyroxene) of Sisson & Grove (1993). Temperatures vary over a large range but are lower than liquidus temperatures calculated using MELTS. The highest temperatures are given by Sulphur Creek and Lake Shannon, and the lowest temperatures by Park Butte, consistent with textural evidence for slower cooling in the latter flow.

Olivine þ spinel þ pyroxene oxygen barometer For olivine^spinel^pyroxene assemblages, phase compositions can be used to compute oxygen fugacities using temperatures from olivine^spinel pairs as input. Tarn Plateau and Cathedral Crag contain clinopyroxene phenocrysts coexisting with olivine and chromian spinel. Park Butte

and Sulphur Creek contain groundmass intergrowths of olivine þ clinopyroxene þ spinel. No appropriate assemblages were found in the Hogback or Lake Shannon samples. Oxygen fugacities were calculated via the following reaction: Mg2 Si2 O6 px þ 2=3Fe3 O4 sp ¼ Mg2 SiO4 ðolÞ ð1Þ þ Fe2 SiO4 ðolÞ þ 1=3O2 : High-quality thermochemical data are available for all end-member phases in equation (1) (Berman, 1988; Robie & Hemingway, 1995) and log fO2 is calculated using activity^composition models for each phase: sp 2=3 ol =ðaol ð2Þ log fO2 ¼ 3 log K1 þ log ½apx en ðamt Þ fo afa Þ where [K1 ¼exp(G18/RT)] is the equilibrium constant sp ol ol for reaction (1) and apx en , amt, afo and afa are, respectively, the activities of the enstatite component in pyroxene, the magnetite component in spinel, and the forsterite and fayalite components in olivine. Using input temperatures from sp olivine^chromite thermometry, we computed amt, aol fo and px aol fa using the activity models of Wood (1991) and aen from data provided in MELTS (Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998). Resulting fO2 values (Supplementary Data Electronic Appendix 2, Table B) are shown in Fig. 12 and range from quartz^fayalite^magnetite (QFM) þ 0·7 to þ1·2.

260


-16.0

MULLEN & McCALLUM


Fe^Ti oxide oxygen barometer and thermometer

Pressures and temperatures of mantle melt segregation Liquidus temperatures and pressures of the Mt. Baker reconstructed primary basalts were determined using the olivine^liquid geothermometer coupled with the silica activity

Comparison with experimentally determined phase equilibria Independent constraints on the pressures of melt segregation of Mt. Baker primary basalts are provided by experimental phase equilibria on compositionally comparable systems. In addition to the bulk composition and

Table 5: Intensive parameters and conditions of magma generation for the Mt. Baker basalts Park Butte

Group1 mantle T (8C)2

Cathedral

Hogback

Lk. Shannon

Lk. Shannon

Sulphur

Tarn

Crag

Camp

122

114

Creek

Plateau

III

II

I

I

I

I

II

1326

1274

1309

1307

1326

1354

1273

mantle P (kbar)3

14·0

11·2

% partial melt4

11

10

Mantle residue4

11·8

11·2

13·7

15·1

5

–

7

5

–

Harzburgite:

Harzburgite:

Lherzolite:

Lherzolite:

Lherzolite:

40% opx

50% opx

7% cpx

3% cpx

5% cpx

50% opx

60% oliv

50% oliv

34% opx

43% opx

37% opx

50% oliv

59% oliv % SC added4 wt % H2O at liquidus fO2 (QFM)5

9·0 12

11 2·1 0·2 to 1·2

26 3·7 þ0·0 to 1·2

54% oliv

58% oliv

9

–

8

7

2·1

–

1·5

1·5

–

–

–

1

þ1·0

Harzburgite:

20 2·7 þ0·9 to 1·4

Group I, CAB (calc-alkaline basalt); Group II, HMBA (high-Mg basaltic andesite); Group III: LKOT (low-K olivine tholeiite). 2 Olivine–liquid geothermometer (Putirka, 2008). 3 Melt silica activity barometer (Putirka, 2008). 4 Results of trace element and isotope modeling. ol, olivine; sp, spinel; cpx, clinopyroxene; plag, plagioclase; SC, subduction component. For all samples except Group III (Park Butte), the SC is composed of 84% metabasalt fluid, 7% sediment melt, 10% metabasalt melt. The Group III SC is 90% metabasalt fluid, 3% sediment melt, 7% metabasalt melt (see text). 5 Oxygen fugacities are reported relative to QFM buffer of Frost (1991).

261


Coexisting titanomagnetite and ilmenite microphenocrysts in Tarn Plateau, Cathedral Crag and Park Butte basalts were used to calculate near-solidus redox states and temperatures using the calibration of Ghiorso & Evans (2008), assuming P ¼ 200 MPa. T^fO2 values (Fig. 12; listed in Supplementary Data Electronic Appendix 2, Table E) were calculated only for Fe^Ti oxide pairs in direct contact, and data collected on grains exhibiting oxy-exsolution textures were not used. Temperatures based on Fe^Ti oxides are lower than the temperatures given by olivine^chromite pairs in the same rock, consistent with textural evidence for late crystallization of Fe^Ti oxides. Oxygen fugacities range from QFM 0·2 to þ1·5, overlapping with the fO2 range given by olivine^spinel^pyroxene assemblages. Oxygen fugacities calculated for the three Mt. Baker basalt groups are indistinguishable, as also observed for LKOT and CAB in the central Oregon Cascades (Rowe et al., 2009), but contrasting with other results from the Cascade and Mexican arcs showing that LKOT record more reducing conditions than CAB (Bacon et al., 1997; Righter, 2000).

geobarometer for hydrous basalt compositions (Putirka, 2008). T and P were solved simultaneously by iteration. Input water contents were obtained fromTable 4.The results represent the T^P conditions at which the primary liquids were in equilibrium with Fo90^92 olivine þ orthopyroxene, which we interpret as the conditions of melt segregation from the mantle. Temperatures and pressures are well correlated and range from 12738C at 0·9 GPa (Tarn Plateau) to 13548C at 1·5 GPa (Sulphur Creek) (Table 5, Fig. 13). The standard estimates of error are 438C and 290 MPa (Putirka, 2008). The computed pressures of final equilibration of the basaltic melts with mantle assemblages lie within the depth limits of the Moho and the hot core of the mantle wedge. Our results are broadly consistent with thermobarometric results for other Cascades primitive lavas, which indicate that HMBA represent the lowest T^P conditions of final equilibration with the mantle, typically near the base of the crust, and that CAB and LKOT (except at Mt. Shasta) record progressively higher T^P conditions (e.g. Elkins-Tanton et al., 2001; Leeman et al., 2005; Ruscitto et al., 2010, 2011).


VOLUME 55

1.6


P (GPa)

1.2

0.8 1200

FEBRUARY 2014

Phase equilibria of Mt. Baker basalts

1.4

1.0

NUMBER 2

mantle crust

1250

1300

1350

1400

T (°C)

T R AC E E L E M E N T A N D I S O T O P E C ON ST R A I N T S mineralogy of the mantle source, major element compositions of mantle melts are controlled by the pressures and temperatures of melting and the fugacities of volatile components. Because the basaltic magmas at Mt. Baker were hydrous, data used to construct the phase diagram were obtained primarily from experiments conducted on Cascade arc basalts under water-undersaturated conditions that are representative of melting conditions in the Cascade subarc mantle (Baker et al., 1994; Gaetani & Grove, 1998; Mu«ntener et al., 2001; Parman & Grove, 2004). Difficulties inherent in experimental studies on naturally occurring basalts at high pressure under fluid-undersaturated conditions restrict P^T^X conditions investigated to the immediate vicinity of multi-saturation points. Experimental studies on the anhydrous CMAS system (Presnall et al., 1979; Longhi, 1987; Gudfinnsson & Presnall, 1996) provide a template for interpreting the results of experiments on the compositionally more complex natural systems. The effects of water on the phase equilibria are illustrated in Fig. 14, in which the anhydrous phase equilibria at 0·1MPa (Longhi, 1991) and 1GPa (Longhi, 2002; Wasylenki et al., 2003) are superimposed on the hydrous phase equilibria at 1GPa and 2 GPa. The effects of water include the following: (1) reductions in solidus and liquidus temperatures; (2) stabilization of garnet at the solidus at lower pressures; (3) contraction of the primary crystallization volumes of plagioclase, resulting in more aluminous initial melts; (4) a slight increase in the SiO2 content of liquids (calculated on an anhydrous basis) and major decreases in MgO and FeO; (5) ‘flattening’ of the plagioclase loop such that plagioclase is more anorthitic in water-bearing melts relative to anhydrous melts of similar composition.

Source reservoirs that contribute to primitive arc basalts are the mantle wedge, subducting sediment and oceanic crust, and possibly subducting serpentinized mantle (e.g. Tatsumi & Eggins, 1995; Ulmer & Trommsdorff, 1995). Previous trace element and isotope studies of primitive Cascade arc basalts indicate contributions to the mantle wedge from sediment and altered oceanic crust (Borg et al., 1997; Rose et al., 2001; Grove et al., 2002; Leeman et al., 2004; Green & Sinha, 2005; Jicha et al., 2008; Le Voyer et al., 2010), but consensus has not been reached on their relative inputs or whether the slab-derived components are fluids, melts, or a combination. At Mt. Baker, the Pb isotope compositions of the basalts plot between the fields defined by northern Cascadia subducting sediment and Juan de Fuca MORB (Fig. 6b and c), indicating a depleted source for the Mt. Baker basalts that was modified by a sediment-derived melt and/or fluid. A component derived from altered oceanic crust, however, can be masked in Pb^Pb plots because its isotopic composition is probably similar to that of the mantle wedge. We constrained the inputs from both reservoirs, and from the mantle wedge, through least-squares forward modeling of Pb isotope ratios and 28 trace elements in the Mt. Baker basalts.

Model framework and input parameters In the first step of the model, the compositions of four possible metasomatic agents that could be released from the slab (sediment melt and fluid; metabasalt melt and fluid) are determined by forward modeling, beginning with compositions of northern Cascadia sediment and metamorphosed altered oceanic crust (metabasalt) (Table 6). Fluid

262


Fig. 13. Pressure vs temperature plot illustrating the conditions at which the Mt. Baker basalts segregated from the mantle. P and T were calculated using the silica activity barometer and olivine^liquid geothermometer calibrations of Putirka (2008) and use the liquidus H2O contents fromTable 4. Error bars indicate the standard estimates of error (438C and 0·29 GPa; Putirka, 2008). The crust^mantle boundary equivalent pressure is based on seismic data from Miller et al. (1997).

The reconstructed Mt. Baker primary basalts plot between the 1 and 2 GPa hydrous field boundaries, in agreement with the pressures calculated by thermobarometry. We interpret this result to mean that the basalts segregated from the mantle at these pressures. There is no evidence for equilibration of the Mt. Baker primary basalt magmas with garnet-bearing mantle assemblages at pressures in excess of 2 GPa, a conclusion supported by trace element data (see the next section). The Sulphur Creek, Lake Shannon, and Hogback basalts appear to be close to multi-phase saturation with mantle lherzolite at 1·5 GPa. The Tarn Plateau magma, and possibly Cathedral Crag, appear to have formed by higher degrees of melting, in equilibrium with a harzburgitic residue at a pressure of 1GPa. Park Butte may also have a harzburgite residue, but at a slightly higher pressure of 1·4 GPa.

MULLEN & McCALLUM

Qtz50


[Wo] projection,oxygen units ± cpx silica opx

40

Opx 30

opx ol

pl

20

opx

ol

pl

10

opx pl

Ol

20

60

40

sp

sp 80

Pl

Group I

opx

Sulphur Creek Lk Shannon 122 Lk Shannon 114 Hogback Camp Group II

pl

ol

opx pl

ol sp

sp

Tarn Plateau Cathedral Crag Group III Park Butte

1.2 GPa multisaturation point (hydrous) 2 GPA multisaturation point (hydrous) Field boundaries at 1.2 GPa (hydrous) Field boundaries at 2 GPa (hydrous) Field boundaries at 1 bar, Mg# 0.7 Field boundaries at 1 GPa (anhydrous) Fig. 14. Projection in oxygen units from [Wo] on to the Ol^Pl^Qtz plane illustrating phase relations in the basalt^lherzolite system. Clinopyroxene is a possible additional phase in all assemblages. The Mt. Baker basalts have been adjusted into equilibrium with mantle olivine. Fine black dotted lines are field boundaries at 0·1MPa for liquids with Mg# 0·7 from Longhi (1991). Bold gray dotted lines are field boundaries at P 1·2 GPa in the anhydrous system, based on the BATCH algorithm (Longhi, 2002) and experimental data from Bartels et al. (1991), Kinzler & Grove (1992), Longhi et al. (1999) and Wasylenki et al. (2003). Temperatures decrease in the direction of the arrows. Bold black lines and gray star indicate field boundaries and multisaturation point, respectively, in the hydrous system at P 1·2 GPa based on data from Baker et al. (1994) at 1GPa, Mu«ntener et al. (2001) at 1·2 GPa, Gaetani & Grove (1998) at 1·2 GPa, and Parman & Grove (2004) at 1·5 GPa. Bold black dashed lines and white star indicate field boundaries and multisaturation point, respectively, in the hydrous system at 2 GPa from Gaetani & Grove (1998). In the hydrous system, the Mg# of experimental liquids ranges from 0·71 to 0·54 and the H2O contents of liquids range from 4 to 6 wt %. Data are based primarily on experiments on calc-alkaline basalt (85-44) from Mt. Shasta (see Fig. 3). [Wo] projection: Ol ¼ 2(MgO þ FeO þMnO þ 2Fe2O3)/; Pl ¼ 8(Al2O3 þ Na2O K2O)/; Qz ¼ 2SiO2 (MgO þ FeO þMnO þ 2Fe2O3) 2Al2O3 CaO 10(Na2O þ K2O)/. ¼2SiO2 þ (MgO þ FeO þMnO þ 2Fe2O3) þ 6Al2O3 2CaO 2Na2O 18K2O. Oxide components are calculated as mole fractions.

and melt compositions are calculated using experimentally determined bulk partition coefficients (D values) (Table 6) and mass-balance equations for equilibrium melting^dehydration (Shaw, 1970): Cil=f 1 ¼ iÞ þ D i Ci0 ½Fl=f ð1 D

ð3Þ

where Cil=f is the concentration of a trace element (i) in the melt (l) or fluid (f), Ci0 is the initial concentration of trace element (i) (in the metabasalt or sediment), Fl/f is the frac i is the bulk partition tion of melt or fluid generated, and D coefficient, defined as crystalline assemblage/fluid or crystalline assemblage/melt.

