madeira island alkaline lava spinels: petrogenetic implications

Mineralogy and Petrology (2004) 81: 85–111DOI 10.1007/s00710-004-0029-5

Madeira Island alkaline lava spinels:petrogenetic implications

J. Mata and J. Munha

Centro de Geologia, Departamento de Geologia da Universidade de Lisboa,Faculdade de Cieencias, Lisboa, Portugal

Received May 17, 2002; revised version accepted November 28, 2003Editorial handling: E. F. Stumpfl

Summary

Four groups of spinels have been identified in Madeira Island alkaline basalts: 1) Al-rich chromian-spinels (Cr#<0.15; 51<Mg#<62) characteristic of olivine xenocrystsfrom dismembered, high-pressure, cognate, ultramafic xenoliths; 2) Al-poor chromian-spinel=magnesiochromite=chromite (Cr#>31; 30<Mg#<61) included in olivinephenocryst cores; 3) chromian-titanomagnetites included in olivine phenocryst rims(11<Mg#<34) and chromian-titanomagnetites to titanomagnetites in clinopyroxenephenocrysts (6<Mg#<25) and in the groundmass (3<Mg#<36); 4) rare, TiO2-poor,MnO-rich titanomagnetite inclusions in green-core clinopyroxene xenocrysts, derivedfrom metasomatized upper mantle xenoliths. Chemical variations among spinel groups1) to 3) largely reflect physical conditions and the extent of fractionation of crystallisingmagmas. Extensive solid solution among chromian-spinel – ulvospinel – magnetiteand the ulvospinel enrichment exhibited by titanomagnetite evolutionary trends areattributed to the combined effects of low-aSiO2

and oxidizing conditions (0.2��log( f O2)NNO�1.8) during crystallisation from Madeira alkaline magmas. Pressuredoes not seem to have a direct influence on the stability of high-Al spinels; however,Cr=Al contrasts between spinels in high-pressure olivine xenocrysts and those in olivinephenocryst cores are envisaged as reflecting compositional effects of distinct crystal-lisation sequences during polibaric fractionation of Madeira magmas.

Introduction

Notwithstanding their minor modal proportions in common igneous rocks, spinelshave long been acknowledged as carrying useful information regarding the petro-genesis of their host lavas (Irvine, 1965, 1967). Spinels exhibit an important rangeof solid solution, expressed by wide variations on di-, tri-, and tetravalent cations,reflecting the possibility of their crystallisation over highly variable conditions;

thus, spinel composition may be a sensitive indicator of intensive and composi-tional properties of their host magmas (e.g. Arai, 1992; Toplis and Carrol, 1995;Kamenetsky et al., 2001; Barnes and Roeder, 2001).

In this paper we present and discuss the compositional variation of spinelsoccurring in Madeira alkaline lavas as a way to better understand and constrainthe petrogenetic processes operating in these magmas. Additionally, Madeira dataare used to assess the importance of some of the factors controlling the composi-tion and the crystallisation sequence of spinel minerals.

Synopsis of Madeira lavas geology and petrochemistry

Madeira Island (32� N, 17� W) is located in the Atlantic at the end of a northeast–southwest oriented chain of islands and seamounts, reflecting lithospheric drift(�1.2 cm=year) over a stationary long lived intermittent mantle plume(Geldmacher et al., 2000; Geldmacher and Hoernle, 2000). Mata et al. (1998)explained the origin of this plume by the destabilization of the S-wave anomalouslayer defined by Hoernle et al. (1995) as extending to depths of 500 km and cover-ing an area of 2500 by 4000 km. Isotopic age determinations indicate that theMadeira sub-aerial development occurred mostly during post-Miocene times, withthe last eruption occurring at 6000–7000 years before present (Mata and Munha,1999; Geldmacher et al., 2000).