263


ol


VOLUME 55

NUMBER 2

FEBRUARY 2014

Table 6: Model input compositions for mantle, sediment and oceanic crust, and slab D values sediment1

5008C eclogite2

9008C eclogite2

mantle3

Slab partition coefficients Dmb/fluid4

Cs

3·0

Rb

58

Ba

860

Th

5·5

0·0015

0·00132

0·0053

0·011

1·60

0·79

0·13

0·088

0·011

0·016

2·00

0·57

8·4

1·2

1·20

0·16

0·034

0·84

0·48

0·18

0·17

0·0137

8·4

0·057

4·81

0·89

7·0

0·12

1·37

0·93

0·24

2·65

1·21 1·20

1·8

0·30

0·12

0·0047

8·9

3·5

3·5

0·210

32

0·0138

107

0·61

La

20

Ce

40

Pb

12

Pr

5·0

0·26 420 3·4 11

0·26 94

0·74

0·026

4·00

1·34

9·0

0·772

8·2

0·064

4·01

1·35

0·36

0·20

0·0232

0·31

0·043

0·64

1·29

2·1

1·7

0·131

0·1412

3·67

1·48

0·046

0·53

0·51

97

Nd

20

11

10

135

2·61

3·6

103

Zr

2·00

0·069

0·234

327

4·2

0·32

0·028

2·4

Sr

Sm

60

4·0 104

9·80

3·4 104

1212 2·9

0·713

18

0·31

3·26

1·53

0·27

39

0·78

2·41

1·6

7·94

53

0·47

4·1710

1·3711

Hf

3·5

2·9

2·9

0·199

64

0·40

2·67

0·95

Eu

1·1

1·5

1·3

0·107

46

1·0

2·27

1·68

59

2·3

2·21

1·31

Gd

Ti

4406 3·9

5·6

4·7

0·395

82

1·9

1·76

1·65

Tb

0·60

1·0

0·85

0·075

9612

3·612

1·52

1·66

Dy

3·7

6·6

5·9

0·531

113

6·8

1·34

1·68

6·6

1·16

1·69

Y

21

8753

43

8753

798

34

3·55

47

Ho

0·75

1·4

1·3

0·122

13712

1012

1·20

1·69

Er

2·2

4·1

3·9

0·371

167

15

1·14

1·69

Yb

2·1

4·0

4·0

0·401

216

23

1·06

1·68

Lu

0·31

0·64

0·64

0·063

238

29

1·05

1·72

0·706701

0·7025

0·7025

0·7025

0·512513

0·51312

0·51312

0·51312

9·4

9·4

9·4

87

Sr/86Sr

143

Nd/144Nd

eNd

2·4

206

Pb/204Pb

19·067

18·650

18·650

18·610

207

Pb/204Pb

15·645

15·490

15·490

15·505

208

Pb/204Pb

38·856

37·930

37·930

38·020

1

Northern Cascadia sediment trace elements are the average of bulk sediment compositions determined for ODP sites 1027 and 888 (Carpentier et al., 2013); isotope data are the average of all analyses from the same two sites (Carpentier et al., 2010). 2 Average 500–6008C and 9008C altered N-MORB eclogites (Becker et al., 2000) were used as sources of fluid and melt, respectively, released by Juan de Fuca metabasalt. LREE and HREE were interpolated and extrapolated from Sm and Nd. Hf and Ta were calculated from Zr and Nb and the average Zr/Hf and Ta/Nb ratios of altered MORB (Kelley et al., 2003). Eu was calculated such that Eu/Eu* ¼ 1. Isotope ratios were selected from within the arrays defined by Juan de Fuca MORB in Fig. 6. 3 Trace elements from average DM (depleted mantle) of Salters & Stracke (2004). 4 7008C, 4 GPa dataset of Kessel et al. (2005a). 5 10008C, 4 GPa dataset of Kessel et al. (2005a). 6 6508C, 2·5 GPa dataset of Johnson & Plank (1999). 7 9008C, 2·5 GPa dataset P47* of Johnson & Plank (1999) (P47* is the second set of partition coefficients reported for this experiment). 8 From Adam et al. (1997). 9 Average of Klimm et al. (2008). 10 From lowest-temperature (7508C) experiment of Hermann & Rubatto (2009). 11 From 9008C, 3 GPa experiment C1699 of Hermann & Rubatto (2009). 12 Interpolated.

264


16271

Dsed/melt7

0·60

U

K

Dsed/fluid6

0·0090

Nb Ta

Dmb/melt5

MULLEN & McCALLUM


subject to the following constraints: (1) olivine and orthopyroxene are required in the residue (as per phase equilibria constraints); (2) the sum of the weight fractions of the minerals in the residual assemblage is unity; (3) olivine is the most abundant phase in the residue, as in common mantle xenoliths (Pearson et al., 2003); (4) predicted Pb isotope compositions are within 0·25% of measured 208 Pb/204Pb and 206Pb/204Pb and within 0·1% of measured 207 Pb/204Pb; and (5) Fl and Ff in equation (3) are 50·1. The denominator in equation (4) normalizes the concentrations of the elements so that each trace element has an equivalent impact on the solution regardless of its absolute concentration. The full equation solved in the model, and a schematic representation of the steps in the process, are given in Supplementary Data Electronic Appendix 4A and B. Variable model parameters are (1) proportion of SC added to the mantle wedge, (2) partial melt fraction of metasomatized mantle, (3) residual mantle mineral assemblages and modal abundances for melting metasomatized mantle, (4) proportion of subduction component derived from sediment vs metabasalt, and (4) proportion of fluid vs melt in the subduction component. The isotopic and trace element compositions used for the mantle wedge, metabasalt, and sediment are listed in Table 5. For sediment we used the average of bulk compositions determined for hemipelagic and turbiditic sediment at Ocean Drilling Program (ODP) sites 888 and 1027 (Fig. 6) (Carpentier et al., 2010, 2013). For metabasalt we first tested the average altered N-MORB of Kelley et al. (2003) but this composition gave a poor fit to the Mt. Baker basalts because of its very high (U/Th)N ratio (normalized to N-MORB) relative to arc basalts. Fluids or melts derived from this starting material will inherit very

high (U/Th)N because U and Th exhibit similar partitioning behavior (Brenan et al., 1995; Kessel et al., 2005a; Feineman et al., 2007). Because natural eclogites have lower (U/Th)N values that are more similar to those of arc basalts (Becker et al., 2000), we used the average 500^ 6008C and 9008C N-MORB eclogites reported by Becker et al. as the source of fluids and partial melts, respectively, released by metamorphosed altered oceanic crust. Isotope ratios were selected from the arrays defined by Juan de Fuca MORB (Fig. 6). Because the Mt. Baker basalts have similar isotopic compositions, we assumed that the basalts tap a single mantle wedge composition. We used DM (average depleted NMORB mantle) of Salters & Stracke (2004) for the trace element composition of the wedge and selected an isotopic composition lying at the primitive end of the Cascade arc array, within the Juan de Fuca MORB field (Fig. 6). Slab fluid and melt partitioning data for a sufficiently large number of elements for our modeling purposes have been reported by Johnson & Plank (1999), Kessel et al. (2005a), Spandler et al. (2007) and Hermann & Rubatto (2009). Of these studies, only Kessel et al. (2005a) reported D values for metabasalt, so we used their 4 GPa, 7008C and 10008C datasets for all metabasalt dehydration and melting calculations, respectively. For sediment melting and dehydration, we tested all possible combinations of the available datasets to determine which best reproduce [i.e. lowest residuals, equation (4)] the Mt. Baker basalts (Supplementary Data Electronic Appendix 4C). As discussed below, the best-fit solutions are provided by the 6508C (sediment dehydration) and 9008C (sediment melting) datasets of Johnson & Plank (1999).

Model limitations and sources of error With large discrepancies (up to several orders of magnitude) among published D values for melting and dehydration of subducting materials, uncertainties in these values are the largest source of error in our trace element model (see the next section). These inconsistencies may be related to differences in experimental design, quenching methods, measurement protocols, and starting material compositions. Other sources of error in our model are uncertainties in the compositions and mineral assemblages of slab materials under the conditions at which dehydration and/or melting take place. As discussed above, the composition of subducting oceanic crust can be approximated with fresh MORB, altered oceanic crust, or eclogite, and there are large discrepancies in the abundance of some trace elements among these rock types (Becker et al., 2000; Kelley et al., 2003). Similarly, progressive metamorphism may result in changes to the composition of subducting sediment. The isotopic and trace element composition of modern Cascadia sediment (Carpentier et al., 2010, 2013) is virtually identical to the composition of Darrington phyllite (Mullen, 2011), an Early Cretaceous blueschist-

265


In the second step of the model, calculated slab fluid and melt compositions are mixed in variable proportions to form a total subduction component (SC). In the third step, a metasomatized mantle composition is calculated by mixing the SC (in variable proportions) with the mantle wedge. In the fourth step, hypothetical primary basalt compositions are calculated from metasomatized mantle using equation (3) for equilibrium melting. Bulk partition coefficients for mantle melting are calculated i values from Table 2. using mineral/melt D Hypothetical basalt compositions are then compared with the Mt. Baker basalts (adjusted into mantle equilibrium) and best-fit solutions are obtained using the Generalized Reduced Gradient (GRG2) nonlinear optimization code provided in Microsoft Excel Solver. Model parameters are varied such as to minimize the sum of squares of residuals (calculated from the differences between the calculated and observed trace element abundances): X liq X 2 liq liq ð4Þ ½Ci ðcalcÞCi ðobsÞ=Ci ðobsÞ r2 ¼ i


VOLUME 55

grade metasediment in the Northwest Cascades System that was subducted in Mesozoic times (Brown et al., 1982), lending confidence to the composition used in the model for the sediment end-member. To simplify the modeling as much as possible, we did not attempt to include the effects of interaction between the SC and the mantle wedge, such as zone refining or chromatographic effects (Navon & Stolper, 1986; Hawkesworth et al., 1993; Ayers, 1998), which could affect the composition of the SC. We have also not included possible fluid input from the subducting mantle.

The dependence of model outcomes on various input parameters is illustrated in sequential modeling scenarios in Supplementary Data Electronic Appendix 4C for the example of Tarn Plateau. In each of four scenarios, one or more constraints are applied or removed from the model to demonstrate the effect on the outcome. In each scenario, all possible pairings of slab D value datasets are shown and the three best-fit solutions are highlighted. Model results depend most heavily on the % SC, the Pb isotope composition of the basalts, and the choice of partition coefficients for slab melting and dehydration. Scenario 1 requires that (1) the Pb isotope ratios of the basalts are reproduced well by the model (see above) and (2) the % SC is limited to a maximum of 5% [low SC inputs of 2^4% for Mt. Baker were proposed by Moore & DeBari (2012)]. Regardless of which combination of partition coefficient datasets is utilized, none of the Scenario 1 outcomes provides an acceptable fit to the trace element abundances in the Tarn Plateau basalt (residuals all 41). Allowing garnet as part of the mantle residue (Scenario 2) slightly reduces the residuals, but none of the solutions provides a good fit to LILE or LREE abundances in the actual basalt because of the Pb isotope and % SC constraints. In Scenario 3, the Pb isotope constraint is removed, resulting in significant reduction of residuals. However, all solutions are associated with relatively high sediment input, leading to predicted Pb isotope ratios that are identical to the subducting sediment end-member (Fig. 6a and b), which is much more radiogenic than the actual basalts. Therefore, the Scenario 3 solutions are not our preferred solutions. Removing the constraint of 55% SC from the model (Scenario 4) appears to be critical for obtaining solutions that most reasonably match both the trace element abundances and Pb isotope ratios of the actual basalts. Other than the partition coefficient datasets, % SC exerts the largest control on model outcomes. Scenario 4 is associated with the lowest residuals observed. All possible pairings of partition coefficient datasets result in a relatively consistent range of values for % partial melt for Tarn Plateau (9^ 14%). In addition, all solutions predict residual harzburgite and the absence of residual garnet, in agreement with

FEBRUARY 2014

phase equilibria. The % SC is notably higher than 5%, ranging from 14 to 29% depending on the partition coefficient datasets used. All partition coefficient pairings also indicate that the SC contains both fluid and melt derived from metabasalt, but the sediment-derived component is variable (melt, fluid, or both). The SC consistently contains a significantly smaller contribution from sediment (1^25%) than from metabasalt. Of all Scenario 4 solutions, the best fit is provided by the 6508C (dehydration) and 9008C (melting) datasets of Johnson & Plank (1999). The best-fit results indicate that Tarn Plateau is a 12% partial melt of mantle metasomatized by an SC consisting of three components: sediment melt, metabasalt melt, and metabasalt fluid. The % SC is 20%. Because Scenario 4 is associated with the lowest residuals and the predicted Pb isotope ratios closely match the actual basalts, we consider the Scenario 4 solutions as our preferred best fit and use this scenario to model all of the Mt. Baker basalts.

Best-fit model results Best-fit model results are given in Table 5 and shown in Figs 15^17. The mantle source of Groups I and II (CAB, HMBA) was metasomatized by a subduction component (SC) consisting of 84% metabasalt fluid, 7% sediment melt, and 10% metabasalt melt. The % SC added to the depleted mantle is lowest in the CAB and highest in the HBMA, ranging from 7% (Sulphur Creek) to 26% (Cathedral Crag) (Figs 15 and 16). Mantle melt fractions generally correlate with degrees of metasomatism and range from 5% in the CAB (Sulphur Creek) to 12% in the HMBA and LKOT (Tarn Plateau and Park Butte). The Group III LKOT (Park Butte) can also be reproduced fairly well with the same SC, but a better fit is obtained with a slightly different SC (90% metabasalt fluid, 3% sediment melt, and 7% metabasalt melt) at 11% partial melt and 11% SC input (Figs 15c, 16c and 17b inset). For all of the basalts, the Pb concentration in the total SC is high relative to the depleted mantle, meaning that the metasomatized mantle acquires a Pb isotopic composition nearly identical to that of the total SC at 5%þ SC addition, resulting in uniform isotopic ratios for all of the Mt. Baker basalts as observed (Fig. 17a). Mantle melting residues (Figs 15 and 16) are olivine þ orthopyroxene for the HMBA and LKOT, and olivine þ orthopyroxene þ clinopyroxene ( spinel) for the CAB, consistent with the conclusions of the phase equilibria analysis. The absence of clinopyroxene from the HMBA and LKOT mantle residues is consistent with larger partial melt fractions calculated for these samples; mantle melting reactions predict a reduction in clinopyroxene with progressive melting (Walter, 2003).

Discussion of best-fit results Cs and Rb abundances predicted by the best-fit solutions (in all four modeling scenarios tested; see Supplementary

266


Influence of input parameters on model outcomes

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MULLEN & McCALLUM

1000 actual basalt best-fit model metasomatized mantle DM

Rock / N-MORB

100

10


Tarn Plateau

Cathedral Crag

Group II 12% partial melt 20% SC 50 ol / 50 opx ∑r2=0.78


1

0.1

ParkButte Butte Park

Lake Shannon 114

10

1

0.1

(c)

0.01 1000

Sulphur Creek

Hogback Camp

Group I 5% partial melt 7% SC 58 ol / 37 opx / 5 cpx ∑r2=0.20

100

Rock / N-MORB

(d)

10


1

0.1

(e)

(f)

0.01 Cs Ba U Ta La Pb Sr Sm Hf Ti Tb Y Er Yb Cs Ba U Ta La Pb Sr Sm Hf Ti Tb Y Er Yb Rb Th Nb K Ce Pr Nd Zr Eu Gd Dy Ho Tm Lu Rb Th Nb K Ce Pr Nd Zr Eu Gd Dy Ho Tm Lu 1000

Rock / N-MORB

100 sediment fluid sediment melt metabasalt fluid metabasalt melt

10 1

Total SC (Groups I & II) Total SC (Group III)

0.1

SN94 MG91

0.01

(g)

0.001 Cs Rb Ba Th U Nb Ta K La Ce Pb Pr Sr Nd Sm Zr Hf Eu Ti Gd Tb Dy Y Ho Er Tm Yb Lu

Fig. 15. N-MORB normalized (Sun & McDonough,1989) trace element diagrams for (a^f) mantle-adjusted Mt. Baker basalts (bold black lines with symbols). Best-fit model solutions are dashed lines and metasomatized mantle source compositions are dotted lines. The bold gray line in each panel is identical and is the initial DM mantle source composition used in the model (Table 6). r2 values are the sum of squares of residuals for all elements shown except Cs and Rb. For all samples except Park Butte (c), the best-fit solutions are based on a total SC composed of 84% metabasalt fluid, 7% sediment melt, and 10% metabasalt melt. The Park Butte total SC is 90% metabasalt fluid, 3% sediment melt, 7% metabasalt melt. (g) Total subduction component (SC) composition calculated for Mt. Baker, and the melts and fluids from sediment and metabasalt that form the total SC. Also plotted are the subduction components of McCulloch & Gamble (1991) (MG91) and Stolper & Newman (1994) (SN94).

267



Group III 11% partial melt 11% SC 60 ol / 40 opx ∑r2=0.62

100

Rock / N-MORB

(b)

(a)

0.01 1000


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NUMBER 2

FEBRUARY 2014

100

REE/chondrite

actual basalt best-fit model

Cathedral Crag



(b)

(a)

100

Park Butte

Lake Shannon 114 Group I 7% partial melt 8% SC 54 ol / 43 opx / 3 cpx ∑r2=0.04

REE/chondrite

Group III 11% partial melt 11% SC 60 ol / 40 opx ∑r2=0.08

(c)

(d)

10 100

Sulphur Creek

Hogback Camp


REE/chondrite


(e)

(f)

10 La Ce Pr Nd PmSm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd PmSm Eu Gd Tb Dy Ho Er Tm Yb Lu Fig. 16. REE diagrams for mantle-adjusted Mt. Baker basalts (bold black lines with symbols). Best-fit model solutions are dashed black lines. r2 values are the sum of squares of residuals for all elements shown. Best-fit solutions are based on the same total SC for all samples except Park Butte (see Fig. 15 caption).

Data Electronic Appendix 4C) are invariably higher than observed; Sr abundances are consistently slightly lower than observed (Figs 15 and 16). High predicted Cs and Rb values are most probably related to the metabasalt melting and dehydration partition coefficients used for Cs and Rb, which were obtained from experiments in which the starting material was a K-free MORB doped with trace

amounts of 26 elements (Kessel et al., 2005a). The absence of K prevented the crystallization of phengite, the major host of alkalis, resulting in Dslab=fluid values that are too Cs,Rb low (Hermann & Rubatto, 2009). Phengite is a stable high-pressure phase in subducted metasediment (Spandler et al., 2007; Hermann & Rubatto, 2009) as well as metabasalt (Pawley & Holloway, 1993; Kerrick & Connolly, 2001).

268


10

Tarn Plateau

MULLEN & McCALLUM

(a)


total SC


model sediment

207

Pb/204Pb

15.65

model mantle wedge model oceanic crust addition of sediment and oceanic crust to form total SC addition of SC to mantle wedge to form metasomatized mantle

15.60

Mt. Baker basalts

15.55

N. Juan de Fuca MORB

18.50

18.70

18.90 206

19.10

19.30

204

Pb/ Pb

(b) mb fluid

1400

.999 .99

1200

.98 .97 .96

Ba/Th

1000

.95

.995

800

.9

sed fluid

.6 .5 .3

partial melts of DM+total SC same as above, but with different SC (for Park Butte)

.05 .2

400 200

.2 .1

DM mb melt

.90

0 0.00

.7 .9 .8

.3 .6

.5

400

.4

300

sed bulk 200

Mt. Baker basalts 0.50

1.00

Pb/Ce 206

500

.2

sed melt

.6 .5 .6 .7

.1

.98

.4 .5 .7

sed melt 2

.1

.3

204

.7

.99

600

207

.8


204

1.50

total SC

100

DM mb melt 2.00 0 0.00 0.10

0.20

Fig. 17. (a) Pb/ Pb vs Pb/ Pb diagram illustrating the best-fit solutions for the Mt. Baker basalts. Gray field is for all Mt. Baker basalt data (from Fig. 6c). Compositions used in modeling (Table 6): white square, depleted mantle beneath Mt. Baker; gray circle, Juan de Fuca oceanic crust; gray square, bulk northern Cascadia subducting sediment. White star, total Mt. Baker subduction component. Gray arrows indicate formation of the total subduction component (SC) by mixing fluid melt from Juan de Fuca metabasalt with fluid melt from sediment. Black arrow indicates addition of the total subduction component to the depleted mantle; tick marks indicate intervals of 2% metasomatism of the mantle. Data for northern Juan de Fuca MORB and northern Cascadia sediment are from Fig. 6c. Because Pb is high in the total SC relative to the mantle, the metasomatized mantle wedge (and partial melts thereof) will acquire the Pb isotope ratios of the total SC at 10% addition and will plot at the white star (the isotopic signature of the total SC). (b) Ba/Th vs Pb/Ce plot showing how the trace element model (discussed in the text) can account for the Mt. Baker basalts with a total subduction component composed predominantly of metabasalt fluid plus small amounts of metabasalt melt and sediment melt, but no sediment fluid. Black curves are mixing curves for slab-derived fluids and melts; mixing proportions are indicated by tick marks. Slab melt and fluid compositions are calculated using the sediment and metabasalt compositions and partition coefficient datasets in Table 6 (details in text). Gray field is for all Mt. Baker basalt data (single data points shown in inset). The black arrow (inset) shows the compositional path of melts formed by melting depleted mantle metasomatized by variable amounts of the total SC (white star). The black dotted arrow (inset) provides a better fit for Park Butte using an SC with a slightly higher proportion of metabasalt fluid (see text for discussion). mb, metabasalt; sed, sediment; DM, depleted mantle.

269


15.50


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FEBRUARY 2014

the past 14 Myr. Seismic reflection profiles also reveal that little to no sediment subducts outboard of northern Washington (Davis & Hyndman, 1989).

Comparison with other subduction component estimates Our SC composition is broadly consistent with estimates for other arcs, lying between those of McCulloch & Gamble (1991) and Stolper & Newman (1994) (Fig. 15g). However, our estimate of 7^26% SC is significantly higher than determined by some previous studies of arc basalts in arcs globally, which have concluded that subduction components are typically added to the mantle in quantities of 55% (e.g. McCulloch & Gamble, 1991; Stolper & Newman, 1994; Tatsumi, 2000; Portnyagin et al., 2007; Johnson et al., 2009; Moore & DeBari, 2012; Ruscitto et al., 2012). Stolper & Newman (1994) pioneered a method based upon a mass-balance approach that relies upon preeruptive water contents measured in natural glasses or in olivine-hosted melt inclusions, whereas our estimates are based upon experimentally determined trace element disand Drock=melt , where rock tribution coefficients Drock=fluid i i refers to metabasalt or sediment. Fluid compositions calculated by Stolper & Newman (1994) are highly concentrated (Fig. 15g), containing 450% dissolved components with values much lower than those meacalculated Drock=fluid i sured experimentally. Studies based upon pre-eruptive water contents also do not account for volatiles that may have been lost from the melt prior to entrapment in olivine (Me¤trich & Wallace, 2008), or for water not retained in the mantle wedge. In our model, mantle metasomatism is dominated by relatively dilute aqueous fluids containing 480 wt % H2O (Kessel et al., 2005b), and requires large fluxes to generate the observed high incompatible element abundances. The relatively high SC estimates in our preferred model are consistent with those of Ayers (1998), who used experimentally determined peridotite/fluid D values to show that 17% SC input to the mantle wedge is required to generate an average island arc basalt. It is also noteworthy that the trace element range for our calculated metasomatized mantle at Mt. Baker lies within the range of natural metasomatized mantle xenoliths (Pearson et al., 2003). Ratios of highly incompatible elements in arc magmas, such as Ba/Th, Sr/Th, Pb/Ce, and U/Th, are commonly used as indicators of slab-derived aqueous fluid (Labanieh et al., 2012, and references therein). The best-fit model solutions for Mt. Baker predict that the SC is composed mainly of metabasalt fluid, and based on the partition coefficients utilized in our model, metabasalt fluid does have higher Ba/Th and Pb/Ce ratios than the basalts (as does sediment fluid) (Fig. 17b). However, the dilute nature of the metabasalt fluid (Fig. 15g) means that large quantities can be added to the mantle wedge and still fit the Ba/ Th and Pb/Ce of the basalts. Sediment fluid, which is not

270


Most natural eclogites and blueschists contain significant quantities of phengite (Becker et al., 2000; Li et al., 2008). The relatively poor fit for Sr may be related to the relatively high Dslab=fluid value of 2·9 for Sr reported by Kessel Sr slab=fluid et al. (2005a). Brenan et al. (1995) reported a DSr value of 0·19 at 2 GPa and 9008C. As shown in Fig. 6a, the model predicts higher 87Sr/86Sr values than given by the actual Mt. Baker basalts unless a lower value (0·6) is slab=fluid selected for DSr . The HFSE depletions exhibited by the Mt. Baker basalts do not require residual slab rutile in excess of that present in experimental results for oceanic crust dehydration (0·15% rutile) and melting (0·09% rutile) (Kessel et al., 2005a), nor residual mantle amphibole or rutile (e.g. Borg et al., 1997). The HFSE depletions are explainable as a consequence of ‘lack of addition’ by the subduction component (e.g. Thirlwall et al., 1994). The absence of garnet from residual assemblages contrasts with the results of Moore & DeBari (2012), who proposed residual garnet to explain the REE patterns of the Mt. Baker HMBA. However, a lack of residual garnet is consistent with the pressure ranges obtained from geobarometric calculations and phase equilibria, which indicate a maximum pressure of 1·5 GPa. Garnet is not stable at the solidus of hydrous mantle at pressures less than 1·6 GPa (Gaetani & Grove, 1998). In addition, Dy/Yb values for the HMBA are no higher than in the Mt. Baker LKOT or CAB; all three groups have similar Dy/Yb values (1·9^2·2) that fall within the MORB range, implying no residual garnet (Davidson et al., 2013). For spinel, partition coefficients are not available for many trace elements; therefore, spinel was assumed ‘inert’ for the purposes of trace element modeling and its presence or absence in the residual lherzolite could not be evaluated. However, depleted mantle xenoliths contain up to 2·5% spinel (Pearson et al., 2003) and BATCH algorithm computations (Longhi, 2002) predict that basalts formed at the highest melt fractions are in equilibrium with residual spinel lherzolite. Based on our modeling results, the total sediment contribution to the mantle wedge ranges from 0·5% (Sulphur Creek) to 1·8% (Cathedral Crag), much less than the total metabasalt contribution (6^25%). Such a low sediment input is consistent with the Pb isotope compositions of the Mt. Baker basalts, which limit the amount of sediment input because of the relatively high Pb concentration in sediment. Less sediment input to the mantle beneath Mt. Baker than the rest of the arc may explain why the Mt. Baker basalts have Pb isotope ratios more similar to MORB than many other Cascade arc basalts (Fig. 6a and b). Low sediment input is also consistent with the calculations of Batt et al. (2001) that indicate that at least 80% of the incoming sediment at the trench has been transferred to the accretionary prism in the Olympic Mountains over

NUMBER 2

MULLEN & McCALLUM


predicted by the model, would result in higher Pb/Ce than observed at Mt. Baker (Fig. 17b).