Trace element and isotopic data (Mata et al., 1989, 1998; Mata and Kerrich,2000; Kogarko et al., 2000; Geldmacher and Hoernle, 2000) suggest that theMadeira mantle source was dominated by DMM and HIMU components; traceamounts of EM I were also suggested by Mata et al. (1998). The HIMU componentwas interpreted by Halliday et al. (1995) as a consequence of recent (<150 Ma)near surface mantle fractionation (see also Thirlwall, 1997). However, Mataet al. (1998) and Geldmacher and Hoernle (2000) rejected this hypothesis byconsidering the HIMU signatures inherited from significantly older (>850 Ma)recycling and incubation of hydrothermally altered oceanic lithosphere. Theinteraction of the ascending plume melts with the old (130 Ma) and thick(�100 km) Madeira lithosphere (Monnereau and Cazenave, 1988) is endorsed bythe trace element chemical evidence for magma equilibration with residual amphi-bole (Mata et al., 1998) and by the Os isotopic signatures of lavas (Widom et al.,1999).

Madeira lavas are predominantly of moderate alkaline affinities ranging incomposition from basanites (and picrobasalts) to mugearites (Table 1). Theircompositional range is more restricted than that of the analysed dykes, whichmay attain trachytic compositions (e.g. Hughes and Brown, 1972; Schminckeand Weibel, 1972; Aires-Barros et al., 1974; Mata et al., 1989). Evolutionaryprocesses have been explained mainly in terms of fractional crystallisation ofolivine, clinopyroxene and, to a lesser extent, plagioclase and oxides. Theseoccur as phenocrysts; the most abundant being, by far, olivine. The fractiona-tion sequence was dominated by early crystallisation of Cr-spinelþ olivine(Fo¼ 88�85 mol%), followed by olivine (Fo¼ 85! 50 mol%)þ diopside (Al!Ti rich)þ spinel (Cr–Ti spinel!Ti-magnetite)� plagioclase assemblages (seeTables 2 and 3); plagioclase (mostly, labradorite) fractionation becomes increas-

86 J. Mata and J. Munha

https://www.researchgate.net/publication/242877800_Seismic_and_geochemical_evidence_for_large-scale_mantle_upwelling_beneath_the_Eastern_Atlantic_and_Western_and_Central_Europe?el=1_x_8&enrichId=rgreq-e83efa0d-df2b-49f2-948d-b47f167c1d89&enrichSource=Y292ZXJQYWdlOzIyNzMzNzg1NztBUzoxMzU0NjAwMjYzMjcwNTlAMTQwOTMwNzU5NTg5Mg==

ingly important at more advanced (hawaiitic=mugearitic) stages of differentiation.Mafic lavas frequently carry ultramafic inclusions (see Munha et al., 1990),ranging from lithospheric harzburgites and lherzolites to dunite, wherlite, web-sterite and clinopyroxenite (DWWC) cumulates. Munha et al. (1990) demon-strated that DWWC Cr-spinelþ orthopyroxeneþ olivine cumulus assemblagescrystallised from Madeira alkaline magmas at 1.2–1.5 GPa, indicating that thefractionation processes were polibaric (see also Green and Ringwood,1967). Mantle heterogeneities, and differences in the degree of partial meltingand=or in the depth of magma extraction explain the main chemical differencesreported for the least evolved lavas (Mata et al., 1998; Geldmacher et al., 2000;Ribeiro et al., 2001).

Analytical procedures

The spinels studied were analysed at Centro de Geologia da Universidade deLisboa, on polished thin sections, using a 5 mm beam in a three channel wave-length dispersive JEOL-JCXA733+ electron microprobe, operated at an acceler-ating voltage of 18 kV and a beam current of 25 nA. Natural mineral and puremetal (Cr and V) standards were used. Compositions were calculated using theJEOL-ZAFO data-reduction routine program. Iron was determined as total iron(FeOT), and Fe3þ and Fe2þ calculated assuming ideal spinel stoichiometry andthat iron is the only element of variable oxidation state. Precision, as indicated byreplicate determinations on an in-house standard, is better than �2% for majorelements.

Systematics of Madeira lava spinels

Oxide minerals in Madeira lavas can be classified as belonging to the ilmenite andthe spinel series (Haggerty, 1976).