C ON D I T I ON S OF M AG M A G E N E R AT I O N AT M T. B A K E R Water contents and phase stabilities in the subducting crust

4

MORB/gabbro protolith

ab of m

Mt. Baker 85-100 km 3 above slab surface

bo

tto

subsolidus dehydration am out

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Fig. 18. P^Tdiagrams for (a) a basalt^gabbro protolith and (b) a psammopelitic composition showing the changing mineral assemblages. (a) Gray lines indicate metamorphic facies boundaries from Hacker et al. (2003a). P^T profiles for the top and bottom of a 7 km thick subducting slab in northern Cascadia (Syracuse et al., 2010) bound the gray shaded area. White star indicates P and T of first contact between slab and wedge. Dashed black line is the H2O saturated basalt^gabbro solidus (Lambert & Wyllie, 1972; Liu et al., 1996). Continuous black line labeled zoisite out is the dehydration solidus of zoisiteâmphibole eclogite and dash^double-dot line labeled am out is the dehydration solidus of amphibole eclogite for basalt with 0·3 wt % H2O (after Vielzeuf & Schmidt, 2001). Black curve labeled phen out is the dehydration solidus of phengite-bearing coesite eclogite (Hermann & Rubatto, 2009). Dotted black lines indicate subsolidus dehydration of zoisiteâmphibole eclogite and zoisite eclogite. Numbers and circles on the ‘top-of-slab’ curve indicate the remaining wt % H2O in subducting MORB in a hot subduction zone with an initial H2O content of 3 wt % (Pawley & Holloway, 1993; Kerrick & Connolly, 2001). (b) The P^T path for the subducting slab is the same as in (a). Stable assemblages are based on the experimental results of Hermann & Spandler (2008); the stability limit of biotite is not well constrained. Dashed black curve is the H2O-saturated solidus for pelitic compositions (Hermann & Rubatto, 2009). Continuous black curve labeled phen out is the dehydration solidus of phengite-bearing coesite eclogite (Hermann & Rubatto, 2009), as in (a). Numbers indicate wt % H2O remaining in a subducting psammopelitic sediment, assuming an initial H2O content of 5 wt % (Li et al., 2008). grt, garnet; cpx, clinopyroxene; coe, coesite; ph, phengite; cld, chloritoid; am, amphibole; qtz, quartz; bi, biotite; kfs, K-feldspar.

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Thermal models for northern Cascadia, coupled with petrological phase stabilities for subducting sediment, basalt, and mantle peridotite, provide constraints on the flux of fluids and melts beneath Mt. Baker. Although the subducting slab in northern Cascadia is young and hot, partial melting of the slab is likely to be limited because of extensive loss of subducted water at relatively shallow depths to the accretionary prism and the forearc mantle wedge. At the trench outboard of the Olympic Mountains, the cover of hemipelagic and turbiditic sediment is 3^5 km thick (Davis & Hyndman, 1989). However, at least 80% and perhaps as much as 100% of the sediment is added to the toe of the accretionary wedge and/or underplated to the base of the wedge (Batt et al., 2001). Sediment initially contains 40^60 wt % H2O in hydrous minerals and pore

spaces (Davis et al., 1990; Plank & Langmuir, 1998). Upon reaching the corner of the mantle wedge, 4 wt % H2O is retained in the remaining subducted sediment, mainly sequestered in white mica and chlorite (Li et al., 2008). The upper 500 m of subducting basaltic^gabbroic crust is extensively altered and hydrated prior to subduction, with 6^9 wt % H2O present in pore spaces, as adsorbed water, and as structural water in hydrous minerals (Staudigel et al., 1995). As with subducting sediment, most of the water is released to the accretionary wedge. In northern Cascadia, 3 wt % H2O remains in the metabasalt in the slab at 4008C, the temperature at which the slab surface first makes contact with the corner of the mantle wedge (Kerrick & Connolly, 2001; Hyndman & Peacock, 2003). At this point, the subducting oceanic crust has been metamorphosed to chloriteêpidote blueschist (Hacker et al., 2003a). At deeper levels in the subducted crust and underlying mantle, hydrated zones are likely to be less uniform and concentrated along fractures and pseudofaults (Hacker et al., 2003a). Phase stabilities and water contents in the slab after contact with the mantle wedge are illustrated in Fig. 18 for MORB (Vielzeuf & Schmidt, 2001) and psammopelitic

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Phase stabilities in the mantle wedge The effect of water on mantle wedge phase stabilities is shown in Fig. 19, a cross-section along a SW^NE transect

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through the Mt. Baker stratocone. Hydration zones in the forearc mantle and subducted mantle are based on the phase stabilities shown in Fig. 20. At temperatures up to 6808C and depths of up to 70 km at the slab^wedge interface, water released from the subducting slab produces serpentine in the forearc mantle that is clearly seen in magnetic anomalies (Blakely et al., 2005) and seismic velocities (Brocher et al., 2003). Given that the mantle wedge is largely decoupled from the slab between 40 and 60 km (Syracuse et al., 2010), the static corner (‘cold nose’) of the wedge would be continuously hydrated and reach a steady state with a maximum degree of serpentinization [60% according to Hyndman & Peacock (2003)]. Approximately 10% of the serpentinized zone lies within the region of slab^wedge coupling and could be down-dragged as subduction proceeded. Farther down-dip, chlorite replaces serpentine and is stable to 7808C and 80 km at the slab^wedge interface. In the chlorite stability field, the Al2O3 content of depleted mantle (3 wt %) constrains the chlorite content to a Mount Baker

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distance from trench (km) Fig. 19. Cross-section along a SW^NE transect through Mt. Baker. Isotherms (in 8C) are from the Cascadia models of Kirby & Wang (2002), Currie et al. (2004), Wada & Wang (2009) and Syracuse et al. (2010). The Moho beneath Mt. Baker is at a depth of 40 km and the depth to the top of the slab is 95 km. The temperature at the Moho is 8008C (hotter along the magma conduit) and the intersections of isotherms with the top and bottom of the subducting slab are consistent with the model of Syracuse et al. (2010) for the northern Cascade arc. This model indicates a maximum temperature beneath Mt. Baker of 13208C at a depth of 82 km. Hydration zones in the forearc mantle and subducted mantle are based on phase stabilities shown in Fig. 20. SP (diagonally shaded field), serpentine peridotite; CP (dotted field), chlorite peridotite; AP (gray field), amphibole peridotite; GBS, greenschist^blueschist; HE, hydrous eclogite; AE, anhydrous eclogite. Mantle flow paths are shown as bold dashed curves with arrows. Dehydration melting of amphibole peridotite occurs where the mantle flow path crosses the amphibole dehydration curve, at P 2·8 GPa and T 10208C. Mantle closer to the slab^wedge interface than 3 km does not melt (see text for explanation).

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protoliths (Hermann & Spandler, 2008). The P^T pathway shown for the subducting slab is the Cascadia thermal model of Syracuse et al. (2010) based on 10·6 Myr old subducting oceanic crust, the same age as the incoming slab beneath Mt. Baker (Wilson, 2002). Hydrous fluids are released continuously from the subducting oceanic crust with a spike in the flux rate at the transition from blueschist to eclogite at a depth of 50 km, as chlorite is replaced by garnet (Hacker et al., 2003a). At 6808C and 2·4 GPa, the H2O-saturated solidus is crossed (Fig. 18a) but melting is likely to be limited because the water content is low enough (0·75%) to be sequestered in zoisite, amphibole, and phengite, although a small amount of water may be present in fluid inclusions and in fluids infiltrating from a deeper crustal level. At 7658C, 2·5 GPa (83 km), the zoisite breakdown curve is crossed, triggering dehydration melting. The amount of melt produced will be small, as melting is limited by the abundance of zoisite [512% according to Li et al. (2008)]. A second dehydration melting event occurs at the amphibole breakdown curve, which is crossed at 8158C, 2·6 GPa (86 km). Again, the melt fraction is expected to be small, as the amphibole content of the eclogite is 510% (Li et al., 2008). Beyond this point, any remaining water is sequestered in phengite, the abundance of which is limited by the K content of the eclogite. A maximum phengite content of 5 wt % is possible if all K2O in the altered oceanic crust [0·6 wt % according to Kelley et al. (2003)] is stored in phengite. The phengite abundance is likely to be lower, as some K2O would be lost during metamorphic dehydration reactions (Becker et al., 2000). Release of water from subducting sediment is continuousto a depth of 79 km (2·35 GPa, 6808C), at which point fluid loss diminishes because any remaining H2O (1·2 wt %) is sequestered largely in phengite (Li et al., 2008). Melting at the H2O-saturated solidus is limited by the amount of unbound water.The slab must subduct to 2·9 GPa (96 km), 9108C (Fig. 18b) before dehydration melting of phengite occurs (Hermann & Spandler,2008).This depth is essentially coincident withthe arc axis, but is subjecttolarge uncertainty because the slopes of the P^T path and the dehydration curve of phengite are similar (Fig.18b). Release of hydrous fluids from the slab is essentially complete before the slab reaches sub-arc depths. The only slab-derived component that could be released directly beneath Mt. Baker is slab melt. However, trace element modeling results indicate that fluid is the main metasomatic agent of the mantle source of the Mt. Baker basalts. We conclude that mantle hydrated by fluids at shallower depths must be down-dragged to sub-arc depths prior to melting.

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T (°C) Fig. 20. P^T phase relations for lherzolite^H2O. Curves 1, 2, 3 and 4 from Fumagalli & Poli (2005). Curves 5 and 6 are upper stability limit of amphibole in pyrolite with 0·6 wt % H2O (Niida & Green, 1999). Curve 7 is dehydration solidus of amphibole pyrolite from Niida & Green (1999). Contours in the amphibole field indicate modal abundances of amphibole. Curve A, P^T path at the slab^ mantle wedge interface for the northern Cascade arc from Syracuse et al. (2010); curve B, P and T along mantle flow path at 3 km above the slab^wedge interface. White circle labeled M indicates amphibole-out point on Curve B. Curve G is vapor-saturated peridotite solidus from Green et al. (2008, 2010). Abbreviations as in Figs 18 and 19, plus ol, olivine; opx, orthopyroxene; sp, spinel; chl, chlorite; atg, antigorite.

maximum of 20%, corresponding to a bulk water content of 2·5 wt %. At depths of 80^100 km, amphibole is stable and peridotite with 30% amphibole has a bulk H2O content of 0·6 wt % (Niida & Green, 1999). Regions of serpentine and chlorite stability in the subducting mantle are shown in Fig. 19, but the occurrence of these minerals is most probably restricted to fracture zones along which seawater migrated. Assuming that metabasalt and metasediment in hot subducting slabs release 3 wt % fluid containing 80% H2O (Kessel et al., 2005b) over the depth interval from 60 to 100 km, mass-balance calculations based on water contents show that dehydration of 100 g of slab is capable of generating 100 g of chlorite peridotite containing 20 wt % chlorite or 400 g of amphibole peridotite containing 30% amphibole. The total amount of hydrated peridotite depends on the thickness of the slab that is dehydrating.

Mantle melting models Melting models for hydrated peridotite depend critically on knowledge of the P^T conditions of the vapor-saturated peridotite solidus and the stable subsolidus mineral assemblage. Experiments to define the solidus began with Kushiro et al. (1968) but the solidus P and T remains unresolved, as experiments reported by Till

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et al. (2012a) place the vapor-saturated solidus at 810 108C between 3 and 6 GPa, which is 200^5008C lower than most other experimentally determined solidi (e.g. Green et al., 2010). For details on the debate, the reader is directed to comments on Till et al. (2012a) by Green et al. (2012) and Stalder (2012) and replies by Till et al. (2012b, 2012c). Until this important issue is resolved, we use the Green et al. (2010) solidus, as these researchers presented clear evidence that melting has occurred in the supersolidus region investigated in their experiments. As shown in Fig. 20, the P^T pathway for the slab^ wedge interface in the northern Cascades (Syracuse et al., 2010) does not intersect the vapor-saturated solidus of Green et al. (2010). However, a profile 3 km above the interface encounters the dehydration solidus of amphibole peridotite at P 2·8 GPa and T 10208C (Fig. 20). The abundance of amphibole depends on the bulk composition and water content of the peridotite. In the example discussed by Niida & Green (1999), 10% pargasitic amphibole remains at the solidus of a peridotite that initially contained 0·6 wt % water (peridotite with 30% amphibole). Some melting may take place between the vaporsaturated solidus at T 10008C and the dehydration solidus in response to water released as the abundance of pargasite progressively decreases. At the dehydration solidus, amphibole releases its water over a 158C interval until the upper stability limit of pargasite is crossed. Because amphibole contains 2 wt % structural H2O, the mantle assemblage at the dehydration solidus contains 0·2 wt % H2O. Based on a water solubility of 25 wt % in a mantle melt at 2·5 GPa (Green et al., 2008), 1^2 wt % melt will be generated containing 20 wt % dissolved H2O (see isopleths in Fig. 21). This estimate assumes that the water (and most of the incompatible trace elements) sequestered in the amphibole enters the melt and that a roughly equal amount of anhydrous phases is involved in the melting reaction. The stoichiometry of the melting reaction is not well known, but amphibole most probably melts incongruently, with olivine being the most abundant product phase (Niida & Green, 1999). In the partial melt zone, small amounts of hydrous melt will reduce the overall density of the mantle and render it buoyant. As such, the partially molten zone will become gravitationally unstable and rise, either as a diapir (Tatsumi, 1983) or by porous or channelized flow (Daines & Kohlstedt, 1994; Grove et al., 2002). A vertically ascending diapir beneath Mt. Baker will encounter lower-P and higher-T conditions, resulting in additional melting in the diapir and reduction of the water content of the aggregate melt. At 10% melting, pMELTS calculations (Ghiorso et al., 2002) indicate 2 wt % H2O in the melt (Fig. 21). The extent and composition of the melt in the ascending diapir depends on how rapidly the diapir equilibrates


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T (ºC) Fig. 21. Pressure vs temperature diagram illustrating melt generation and ascent beneath Mt. Baker. Bold black curve is P^T path of diapirs ascending beneath Mt. Baker. Inset shows schematic cross-section of the magma transit path. The phase relations of amphibole peridotite (gray) and the P^T pathway for mantle flow (bold gray curve) are from Fig. 20. Abbreviations as in Fig. 20. White circles, modal % amphibole. Mantle melting is enhanced at the intersection of the mantle pathway with the dehydration solidus. Temperatures and pressures along the ascent path are based on the northern Cascadia model of Syracuse et al. (2010). Additional melting in the diapir during ascent reduces the water content in the aggregate melt. The wt % H2O isopleths (dashed curves) of vapor-undersaturated mantle melts are based on pMELTS calculations at 1·5 and 2 GPa (Ghiorso et al., 2002). White squares, P^Tdata for Mt. Baker basalts (Fig. 13).

internally as heat is transferred from the surrounding hotter mantle. Based on calculated temperatures and H2O contents of the Mt. Baker primary magmas, the diapirs are thermally equilibrated with the hotter mantle through which they pass. In addition to reducing the H2O contents of aggregate melts, re-equilibration at progressively lower pressures has significant effects on the major and trace element compositions of the basalts and explains why melting is initiated in the garnet stability field yet the basalts show no evidence for residual garnet. Thermal re-equilibration with the mantle during ascent will continue until the melt segregates from its residue. Thermobarometry and phase equilibria indicate that the basalts segregated from the mantle between 1·5 GPa (52 km) and 1GPa (40 km), corresponding to depths ranging from just above the hot core of the mantle wedge to the Moho and within the stability field of spinel lherzolite (Fig. 21). It is possible that a diapir could rise adiabatically through the overlying mantle, and, if so, it would follow a path of near constant temperature and decreasing pressure. However, the temperatures calculated for the Mt. Baker primary basalts suggest that adiabatic uprise did not occur until the diapirs reached a critical melt percentage.

S U M M A RY A N D C O N C L U D I N G REMARKS Reconstructed Mt. Baker primary melts define three geochemically and petrographically distinct groups that are similar to end-member basalt groups previously defined for the Cascade arc: Group I, calc-alkaline basalts (CAB); Group II, high-Mg basaltic andesites HMBA); Group III, low-K olivine tholeiite (LKOT). The geochemical and petrographic differences between the Mt. Baker basalt groups reflect differences in mantle melt fractions, which are directly related to the quantity of subduction component added from the subducting slab.The basalt compositions are best fit with an SC consisting mainly of metabasalt fluid, 84% except in the case of Group III (Park Butte) at 90%. The difference may reflect the fact that Park Butte is older (745 ka) than the other basalts (maximum 331ka). The remainder of the total SC consists of 7% sediment melt and 10% metabasalt melt (3% and 7%, respectively, for Park Butte). The Mt. Baker basalts do record the signature of sediment sampled at the trench (most clearly by Pb isotopes) despite seismic evidence indicating that little sediment has subducted in northern Cascadia since the inception of the arc (Davis & Hyndman,1989).

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isotopically dissimilar, pointing to the presence of multiple mantle components (Bacon et al., 1997; Jicha et al., 2008; Schmidt et al., 2008). The Mt. Baker basalts record variable extents of mantle metasomatism by the SC, but SC inputs are high enough such that all of the basalts have uniform isotope ratios equal to the isotopic ratios of the SC. The low proportion of slab melt in the SC (up to 17%) is consistent with the low H2O contents predicted for subducting sediment and metabasalt beneath the Cascadia fore-arc and arc axis, precluding vapor-saturated slab melting and limiting the extent of slab dehydration melting. Owing to the high temperatures associated with northern Cascadia, release of fluid from the slab is largely complete before reaching sub-arc depths. Melts may be released from the slab directly beneath Mt. Baker upon dehydration melting of phengite. However, only a small quantity of slab melt is recorded by the basalts, and most of the subduction component is hydrous fluid. Therefore, mantle hydrated by fluid under shallower conditions must be down-dragged to conditions at which the dehydration solidus is intersected, consistent with predictions for slab^ mantle wedge coupling based on geophysical models (Wada et al., 2008). Initiation of mantle melting beneath Mt. Baker occurs at 3·5 GPa, above the slab^mantle wedge interface, by dehydration melting of amphibole lherzolite. Initial melts have high water contents, up to 20 wt %. Lower-density domains of melt plus mantle residue will be gravitationally unstable and rise buoyantly through the hot core of the convecting mantle. Additional melting and re-equilibration during ascent through the hot core of the mantle wedge, up to segregation depths, results in aggregate melts with reduced water contents (consistent with those determined for the Mt. Baker basalts) and eliminates any evidence of residual garnet.

AC K N O W L E D G E M E N T S This paper is dedicated to the memory of Nathan Green, whose pioneering research on the mafic volcanic rocks of the Garibaldi Belt has stimulated and informed much of our own research. The research presented in this paper formed part of the first author’s Ph.D. dissertation completed at the University of Washington. We thank members of the supervisory committee, who put in a great deal of time and effort to review and critique several drafts of dissertation chapters. Don Swanson at the Hawaiian Volcano Observatory and Bernard Evans at the University of Washington graciously served as members of the thesis reading committee. The senior author thanks Dominique Weis of the University of British Columbia for financial and moral support during the preparation of this paper. We gratefully acknowledge Dave Tucker and Jeff Tepper for providing advice and samples from the Hannegan caldera and Chilliwack batholith. The paper benefited greatly from comprehensive reviews by Susan DeBari and two

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Of the Mt. Baker basalts, the CAB have the lowest liquidus H2O contents (1·5^2·1wt %), melt fractions (5^7%), and SC inputs. They record lherzolite ( spinel) mantle residues and segregated at the highest P^T conditions (up to 13548C, 1·5 GPa), slightly shallower than the hot core of the mantle wedge. The HMBA have the highest H2O contents (2·7^3·7 wt %) and largest SC inputs, represent the largest melt fractions (10^12%), and are associated with exhaustion of clinopyroxene in the mantle source, resulting in depleted (harzburgitic) residues. These lavas segregated from the mantle wedge at shallower (1GPa) and cooler (12708C) conditions, near the base of crust. These results are consistent with the origins proposed for HMBA in the southern Cascades by most workers (Baker et al., 1994; Borg et al., 1997; Clynne & Borg, 1997; Grove et al., 2003; Ruscitto et al., 2011). An alternative model for Mt. Shasta HMBA is contamination of basaltic magmas by ultramafic country rock (Streck et al., 2007); this process has been ruled out for Mt. Baker HMBA (Moore & DeBari, 2012). Mt. Baker HMBA also lack the unusually high Ni contents in olivine that characterize HMBA derived from pyroxenite mantle sources (Straub et al., 2008). Origination as a slab melt, as proposed for HMBA in other arcs (e.g. Kay et al., 1978), is also inconsistent with the similarity of Dy/Yb and isotope ratios of Mt. Baker HMBA to those of the other Mt. Baker basalts. The Mt. Baker LKOTrecords among the highest P^Tof mantle equilibration (1·4 GPa, 13268C), a relatively high melt fraction (11%), and intermediate values for H2O (2·1wt %) and SC input. These results contrast with LKOT in the central and southern Cascades, which represent near-anhydrous decompression melts of lherzolite with minimal slab input (Bartels et al., 1991; Sisson & Layne, 1993; Baker et al., 1994; Elkins-Tanton et al., 2001). Redox states are similar in all of the basalts, with average fO2 values approximately equivalent to the (QFM þ1) buffer, indicating that the redox state of the mantle may be buffered to the input of water. The fO2 values for Mt. Baker are in the middle of the range defined by other primitive basalts in the Cascade arc (QFM 2 to þ4) and at the lower end of the global fO2 range for arc lavas (QFM to QFM þ6) (Carmichael, 1991; Parkinson & Arculus, 1999; Righter, 2000; Rowe et al., 2009). The Mt. Baker basalts are more reduced than coeval andesites (Mullen & McCallum, 2013). The isotopic compositions of the basalts are tightly grouped at the depleted end of the isotopic arrays defined by the Mt. Baker volcanic field and by other primitive Cascade arc magmas. The isotopic similarity among the Mt. Baker basalts points to a relatively uniform depleted mantle wedge, similar to the source of northern Juan de Fuca MORB (but overprinted by slab input). At other Cascade volcanic centers, distinct basalt groups tend to be


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anonymous reviewers, as well as the careful editorial handling of Wendy Bohrson.

F U N DI NG This work was supported by funding to the first author from a National Science Foundation Graduate Research Fellowship, a University of Washington NASA Space Grant Fellowship, and several graduate research fund awards and fellowships (Peter Misch, David A. Johnston, and Wernicke Livingston) from the Department of Earth and Space Sciences at the University of Washington.

Supplementary data for this paper are available at Journal of Petrology online.

R EF ER ENC ES Adam, J. & Green, T. (2006). Trace element partitioning between mica- and amphibole-bearing garnet lherzolite and hydrous basanitic melt: 1. Experimental results and the investigation of controls on partitioning behavior. Contributions to Mineralogy and Petrology 152, 1^17. Adam, J., Green, T. H., Sie, S. H. & Ryan, C. G. (1997). Trace element partitioning between aqueous fluids, silicate melts and minerals. EuropeanJournal of Mineralogy 9, 569^584. Armstrong, J. T. (1988). Quantitative analysis of silicate and oxide minerals: comparison of Monte Carlo, ZAF, and j(rz) procedures. In: Newbury, D. E. (ed.) Microbeam Analysisç1988. San Francisco, CA: San Francisco Press, pp. 239^246. Asimow, P. D. & Ghiorso, M. S. (1998). Algorithmic modifications extending MELTS to calculate subsolidus phase relations. American Mineralogist 83, 1127^1131. Audet, P., Bostock, M. G., Mercier, J.-P. & Cassidy, J. F. (2008). Morphology of the Explorer^Juan de Fuca slab edge in northern Cascadia: Imaging plate capture at a ridge^trench^transform triple junction. Geology 36, 895^898. Ayers, J. (1998). Trace element modeling of aqueous fluid^peridotite interaction in the mantle wedge of subduction zones. Contributions to Mineralogy and Petrology 132, 390^404. Bacon, C. R., Bruggman, P. E., Christiansen, R. L., Clynne, M. A., Donnelly-Nolan, J. M. & Hildreth, W. (1997). Primitive magmas at five Cascade volcanic fields: Melts from hot, heterogeneous subarc mantle. Canadian Mineralogist 35, 397^423. Baker, M. B., Grove, T. L. & Price, R. (1994). Primitive basalts and andesites from the Mt. Shasta region, N. California; products of varying melt fraction and water content. Contributions to Mineralogy and Petrology 118, 111^129. Bartels, K. S., Kinzler, R. J. & Grove, T. L. (1991). High pressure phase relations of primitive high-alumina basalts from Medicine Lake Volcano, Northern California. Contributions to Mineralogy and Petrology 108, 253^270. Batt, G. E., Brandon, M. T., Farley, K. A. & Roden-Tice, M. (2001). Tectonic synthesis of the Olympic Mountains segment of the Cascadia wedge, using two-dimensional thermal and kinetic modeling of thermochronological ages. Journal of Geophysical Research 106(B11), 26731^26746. Becker, H., Jochum, K. P. & Carlson, R. W. (2000). Trace element fractionation during dehydration of eclogites from high-pressure