Ilmenite series members are very rare, a plausible consequence of the SiO2-undersaturated character of the Madeira lavas (e.g. Carmichael et al., 1974). This,and their extremely fine-grained nature, which precludes reliable microprobe ana-lyses, justify the decision to concentrate our attention on spinels.

Spinels are here considered, sensu lato, as cubic oxide minerals with AB2O4 asgeneral formula where A and B stand for tetrahedral and octahedral cation sites,respectively. Madeira lava spinels range from chromian-aluminous compositions tomembers of the magnetite – ulvospinel solid solution (titanomagnetites); Table 3and Fig. 1. In order to clarify the discussion of their petrogenetic significance,spinels were subdivided into 4 groups, according to textural relationships andchemical characteristics:

– Type I – chromian spinels included in olivine xenocrysts;– Type II – chromian spinels, magnesiochromites and chromites included in oli-

vine phenocryst cores;– Type III – titanomagnetites and Cr-titanomagnetites transitional to Type II;– Type IV – ‘‘titanomagnetites’’ included in ‘‘green diopside’’ xenocrysts.

Madeira Island alkaline lava spinels: petrogenetic implications 87

Tab

le1

.W

hole

rock

com

posi

tions

of

spin

el-b

eari

ng

lava

sfr

om

Madei

raIs

land

Sam

ple

no

M-2

9M

-46

M-6

7M

-12

0M

-14

1M

-21

3M

-22

9M

-24

5M

-25

5M

-29

8M

-31

3M

-35

1R

ock

typ

e�B

NB

BB

NP

CB

BB

NB

NB

NB

NB

N

SiO

2w

t%4

2.7

64

6.2

14

5.8

24

3.3

04

4.5

94

5.6

74

5.5

24

4.3

74

4.0

94

3.8

14

4.0

84

4.1

9T

iO2

3.1

32

.40

2.0

42

.54

2.6

12

.44

2.8

22

.58

2.6

92

.32

2.3

92

.79

Al 2

O3

13

.57

14

.03

13

.50

12

.52

13

.34

14

.02

15

.15

13

.63

14

.06

12

.27

12

.63

13

.69

Fe 2

O3(t

)1

3.5

81

1.7

11

3.1

91

3.4

61

3.4

81

2.3

71

2.9

11

3.1

51

2.9

61

2.4

51

2.4

41

2.9

1M

nO

0.1

80

.18

0.1

80

.19

0.1

90

.18

0.1

80

.20

0.1

90

.19

0.1

80

.19

Mg

O1

0.2

99

.62

11

.09

12

.15

11

.46

10

.12

8.2

51

1.5

61

0.2

11

4.0

51

2.9

71

0.2

0C

aO1

1.6

51

0.1

61

0.9

31

1.6

51

1.2

71

0.7

41

0.1

01

0.9

81

1.4

81

1.4

01

1.6

51

1.4

4N

a 2O

2.2

03

.26

3.1

13

.36

2.1

72

.68

3.4

32

.79

3.2

42

.81

3.3

63

.75

K2O

0.8

11

.10

0.6

10

.77

0.7

80

.80

1.0

20

.75

0.9

10

.84

0.8

30

.92

P2O

50

.50

0.6

80

.38

0.6

00

.51

0.6

30

.78

0.7

00

.80

0.4

80

.49

0.6

0L

.O.I

.1

.39

0.2

9n

.d.

0.3

10

.13

0.1

00

.31

0.0

50

.06

0.1

7n

.d.

0.0

1

Cu

pp

m8

64

88

26

26

55

64

05

56

85

65

96

9V

40

02

95

24

83

65

35

22

84

29

13

27

31

83

18

33

92

99

Cr

36

84

78

50

97

12

54

65

20

23

05

87

49

98

53

79

54

14

Nb

53

76

50

50

43

56

67

50

56

42

46

64

Th

3.4

7.9

2.6

3.1

2.8

n.a

.5

.53

.84

.22

.82

.75

.3L

a3

74

72

63

43

1n

.a.

56

41

48

30

30

55

Ce

72

95

55

73

62

n.a

.1

00

74

84

57

58

81

Sm

6.7

6.0

5.2

7.2

6.3

n.a

.8

.67

.07

.56

.25

.86

.9

� Bb

asal

t;B

Nb

asan

ite;

PC

pic

rob

asal

t


Tab

le2

.R

epre

sen

tati

vem

icro

pro

be

an

aly

ses

of

clin

opyr

oxe

ne

an

do

livi

ne

Cli

no

py

roxen

eO

liv

ine

Sam

ple

no

M-5

1M

-21

3M

-21

3M

-29

M-2

9M

-67

M-6

7M

-25

5M

-25

5M

-11

9T

yp

e�g

dp

h-c

ph

-rp

h-c

mp

h-c

ph

-rp

h-c

mxen

o-c

SiO

2w

t%5

1.8

85

1.4

14

7.2

8S

iO2

wt%

40

.86

39

.44

39

.94

35

.72

39

.30

35

.20

39

.43

TiO

20

.25

1.0

92

.14

TiO

20

.02

0.0

4n

.d.

0.0

80

.02

0.0

80

.04

Al 2

O3

1.5

53

.60

7.2

3A

l 2O

30

.05

0.0

10

.02

0.0

50

.01

0.0

70

.01

Cr 2

O3

n.d

.0

.23

0.0

4N

iO0

.34

0.1

60

.29

0.0

30

.24

0.0

10

.22

FeO

(t)

10

.05

6.4

77

.77

FeO

11

.63

17

.59

13

.93

35

.10

14

.76

36

.32

17

.11

Mn

O1

.45

0.2

10

.11

Mn

O0

.12

0.2

20

.18

0.5

50

.23

0.7

40

.22

Mg

O1

0.5

61

4.6

81

2.3

8M

gO

47

.35

42

.73

44

.95

27

.86

44

.69

26

.21

43

.29

CaO

22

.43

21

.87

22

.05

CaO

0.1

90

.23

0.2

20

.59

0.2

20

.79

0.0

9N

a 2O

1.2

60

.34

0.4

3T

ota

l1

00

.56

10

0.4

29

9.5

39

9.9

89

9.4

79

9.4

21

00

.41

K2O

0.0

20

.01

n.d

.T

ota

l9

9.4

59

9.9

19

9.4

3F

om

ol%

87

.89

81

.24

85

.19

58

.59

84

.37

56

.26

81

.85

� gd

gre

endio

psi

de;

ph

-c=r

phen

ocr

yst

al-c

ore=ri

m;

mg

rou

nd

mas

s;xe

no

-cxen

ocr

yst

al-c

ore


Tab

le3

a.M

icro

pro

be

an

aly

ses

of

chro

mia

nsp

inel

s(T

ypes

I,II

)a

nd

oli

vin

e-sp

inel

geo

ther

mo

ba

rom

etry

Sam

ple

no

M-2

9M

-11

9M

-29

M-4

6M

-67

M-1

20

M-1

41

M-2

13

M-2

45

M-2

98

M-3

51

M-3

81

Sp

inel

typ

eI

III

IIII

IIII

IIII

IIII

II

TiO

2w

t%2

.43

1.2

21

.43

3.6

42

.18

1.5

81

.98

2.4

61

.79

1.4

81

.53

3.2

6A

l 2O

33

6.3

74

3.1

92

3.4

92

0.5

92

2.1

21

8.6

42

4.5

92

2.4

03

0.1

72

3.0

81

7.2

22

6.2

1V

2O

30

.29

0.0

20

.10

0.1

7n

.d.