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terranes and the implications for element fluxes in subduction zones. Chemical Geology 163, 65^99. Berman, R.G. (1988). Internally-consistent thermodynamic data for minerals in the system Na2O^K2O^CaO^MgO^FeO^Fe2O3^ Al2O3^SiO2^TiO2^H2O^CO2. Journal of Petrology 29, 445^522. Blakely, R. J., Brocher, T. M. & Wells, R. E. (2005). Subduction-zone magnetic anomalies and implications for hydrated forearc mantle. Geology 33(6), 445^448, doi:10.1130/G21447.1. Borg, L. E., Clynne, M. A. & Bullen, T. D. (1997). The variable role of slab-derived fluids in the generation of a suite of primitive calc alkaline lavas from the southernmost Cascades, California. Canadian Mineralogist 35, 425^452. Borg, L. E., Blichert-Toft, J. & Clynne, M. A. (2002). Ancient and modern subduction zone contributions to the mantle sources of lavas from the Lassen region of California inferred from Lu^Hf isotopic systematics. Journal of Petrology 43, 705^723. Bostock, M. G. & VanDecar, J. C. (1995). Upper mantle structure of the northern Cascadia subduction zone. Canadian Journal of Earth Science 32, 1^12. Bowen, N. L. & Schairer, J. F. (1935). The system MgO^FeO^SiO2. AmericanJournal of Science 29, 151^217. Brandon, M. T., Cowan, D. S. & Vance, J. A. (1988).The Late Cretaceous San Juan thrust system, San Juan Islands, Washington. Geological Society of America, Special Papers 221, 79 p. Braunmiller, J. & Nabelek, J. (2002). Seismotectonics of the Explorer region. Journal of Geophysical Research 107(B10), 2208, doi:10.1029/ 2001JB000220. Brenan, J. M., Shaw, H. F., Ryerson, F. J. & Phinney, D. L. (1995). Mineralâqueous fluid partitioning of trace elements at 9008C and 2·0 GPa: constraints on the trace element chemistry of mantle and deep crustal fluids. Geochimica et Cosmochimica Acta 59, 3331^3350. Briqueu, L., Bougault, H. & Joron, J. L. (1984). Quantification of Nb, Ta, Ti and V anomalies in magmas associated with subduction zones: petrogenetic implications. Earth and Planetary Science Letters 68, 297^308. Brocher, T. M., Parsons, T., Trehu, A. M., Snelson, C. M. & Fisher, M. A. (2003). Seismic evidence for widespread serpentinized forearc upper mantle along the Cascadia margin. Geology 31, 267^270. Brown, E. H. & Gehrels, G. E. (2007). Detrital zircon constraints on terrane ages and affinities and timing of orogenic events in the San Juan Islands and North Cascades, Washington. Canadian Journal of Earth Science 44, 1375^1396. Brown, E. H., Wilson, D. L., Armstrong, R. L. & Harakal, J. E. (1982). Petrologic, structural, and age relations of serpentinite, amphibolite, and blueschist in the Shuksan Suite of the Iron Mountain^ Gee Point area, North Cascades, Washington. Geological Society of America Bulletin 93, 1087^1098. Carmichael, I. S. E. (1991). The redox states of basic and silicic magmas: a reflection of their source regions. Contributions to Mineralogy and Petrology 106, 121^141. Carpentier, M., Weis, D. & Chauvel, C. (2010). Chemical and isotopic constraints on sedimentary input into the Northern Cascades Arc system. Geochimica et Cosmochimica Acta 74, Supplement 1, A144. Carpentier, M., Weis, D. & Chauvel, C. (2013). Large U loss during weathering of upper continental crust: the sedimentary record. Chemical Geology 340, 91^104. Claeson, D. T. & Meurer, W. P. (2004). Fractional crystallization of hydrous basaltic ‘arc-type’ magmas and the formation of amphibolebearing gabbroic cumulates. Contributions to Mineralogy and Petrology 147, 288^304. Clynne, M. A. & Borg, L. E. (1997). Olivine and chromian spinel in primitive calc-alkaline and tholeiitic lavas from the southernmost

276


S U P P L E M E N TA RY DATA

NUMBER 2

MULLEN & McCALLUM


on the dynamics of subduction zones. Journal of Petrology 46, 555^578. Gaetani, G. A. (2004). The influence of melt structure on trace element partitioning near the peridotite solidus. Contributions to Mineralogy and Petrology 147, 511^527. Gaetani, G. A. & Grove, T. L. (1998). The influence of water on melting of mantle peridotite. Contributions to Mineralogy and Petrology 131, 323^346. Gaffney, A. M., Blichert-Toft, J., Nelson, B. K., Bizzarro, M., Rosing, M. & Albare'de, F. (2007). Constraints on source-forming processes of West Greenland kimberlites inferred from Hf^Nd isotope systematics. Geochimica et Cosmochimica Acta 71, 2820^2836. Galer, S. J. G. & Abouchami, W. (1998). Practical application of lead triple spiking for correction of instrumental mass discrimination. Mineralogical Magazine 62A, 491^492. Ghiorso, M. S. & Evans, B. W. (2008). Thermodynamics of rhombohedral oxide solid solutions and a revision of the P^T two-oxide geothermometer and oxygen-barometer. American Journal of Science 308, 957^1039. Ghiorso, M.S. & Sack, R.O. (1995). Chemical mass transfer in magmatic processes, IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid^ solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119, 197^212. Ghiorso, M. S., Hirschmann, M. M., Reiners, P. W. & Kress, V. C. (2002). The pMELTS: a revision of MELTS for improved calculation of phase relations and major element partitioning related to partial melting of the mantle to 3 GPa. Geochemistry, Geophysics, Geosystems 3, doi:10.1029/2001GC000217. Gill, J. G. (1981). Orogenic Andesites and Plate Tectonics. Berlin: Springer, 390 p. Green, D. H., Hibberson, W. O. & O’Neill, H. S. C. (2008). Clarification of the influence of water on mantle wedge melting. Geochimica et Cosmochimica Acta 72, A325Â325. Green, D. H., Hibberson, W. O., Kovacs, I. & Rosenthal, A. (2010). Water and its influence on the lithosphereâsthenosphere boundary. Nature 467, 448^452. Green, D. H., Rosenthal, A. & Kovacs, I. (2012). Comment on ‘‘The beginnings of hydrous mantle wedge melting’’, C. B. Till, T. L. Grove, A. C. Withers, Contributions to Mineralogy and Petrology, DOI 10.1007/s00410-011-0692-6. Contributions to Mineralogy and Petrology 164, 1077^1081. Green, N. L. (1988). Basalt^basaltic andesite mixing at Mount Baker volcano, Washington, I. Estimation of mixing conditions. Journal of Volcanology and Geothermal Research 34, 251^265. Green, N. L. (2006). Influence of slab thermal structure on basalt source regions and melting conditions: REE and HFSE constraints from the Garibaldi volcanic belt, northern Cascadia subduction system. Lithos 87, 23^49. Green, N. L. & Harry, D. L. (1999). On the relationship between subducted slab age and arc basalt petrogenesis, Cascadia subduction system, North America. Earth and Planetary Science Letters 171, 367^381. Green, N. L. & Sinha, A. K. (2005). Consequences of varied slab age and thermal structure on enrichment processes in the sub-arc mantle of the northern Cascadia subduction system. Journal of Volcanology and Geothermal Research 140, 107^132. Green, N. L., Armstrong, R. L., Harakal, J. E., Souther, J. G. & Read, P. B. (1988). Eruptive history and KÂr geochronology of the late Cenozoic Garibaldi volcanic belt, southwestern British Columbia. Geological Society of America Bulletin 100, 563^579. Green, T. H., Blundy, J. D., Adam, J. & Yaxley, G. M. (2000). SIMS determination of trace element partition coefficients between

277


Cascade range, California: a reflection of relative fertility of the source. Canadian Mineralogist 35, 453^472. Conrey, R. M., Sherrod, D., Hooper, P. & Swanson, D. (1997). Diverse primitive magmas in the Cascade arc, Northern Oregon and Southern Washington. Canadian Mineralogist 35, 367^396. Cousens, B. L., Allan, J. F., Leybourne, M. I., Chase, R. L. & Van Wagoner, N. (1995). Mixing of magmas from enriched and depleted mantle sources in the northeast Pacific: West Valley segment, Juan de Fuca Ridge. Contributions to Mineralogy and Petrology 120(3^4), 337^357. Crider, J. G., Frank, D., Malone, S. D., Poland, M. P., Werner, C. & Caplan-Auerbach, J. (2011). Magma at depth: A retrospective analysis of the 1975 unrest at Mount Baker, Washington, USA. Bulletin of Volcanology 73(2), 175^189. Currie, C. A., Wang, K., Hyndman, R. D. & He, J. (2004). The thermal effects of steady-state slab-driven mantle flow above a subducting plate: The Cascadia subduction zone and back arc. Earth and Planetary Science Letters 223, 35^48, doi:10.1016/ j.epsl.2004.04.020. Daines, M. L. & Kohlstedt, D. L. (1994). The transition from porous to channelized flow due to melt/rock reaction during melt migration. Geophysical Research Letters 21, 145^148. Davidson, J., Turner, S. & Plank, T. (2013). Dy/Dy*: Variations arising from mantle sources and petrogenetic processes. Journal of Petrology 54, 525^537. Davis, A. S., Clague, D. A., Cousens, B. L., Keaten, R. & Paduan, J. B. (2008). Geochemistry of basalt from the North Gorda segment of the Gorda Ridge: Evolution toward ultraslow spreading ridge lavas due to decreasing magma supply. Geochemistry, Geophysics, Geosystems 9, Q04004, doi:10.1029/2007GC001775. Davis, E. E. & Hyndman, R. D. (1989). Accretion and recent deformation of sediments along the northern Cascadia subduction zone. Geological Society America Bulletin 101, 1465^1480. Davis, E. E., Hyndman, R. D. & Villinger, H. (1990). Rates of fluid expulsion across the northern Cascadia accretionary prism: constraints from new heat flow and multichannel seismic reflection data. Journal of Geophysical Research 95(B6), 8869^8889. Donnelly, K. E., Goldstein, S. W., Langmuir, C. H. & Spiegelman, M. (2004). Origin of enriched ocean ridge basalts and implications for mantle dynamics. Earth and Planetary Science Letters 226, 347^366. Du Bray, E. A., John, D. A., Sherrod, D. R., Evarts, R. C., Conrey, R. M. & Lexa, J. (2006). Geochemical database for volcanic rocks of the Western Cascades, Washington, Oregon, and California. US Geological Survey Data Series 155, 49 p. Elkins, L. J., Gaetani, G. A. & Sims, K. W. W. (2008). Partitioning of U and Th during garnet pyroxenite partial melting: Constraints on the source of alkaline ocean island basalts. Earth and Planetary Science Letters 265, 270^286. Elkins-Tanton, L. T., Grove, T. L. & Donnelly-Nolan, J. (2001). Hot, shallow mantle melting under the Cascades volcanic arc. Geology 29, 631^634. Feineman, M. D., Ryerson, F. J., DePaolo, D. J. & Plank, T. (2007). Zoisiteâqueous fluid trace element partitioning with implications for subduction zone fluid composition. Chemical Geology 239, 250^265. Frank, D., Meier, M. F. & Swanson, D. A. (1977). Assessment of increased thermal activity at Mount Baker, Washington, March 1975^March 1976. US Geological Survey, Professional Papers 1022-A, 49 p. Frost, B. R. (1991). Introduction to oxygen fugacity and its petrologic importance. In: Lindsley, D. H. (ed.) Oxide Minerals: Petrologic and Magnetic Significance. Mineralogical Society of America, Reviews in Mineralogy 25, 1^9. Fumagalli, P. & Poli, S. (2005). Experimentally determined phase relations in hydrous peridotites to 6·5 GPa and their consequences


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Jicha, B. R., Hart, G. L., Johnson, C. M., Hildreth, W., Beard, B. L., Shirey, S. B. & Valley, J. W. (2008). Isotopic and trace element constraints on the petrogenesis of lavas from the Mount Adams volcanic field, Washington. Contributions to Mineralogy and Petrology 157, 189^207. Johnson, D. J., Hooper, P. R. & Conrey, R. M. (1999). XRFanalyses of rocks and mineral for major and trace elements on a single low dilution Li-tetraborate fused bead. Advances in X-ray Analysis 41, 843^867. Johnson, E. R., Wallace, P. J., Granados, H. D., Manea, V. C., Kent, A. J., Bindeman, I. N. & Donegan, C. S. (2009). Subduction-related volatile recycling and magma generation beneath Central Mexico: insights from melt inclusions, oxygen isotopes and geodynamic models. Journal of Petrology 50, 1729^1764. Johnson, K. T. M. (1998). Experimental determination of partition coefficients for rare earth and high-field-strength elements between clinopyroxene, garnet, and basaltic melt at high pressures. Contributions to Mineralogy and Petrology 133, 60^68. Johnson, M. C. & Plank, T. (1999). Dehydration and melting experiments constrain the fate of subducted sediments. Geochemistry, Geophysics, Geosystems 1, paper no. 1999GC000014. Kay, R. W., Sun, S.-S. & Lee-Hu, C.-N. (1978). Pb and Sr isotopes in volcanic rocks from the Aleutian Islands and Pribilof Islands, Alaska. Geochimica et Cosmochimica Acta 42, 263^273. Kelley, K. A., Plank, T., Ludden, J. N. & Staudigel, H. (2003). The composition of altered oceanic crust at ODP sites 801 and 1149. Geochemistry, Geophysics, Geosystems 4, doi:10.1029/ 2002GC000435. Kennedy, A. K., Lofgren, G. E. & Wasserburg, G. J. (1993). An experimental study of trace element partitioning between olivine, orthopyroxene and melt in chondrules: equilibrium values and kinetic effects. Earth and Planetary Science Letters 115, 177^195. Kerrick, D. M. & Connolly, J. A. D. (2001). Metamorphic devolatilization of subducted oceanic metabasalts: implications for seismicity, arc magmatism, and volatile recycling. Earth and Planetary Science Letters 189, 19^29. Kessel, R., Schmidt, M. W., Ulmer, P. & Pettke, T. (2005a). Trace element signature of subduction-zone fluids, melts and supercritical liquids at 120^180 km depth. Nature 437, 724^727. Kessel, R., Ulmer, P., Pettke, T., Schmidt, M. W. & Thompson, A. B. (2005b). The water^basalt system at 4 to 6 GPa: Phase relations and second critical endpoint in a K-free eclogite at 700 to 14008C. Earth and Planetary Science Letters 237, 873^892. Kinzler, R. J. & Grove, T. L. (1992). Primary magmas of mid-ocean ridge basalts; 1, Experiments and methods. Journal of Geophysical Research 97, 6885^6906. Kirby, S. H. & Wang, K. (2002). Introduction to a global systems approach to Cascadia slab processes and associated earthquake hazards. In: Kirby, S., Wang, K. & Dunlop, S. (eds) The Cascadia subduction zone and related subduction systems. US Geological Survey, Open-File Reports 02-328, 5^8. Kirby, S. H., Engdahl, E. R. & Villasen‹or, A. (2002). Warm slab subduction as a global process. In: Kirby, S., Wang, K. & Dunlop, S. (eds) The Cascadia subduction zone and related subduction systems. US Geological Survey, Open-File Reports 02-328, 79^80. Klimm, K., Blundy, J. D. & Green, T. H. (2008). Trace element partitioning and accessory phase saturation during H2O-saturated melting of basalt with implications for subduction zone chemical fluxes. Journal of Petrology 49, 523^553. Kress, V. C. & Carmichael, I. S. E. (1991). The compressibility of silicate liquids containing Fe2O3 and the effect of composition, temperature, oxygen fugacity and pressure on their redox states. Contributions to Mineralogy and Petrology 108, 82^92.