0.1

30

.27

0.2

60

.04

0.2

00

.16

0.1

7C

r 2O

31

3.9

41

0.3

53

2.7

02

3.0

62

6.4

93

1.8

82

7.4

12

9.3

42

4.1

33

3.9

84

0.9

01

8.3

2F

e 2O

3�

15

.55

13

.74

12

.01

17

.81

16

.81

17

.70

13

.16

12

.83

13

.17

11

.65

10

.54

18

.96

FeO

16

.95

16

.82

17

.06

26

.36

23

.11

16

.05

20

.20

24

.05

15

.94

15

.99

17

.95

19

.25

Mn

O0

.15

0.1

90

.23

0.4

80

.63

0.3

30

.20

0.4

00

.59

0.1

80

.20

0.2

2M

gO

14

.94

14

.98

12

.76

7.4

18

.89

12

.73

10

.94

8.7

11

4.1

91

3.5

51

1.7

21

2.4

2

To

tal

10

0.6

21

00

.51

99

.77

99

.51

10

0.2

29

9.0

39

8.7

51

00

.45

10

0.0

21

00

.12

10

0.2

29

8.8

1

Mg=(M

gþ

Fe2

þ)

0.6

11

0.6

13

0.5

71

0.3

34

0.4

07

0.5

86

0.4

91

0.3

92

0.6

13

0.6

02

0.5

38

0.5

35

Cr=

(Crþ

Al)

0.2

05

0.1

38

0.4

93

0.4

29

0.4

45

0.5

34

0.4

28

0.4

68

0.3

49

0.4

97

0.6

14

0.3

19

Mg

#(o

l)��

0.8

75

0.8

23

0.8

79

0.8

67

0.8

52

0.8

54

0.8

48

0.8

64

0.8

61

0.8

76

0.8

65

0.8

48

T1(o

l-sp

)�C

82

49

15

93

67

21

80

71

20

58

95

73

81

02

51

02

01

01

71

02

4T

2(M

g-i

n)�

C–

–1

22

11

16

11

17

01

17

81

18

51

19

21

20

31

21

71

21

61

16

4�

FM

Q��

2.3

1.5

1.4

2.4

1.9

1.6

1.3

1.7

1.3

1.2

0.9

2.0

�N

NO��

2.0

1.2

0.7

1.8

1.2

0.9

0.7

1.0

0.6

0.5

0.2

1.3

� Cal

cula

ted

on

the

bas

isof

spin

elst

oic

hio

met

ry;�� M

g=

(Mgþ

Fe2

þ)

inco

exis

tin

go

liv

ine.

T1

–B

all

ha

us

etal

.(1991)

oli

vin

e-sp

inel

tem

per

ature

s;T

2–

Hel

zan

dT

hro

nb

er(1

987)

liquid

us

tem

per

ature

s(s

eete

xt)

.�� l

og

fO2

rela

tive

toF

MQ

(O’N

eil,

19

87a)

and

NN

O(O

’Nei

l,1

98

7b

)bu

ffer

s.T

yp

eI

spin

elfO

2at

12

00� C

,1

.3G

Pa;

Ty

pe

IIsp

inel

fO2

atT

2an

d0.1

GP

a,co

rrec

ted

for�

log

(aS

iO2)¼�

0.2

rela

tive

too

l–o

px

–S

iO2

equil

ibri

a(s

eete

xt

for

calc

ula

tio

nm

eth

od

s)


Tab

le3

b.

Mic

ropro

be

analy

ses

of

tita

nom

agnet

ites

(Typ

eII

I,IV

spin

els)

and

ol-

spoxy

gen

geo

baro

met

ry

Cr–

Ti

mag

net

ite

Ti

mag

net

ite

Sam

ple

no

M-4

6M

-14

1M

-2M

-12

0M

-17

7M

-21

3M

-25

5M

-25

9M

-29

8M

-34

4M

-35

1M

-51

Spin

elty

pe

III=

ol-

rII

I=ol-

rII

I=m

III=

mII

I=m

III=

ph

III=

pl

III=

mII

I=m

III=

mII

I=m

IV=g

d

TiO

2w

t%1

1.8

31

7.0

22

2.2

21

2.5

41

9.6

22

4.2

32

6.6

22

3.7

52

2.4

92

1.0

01

7.5

46

.47

Al 2

O3

7.1

76

.38

2.5

12

.34

2.1

91

.78

2.7

82

.77

3.1

92

.64

1.6

60

.91

V2O

30

.29

0.7

40

.53

0.3

20

.03

0.5

60

.47

0.4

50

.43

0.0

30

.42

0.1

8C

r 2O

31

2.8

61

0.2

20

.09

0.0

30

.47

0.5

10

.07

0.0

80

.11

0.2

90

.09

0.0

1F

e 2O

3�

26

.74

18

.87

23

.82

44

.21

30

.58

20

.58

15

.26

19

.37

21

.89

25

.72

33

.86

55

.13

FeO

35

.17

43

.27

46

.86

32

.49

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Spinels crystallised from Madeira magmas

Type I, Type II and Type III spinels are interpreted as crystallising from Madeiramagmas.