278


garnet, clinopyroxene and hydrous basaltic liquids at 2^7·5 GPa and 1080^12008C. Lithos 53, 165^187. Grove, T. L., Parman, S. W., Bowring, S. A., Price, R. C. & Baker, M. B. (2002). The role of an H2O-rich fluid component in the generation of primitive basaltic andesites and andesites from the Mt. Shasta region, N. California. Contributions to Mineralogy and Petrology 142, 375^396. Grove, T. L., Elkins-Tanton, L. T., Parman, S. W., Chatterjee, N., Mu«ntener, O. & Gaetani, G. A. (2003). Fractional crystallization and mantle-melting controls on calc-alkaline differentiation trends. Contributions to Mineralogy and Petrology 145, 515^533. Gudfinnsson, G. H. & Presnall, D. C. (1996). Melting relations of model lherzolite in the system CaO^MgOÂl2O3^SiO2 at 2·4^ 3·4 GPa and the generation of komatiites. Journal of Geophysical Research 101, 27701^27709. Hacker, B. R., Abers, G. A. & Peacock, S. M. (2003a). Subduction factory 1. Theoretical mineralogy, densities, seismic wave speeds, and H2O contents. Journal of Geophysical Research 108(B1), 2029, doi:10.1029/2001JB001127. Hacker, B. R., Peacock, S. M., Abers, G. A. & Holloway, S. D. (2003b). Subduction factory 2. Are intermediate-depth earthquakes in subducting slabs linked to metamorphic dehydration reactions? Journal of Geophysical Research 108(B1), 2030, doi:10.1029/ 2001JB001129. Halliday, A. N., Lee, D. C., Tommasini, S., Davies, G., Paslick, C. R., Fitton, J. G. & James, D.E. (1995). Incompatible trace elements in OIB and MORB and source enrichments in the sub-oceanic mantle. Earth and Planetary Science Letters 133, 379^395. Harkins, S. A., Appold, M. S., Nelson, B. K., Brewer, A. M. & Groves, I. M. (2008). Lead isotope constraints on the origin of nonsulfide zinc and sulfide zinc^lead deposits in the Flinders Ranges, South Australia. Economic Geology 103, 353^364. Hart, S. R. (1984). A large scale isotope anomaly in the Southern Hemisphere mantle. Nature 309, 743^757. Hart, S. & Dunn, T. (1993). Experimental cpx/melt partitioning of 24 trace elements. Contributions to Mineralogy and Petrology 13, 1^8. Hauri, E. H., Wagner, T. P. & Grove, T. L. (1994). Experimental and natural partitioning of Th, U, Pb and other trace elements between garnet, clinopyroxene and basaltic melts. Chemical Geology 117, 149^166. Hawkesworth, C. J., Gallagher, K., Hergt, J. M. & McDermott, F. (1993). Trace element fractionation processes in the generation of island arc basalts. Philosophic Transactions of the Royal Society of London, Series A 342, 179^191. Hermann, J. & Rubatto, D. (2009). Accessory phase control on the trace element signature of sediment melts in subduction zones. Chemical Geology 265, 512^526. Hermann, J. & Spandler, C. (2008). Sediment melts at sub-arc depths: an experimental study. Journal of Petrology 49, 717^740. Hildreth, W. H. (2007). Quaternary magmatism in the Cascades; geologic perspectives. US Geological Survey, Professional Papers 1744, 125 p. Hildreth, W., Fierstein, J. & Lanphere, M. A. (2003). Eruptive history and geochronology of the Mount Baker volcanic field, Washington. Geological Society of America Bulletin 115, 729^764. Hildreth, W., Lanphere, M. A., Champion, D. E. & Fierstein, J. (2004). Rhyodacites of Kulshan Caldera, north Cascades of Washington; post-caldera lavas that span the Jaramillo. Journal of Volcanology and Geothermal Research 130, 227^264. Hyndman, R. D. & Peacock, S. M. (2003). Serpentinization of the forearc mantle. Earth and Planetary Science Letters 212, 417^432. Hyndman, R. D. & Wang, K. (1993). Thermal constraints on the zone of major thrust earthquake failure: The Cascadia subduction zone. Journal of Geophysical Research 98, 2039^2060.

NUMBER 2

MULLEN & McCALLUM


Washington and British Columbia. Geological Society of America Bulletin 103, 189^209. Me¤trich, N. & Wallace, P. J. (2008). Volatile abundances in basaltic magmas and their degassing paths tracked by melt inclusions. In: Putirka, K. D. & Tepley, F. J., III (eds) Minerals, Inclusions and Volcanic Processes. Mineralogical Society of America and Geochemical Society, Reviews in Mineralogy and Geochemistry 69, 363^402. Miller, K. C., Keller, G. R., Gridley, J. M., Luetgert, J. H., Mooney, W. D. & Thybo, H. (1997). Crustal structure along the west flank of the Cascades, western Washington. Journal of Geophysical Research 102(B8), 17857^17873. Misch, P. (1966). Tectonic evolution of the northern Cascades of Washington Stateça west Cordilleran case history. In: Gunning, H. C. (ed.) Tectonic history and mineral deposits of the western Cordillera. Canadian Institute of Mining and Metallurgy Special Volume 8, 101^148. Miyashiro, A. (1974). Volcanic rock series in island arcs and active continental margins. AmericanJournal of Science 274, 321^355. Monger, J. W. H. (1989). Geology of Hope and Ashcroft map areas. Geological Survey of Canada Maps 41-1989, 42-1989, scale 1:250 ,000. Moore, N. E. & DeBari, S. M. (2012). Mafic magmas from Mount Baker in the northern Cascade arc, Washington: probes into mantle and crustal processes. Contributions to Mineralogy and Petrology 163, 521^546. Mullen, E. K. (2011). Petrology and geochemistry of the Mount Baker volcanic field: constraints on source regions and terrane boundaries, and comparison with other Cascade Arc volcanic centers, Ph.D. dissertation, University of Washington, Seattle, 388 pp. Mullen, E. K. & McCallum, I. S. (2013). Coexisting pseudobrookite, ilmenite, and titanomagnetite in hornblende andesite of the Coleman Pinnacle flow, Mount Baker, Washington: Evidence for a highly oxidized arc magma. American Mineralogist 98, 417^425. Mullen, E. K. & Weis, D. (2013). Sr^Nd^Hf^Pb isotope and trace element evidence for the origin of alkalic basalts in the Garibaldi Belt, northern Cascade arc. Geochemistry, Geophysics, Geosystems 14, 3126^3155, doi:10.1002/ggge.20191. Mu«ntener, O., Kelemen, P. B. & Grove, T. L. (2001). The role of H2O during crystallization of primitive arc magmas under uppermost mantle conditions and genesis of igneous pyroxenites; an experimental study. Contributions to Mineralogy and Petrology 141, 643^658. Navon, O. & Stolper, E. (1986). Geochemical consequences of melt percolation: the upper mantle as a chromatographic column. Journal of Geology 95, 285^307. Nelson, B. K. (1995). Fluid flow in subduction zones: evidence from neodymium and strontium isotope variations in metabasalts of the Franciscan Complex, California. Contributions to Mineralogy and Petrology 119, 247^262. Nichols, M. L., Malone, S. D., Moran, S. C., Thelen, W. A. & Vidale, J. E. (2011). Deep long-period earthquakes beneath Oregon and Washington volcanoes. Journal of Volcanology and Geothermal Research 200, 116^128. Niida, K. & Green, D. H. (1999). Stability and chemical composition of pargasitic amphibole in MORB pyrolite under upper mantle conditions. Contributions to Mineralogy and Petrology 135, 18^40. Parkinson, I. J. & Arculus, R. J. (1999). The redox state of subduction zones: insights from arc-peridotites. Chemical Geology 160, 409^423. Parman, S. W. & Grove, T. L. (2004). Harzburgite melting with and without H2O: experimental data and predictive modeling. Journal of Geophysical Research 109, B02201, doi:10.1029/2003JB002566. Pawley, A. R. & Holloway, J. R. (1993). Water sources for subduction zone volcanism: new experimental constraints. Science 260, 664^667.

279


Kudo, A. M. & Weill, D. F. (1970). An igneous plagioclase thermometer. Contributions to Mineralogy and Petrology 25, 52^65. Kushiro, I., Syono, Y. & Akimoto, S. (1968). Melting of a peridotite nodule at high pressures and high water pressures. Journal of Geophysical Research 73, 6023^6029. Labanieh, S., Chauvel, C., Germa, A. & Quidelleur, X. (2012). Martinique: a clear case for sediment melting and slab dehydration as a function of distance to the trench. Journal of Petrology 53, 2441^2464. Lambert, I. B. & Wyllie, P. J. (1972). Melting of gabbro (quartz eclogite) with excess water to 35 kilobars, with geological applications. Journal of Geology 80, 693^708. Lange, R. A., Frey, H. N. & Hector, J. (2009). A thermodynamic model for the plagioclase^liquid hygrometer/thermometer. American Mineralogist 94, 494^506. Leeman, W. P., Smith, D. R., Hildreth, W., Palacz, Z. & Rogers, N. W. (1990). Compositional diversity of late Cenozoic basalts in a transect across the southern Washington Cascades; implications for subduction zone magmatism. Journal of Geophysical Research 95, 19561^19582. Leeman, W. P., Tonarini, S., Chan, L. H. & Borg, L. E. (2004). Boron and lithium isotopic variations in a hot subduction zone; the southern Washington Cascades. Chemical Geology 212, 101^124. Leeman, W. P., Lewis, J. F., Evarts, R. C., Conrey, R. M. & Streck, M. J. (2005). Petrologic constraints on the thermal structure of the Cascades arc. Journal of Volcanology and Geothermal Research 140, 67^105. Le Voyer, M., Rose-Koga, E. F., Shimizu, N., Grove, T. L. & Schiano, P. (2010). Two Contrasting H2O-rich components in primary melt inclusions from Mount Shasta. Journal of Petrology 51, 1571^1595. Li, Y., Massonne, H.-J., Willner, A., Tang, H.-F. & Liu, C.-Q. (2008). Dehydration of clastic sediments in subduction zones: Theoretical study using thermodynamic data of minerals. Island Arc 17, 577^590. Liu, J., Bohlen, S. R. & Ernst, W. G. (1996). Stability of hydrous phases in subducting oceanic crust. Earth and Planetary Science Letters 143, 161^171. Longhi, J. (1987). Liquidus equilibria and solid solution in the system Anorthite^Forsterite^Wollastonite^Silica at low pressure. American Journal of Science 287, 265^331. Longhi, J. (1991). Comparative liquidus equilibria of hypersthene normative basalts at low pressure. American Mineralogist 76, 785^800. Longhi, J. (2002). Some phase equilibrium systematics of lherzolite melting: I. Geochemistry, Geophysics, Geosystems 3, doi:10.1029/ 2001GC000204. Longhi, J., Vander Auwera, J., Fram, M. S. & Duchesne, J. C. (1999). Some phase equilibrium constraints on the origin of Proterozoic (massif) anorthosites and related rocks. Journal of Petrology 40, 339^362. McCrory, P. A., Blair, J. L., Oppenheimer, D. H. & Walter, S. R. (2004). Depth to the Juan de Fuca slab beneath the Cascadia subduction margin: A 3-D model for sorting earthquakes. US Geological Survey Data Series 91, 13 p. McCulloch, M. T. & Gamble, J. A. (1991). Geochemical and geodynamical constraints on subduction zone magmatism. Earth and Planetary Science Letters 102, 358^374. McDade, P., Blundy, J. D. & Wood, B. J. (2003). Trace element partitioning between mantle wedge peridotite and hydrous MgO-rich melt. American Mineralogist 88, 1825^1831. McDonough, W. F. & Sun, S.-s. (1995). The composition of the Earth. Chemical Geology 120, 223^253. McGroder, M. F. (1991). Reconciliation of two-sided thrusting, burial metamorphism, and diachronous uplift in the Cascades of


VOLUME 55

FEBRUARY 2014

high-Mg andesites. Contributions to Mineralogy and Petrology 162, 109^132. Ruscitto, D. M., Wallace, P. J., Cooper, L. B. & Plank, T. (2012). Global variations in H2O/Ce: 2. Relationships to arc magma geochemistry and volatile fluxes. Geochemistry, Geophysics, Geosystems 13, Q03025, doi:10.1029/2011GC003887. Sack, R. O. & Ghiorso, M. S. (1991). Chromite as a petrogenetic indicator. In: Lindsley, D. H. (ed.) Oxide Minerals: Petrologic and Magnetic Significance. Mineralogical Society of America, Reviews in Mineralogy 25, 323^353. Salters, V. J. M. & Longhi, J. (1999). Trace element partitioning during the initial stages of melting beneath mid-ocean ridges. Earth and Planetary Science Letters 166, 15^30. Salters, V. J. M. & Stracke, A. (2004). Composition of the depleted mantle. Geochemistry, Geophysics, Geosystems 5, doi:10.1029/ 2003GC000597. Schmidt, M. E., Grunder, A. L. & Rowe, M. C. (2008). Segmentation of the Cascade Arc as indicated by Sr and Nd isotopic variation among diverse primitive basalts. Earth and Planetary Science Letters 266, 166^181. Shaw, D. M. (1970). Trace element fractionation during anatexis. Geochimica et Cosmochimica Acta 34, 237^243. Shaw, S. D. (2011). H2O contents in olivine-hosted melt inclusions from primitive magmas in the northern Cascade arc, M.S. thesis, Western Washington University, Bellingham, 85 pp. Sisson, T. W. & Bronto, S. (1998). Evidence for pressure-release melting beneath magmatic arcs from basalt at Galunggung, Indonesia. Nature 391, 883^886. Sisson, T. W. & Grove, T. L. (1993). Temperatures and H2O contents of low-MgO high-alumina basalts. Contributions to Mineralogy and Petrology 113, 167^184. Sisson, T. W. & Layne, G. D. (1993). H2O in basalt and basaltic andesite glass inclusions from four subduction-related volcanoes. Earth and Planetary Science Letters 117, 619^635. Spandler, C., Mavrogenes, J. & Hermann, J. (2007). Experimental constraints on element mobility from subducted sediments using high-P synthetic fluid/melt inclusions. Chemical Geology 239, 228^249. Stalder, R. (2012). Comment on ‘‘The beginnings of hydrous mantle wedge melting’’, C. B. Till, T. L. Grove, A. C. Withers, Contributions to Mineralogy and Petrology, doi 10.1007/s00410-0110692-6. Contributions to Mineralogy and Petrology 164, 1069^1071. Staudigel, H., Davies, G. R., Hart, S. R., Marchant, K. M. & Smith, B. M. (1995). Large-scale isotopic Sr, Nd and O isotopic anatomy of altered oceanic crust: DSDP/ODP sites 417/418. Earth and Planetary Science Letters 130, 169^185. Stolper, E. & Newman, S. (1994). The role of water in the petrogenesis of Mariana trough magmas. Earth and Planetary Science Letters 121, 293^325. Straub, S. M., LaGatta, A. B., Pozzo, M. D., Lillian, A. & Langmuir, C. H. (2008). Evidence from high-Ni olivines for a hybridized peridotite/pyroxenite source for orogenic andesites from the central Mexican Volcanic Belt. Geochemistry, Geophysics, Geosystems 9, Q03007, doi:10.1029/2007GC001583. Streck, M. J., Leeman, W. P. & Chesley, J. (2007). Highmagnesian andesite from Mount Shasta: A product of magma mixing and contamination, not a primitive mantle melt. Geology 35(4), 351^354. Sun, S.-s. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implication for mantle composition and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publication 42, 313^345.

280


Peacock, S. M., Wang, K. & McMahon, A. M. (2002). Thermal structure and metamorphism of subducting oceanic crust: Insight into Cascadia intraslab earthquakes. In: Kirby, S., Wang, K. & Dunlop, S. (eds) The Cascadia subduction zone and related subduction systems. US Geological Survey, Open-File Reports 02-328, 123^126. Pearce, J. A. & Parkinson, I. J. (1993). Trace element models for mantle melting: application to volcanic arc petrogenesis. In: Prichard, H. M., Alabaster, T., Harris, N. B. W. & Neary, C. R. (eds) Magmatic Processes and Plate Tectonics. Geological Society, London, Special Publications 76, 373^403. Pearce, J. A. & Peate, D. W. (1995). Tectonic implications of the composition of volcanic arc magmas. Annual Review of Earth and Planetary Sciences 23, 251^285. Pearson, D. G., Canil, D. & Shirey, S. B. (2003). Mantle samples included in volcanic rocks: xenoliths and diamonds. In: Carlson, R. W. (ed.) The Mantle and Core, Vol. 2. Oxford: Elsevier^ Pergamon, pp. 171^276. Plank, T. & Langmuir, C. H. (1998). The chemical composition of subducting sediment and its consequences for the crust and mantle. Chemical Geology 145, 325^394. Portnyagin, M., Hoernle, K., Plechov, P., Mironov, N. & Khubunaya, S. (2007). Constraints on mantle melting and composition and nature of slab components in volcanic arcs from volatiles (H2O, S, Cl, F) and trace elements in melt inclusions from the Kamchatka Arc. Earth and Planetary Science Letters 255, 53^69. Presnall, D. C., Dixon, J. R., O’Donnell, T. H. & Dixon, S. A. (1979). Generation of mid-ocean ridge tholeiites. Journal of Petrology 20, 3^35. Putirka, K. D. (2008). Thermometers and barometers for volcanic systems. In: Putirka, K. D. & Tepley, F. J., III (eds) Minerals, Inclusions and Volcanic Processes. Mineralogical Society of America and Geochemical Society, Reviews in Mineralogy and Geochemistry 69, 61^120. Reiners, P. W., Hammond, P. E., McKenna, J. M. & Duncan, R. A. (2000). Young basalts of the central Washington Cascades, flux melting of the mantle, and trace element signatures of primary arc magmas. Contributions to Mineralogy and Petrology 138, 249^264. Rhodes, J. M., Dungan, M. A., Blanchard, D. P. & Long, P. E. (1979). Magma mixing at mid-ocean ridges: Evidence from basalts drilled near 228N on the mid-Atlantic ridge. Tectonophysics 55, 35^61. Riddihough, R. P. & Hyndman, R. D. (1991). The modern plate tectonic regime of the continental margin of western Canada. In: Gabrielse, H. & Yorath, C. J. (eds) Geology of the Cordilleran Orogen in Canada, Vol. 2. Geological Survey of Canada, Geology of Canada 4, 435^455. Righter, K. (2000). A comparison of basaltic volcanism in the Cascades and western Mexico: compositional diversity in continental arcs. Tectonophysics 318, 99^117. Robie, R. A. & Hemingway, B. S. (1995). Thermodynamic properties of minerals and related substances at 298·15 K and 1 bar (105 Pascals) pressure and at higher temperatures. US Geological Survey Bulletin 2131, 470 p. Rose, E. F., Shimizu, N., Layne, G. D. & Grove, T. L. (2001). Melt production beneath Mt. Shasta from boron data in primitive melt inclusions. Science 293, 281^283. Rowe, M. C., Kent, A. J. R. & Nielsen, R. L. (2009). Subduction influence on oxygen fugacity and trace and volatile elements in basalts across the Cascade volcanic arc. Journal of Petrology 50, 61^91. Ruscitto, D. M., Wallace, P. J., Johnson, E. R., Kent, A. J. R. & Bindeman, I. N. (2010). Volatile contents of mafic magmas from cinder cones in the Central Oregon High Cascades: Implications for magma formation and mantle conditions in a hot arc. Earth and Planetary Science Letters 298, 153^161. Ruscitto, D. M., Wallace, P. J. & Kent, A. J. R. (2011). Revisiting the compositions and volatile contents of olivine-hosted melt inclusions from the Mount Shasta region: implications for the formation of

NUMBER 2

MULLEN & McCALLUM


lake: The Sulphur Creek lava flow, North Cascades, Washington. Journal of Volcanology and Geothermal Research 185, 311^322. Ulmer, P. & Trommsdorff, V. (1995). Serpentine stability to mantle depths and subduction-related magmatism. Science 268, 858^861. van Keken, P. E., Kiefer, B. & Peacock, S. M. (2002). High-resolution models of subduction zones: Implications for mineral dehydration reactions and the transport of water into the deep mantle. Geochemistry, Geophysics, Geosystems 3(10), 1056, doi:10.1029/ 2001GC000256. Vielzeuf, D. & Schmidt, M.W. (2001). Melting relations in hydrous systems revisited: application to metapelites, metagreywackes and metabasalts. Contributions to Mineralogy and Petrology 141, 251^267. Wada, I. & Wang, K. (2009). Common depth of slab^mantle decoupling: Reconciling diversity and uniformity of subduction zones. Geochemistry, Geophysics, Geosystems 10, Q10009. Wada, I., Wang, K., He, J. & Hyndman, R. D. (2008). Weakening of the subduction interface and its effects on surface heat flow, slab dehydration, and mantle wedge serpentinization. Journal of Geophysical Research 113, B04402, doi:10.1029/2007JB005190. Walter, M. J. (2003). Melt extraction and compositional variability in mantle lithosphere. In: Carlson, R. W. (ed.) The Mantle and Core, Vol. 2. Oxford: Elsevier^Pergamon, pp. 363^394. Wasylenki, L. E., Baker, M. B., Kent, A. J. R. & Stolper, E. M. (2003). Near-solidus melting of the shallow upper mantle: Partial melting experiments on depleted peridotite. Journal of Petrology 44,1163^1191. Weis, D., Kieffer, B., Maerschalk, C., Barling, J., De Jong, J., Williams, G., Hanano, D., Pretorius, W., Mattielli, N., Scoates, J. S., Goolaerts, A., Friedman, R. M. & Mahoney, J. B. (2006). High-precision isotopic characterization of USGS reference materials by TIMS and MC-ICP-MS. Geochemistry, Geophysics, Geosystems 7, Q08006, doi:10.1029/2006GC001283. Wells, R. E., Weaver, C. S. & Blakely, R. J. (1998). Forearc migration in Cascadia and its neotectonic significance. Geology 26, 759^762. Werner, C., Evans, W. C., Poland, M., Tucker, D. S. & Doukas, M. P. (2009). Long-term changes in quiescent degassing at Mount Baker Volcano, Washington, USA: Evidence for a stalled intrusion in 1975 and connection to a deep magma source. Journal of Volcanology and Geothermal Research 186, 379^386. Wheeler, J. O. & McFeely, P. (1991). Tectonic assemblage map of the Canadian Cordillera and adjacent parts of the United States of America. Geological Survey of Canada Map 1712A, scale 1:2 000 000. Wilson, D. S. (2002). The Juan de Fuca plate and slab: Isochron structure and Cenozoic plate motions. In: Kirby, S., Wang, K. & Dunlop, S. (eds) The Cascadia subduction zone and related subduction systems. US Geological Survey, Open-File Reports 02-328, 9^12. Wood, B. J. (1991). Oxygen barometry of spinel peridotites. In: Lindsley, D. H. (ed.) Oxide Minerals: Petrologic and Magnetic Significance. Mineralogical Society of America, Reviews in Mineralogy 25, 323^353. Workman, R. K. & Hart, S. R. (2005). Major and trace element composition of the depleted MORB mantle (DMM). Earth and Planetary Science Letters 231, 53^72.

281


Syracuse, E. M., van Keken, P. E. & Abers, G. A. (2010). The global range of subduction zone thermal models. Physics of the Earth and Planetary Interiors 183, 73^90. Tabor, R. W., Haugerud, R. A., Hildreth, W. & Brown, E. H. (2003). Geologic map of the Mount Baker 30 60 minute quadrangle, Washington. US Geological Survey Map I-2660, scale 1:100 000. Tatsumi, Y. (2000). Slab melting: its role in continental crust formation and mantle evolution. Geophysical Research Letters 27, 3941^3944. Tatsumi, Y. (1983). Generation of arc basalt magmas and thermal structure of the mantle wedge in subduction zones. Journal of Geophysical Research 88(B7), 5815^5825. Tatsumi, Y. & Eggins, S. (1995). Subduction Zone Magmatism. Oxford: Blackwell Science. Tepley, F. J., Lundstrom, C. C., McDonough, W. F. & Thompson, A. (2010). Trace element partitioning between high-An plagioclase and basaltic to basaltic andesite melt at 1 atmosphere pressure. Lithos 118, 82^94. Tepper, J. H. (1996). Petrology of mafic plutons associated with calc-alkaline granitoids, Chilliwack Batholith, north Cascades, Washington. Journal of Petrology 37, 1409^1436. Tepper, J. H., Nelson, B. K., Bergantz, G. W. & Irving, A. J. (1993). Petrology of the Chilliwack batholith, North Cascades, Washington. Generation of calc-alkaline granitoids by melting of mafic lower crust with variable water fugacity. Contributions to Mineralogy and Petrology 113, 333^351. Thirlwall, M. F., Smith, T. E., Graham, A. M., Theodorou, N., Hollings, P., Davidson, J. P. & Arculus, R. J. (1994). High field strength element anomalies in arc lavas: source or process? Journal of Petrology 35, 819^838. Till, C. B., Grove, T. L. & Withers, A. C. (2012a). The beginnings of hydrous mantle wedge melting. Contributions to Mineralogy and Petrology 163, 669^688. Till, C. B., Grove, T. L. & Withers, A. C. (2012b). Reply to ‘Comment on ‘‘The beginnings of hydrous mantle wedge melting’’ by Till et al.’ by Green, Rosenthal and Kovacs. Contributions to Mineralogy and Petrology 164, 1083^1085. Till, C. B., Grove, T. L. & Withers, A. C. (2012c). Reply to ‘Comment on ‘‘The beginnings of hydrous mantle wedge melting’’ by Till et al.’ by Stalder. Contributions to Mineralogy and Petrology 164, 1073^1076. Toplis, M. J. (2005). The thermodynamics of iron and magnesium partitioning between olivine and liquid: criteria for assessing and predicting equilibrium in natural and experimental systems. Contributions to Mineralogy and Petrology 149, 22^39. Tucker, D. S. (2006). Cross section through the Hannegan Caldera. Geological Society of America Digital Map and Chart Series 3, 2 sheets, scale: 1:24 000. Tucker, D. S. & Scott, K. M. (2006). A magmatic component in 19th century Mount Baker eruptions? Geological Society of America, Abstracts with Programs 38, 75. Tucker, D. S. & Scott, K. M. (2009). Structures and facies associated with the flow of subaerial basaltic lava into a deep freshwater