Type I spinels occur in olivine xenocrysts that resulted from fragmentation ofDWWC xenoliths (Munha et al., 1990; Mata and Munha, 1995; Mata, 1996).Indeed, olivine xenocrysts hosting Type I spinels display intra-crystalline deforma-tion and have lower-CaO contents (�0.10 wt.%) than those of coexisting olivinephenocrysts (CaO� 0.2 wt.%; see also Table 2), which are characteristic featuresof olivine from DWWC xenoliths. DWWC xenoliths reflect high pressure cumu-lative processes that affected Madeira magmas (Munha et al., 1990), indicating that

Fig. 1. Compositional variations of Madeira Island spinels. A TiO2–FeO–1=2 (Fe2O3)wt.% relations. B Proportions of trivalent cations (Al–Fe3þ–Cr; p.f.u). Spinels fromhigh-pressure dunite, wherlite, websterite, clinopyroxenite cumulate xenoliths (DWWC)in Madeira lavas are also plotted for comparison


the olivine xenocrysts are cognates in the sense used, for example, by Shelley(1993), i.e. they were derived from within the same magma system as the host.

The origin of the olivine xenocrysts is also endorsed by the chemical charac-teristics of their Type I chromian spinel inclusions. Type I spinels are composi-tionally similar to those occurring in the DWWC cumulate xenoliths (see Figs. 2A,B and Table 3a), but contrast markedly from spinels included in olivine phenocryst

Fig. 2. A Mg=(Mgþ Fe2þ)¼Mg# vs Cr=(CrþAl), B Fe3þ=R3þ vs Cr=(CrþAl) and C Crvs Al (p.f.u) relationships in Madeira spinels. Note the existence of two distinct types ofdominant substitution mechanisms displayed in diagram C (see text for explanation).Symbols are the same as in Fig. 1


cores (Type II); Type II spinels being characterized by lower Al2O3 contents (17.22to 30.17 wt% vs. 33.74–43.19 wt%) and higher Cr2O3 concentrations (21.10 to40.90 wt% vs. 10.14–13.94 wt%) and Cr# values (>0.31 vs.<0.15).

Type II spinels have chromium-aluminous compositions (hereafter referred in abroad sense as chromian spinels; see Table 3a) and occur as euhedral to sub-euhedral crystals included at various positions in the cores of olivine phenocrysts.The absence of chromium-aluminous spinels in the groundmass and the limitationof their occurrence to the interior of olivine phenocrysts, suggest that Type II spinelcrystallisation was restricted to the first stages of magmatic evolution and that itsprecipitation preceded that of olivine, pointing to the chromite-saturated characterof the more primitive magmas.

The occurrence of Type III spinels is more diverse. The great majority of TypeIII spinels are confined to the groundmass, but some crystals were also detected asinclusions in olivine (rims), pyroxene and plagioclase. Occasionally, both types ofspinels occur inside the same olivine phenocryst; in these cases, Type II spinels arerestricted to the olivine core zones, whereas those of Type III predominate near therims (Table 3b and Plate I). Such mineralogical features clearly emphasize the needfor combined chemical and textural analysis in order to understand their chemicalevolution.

Chromium spinel compositional variations essentially obey the effects of octa-hedral Cr$Al substitution, whereas 2Fe3þ $ Fe2þ þTi4þ exchange dominatesthe magnetite-ulvospinel series. That the Cr$Al substitution is not the only octa-hedral exchange mechanism characterising Madeira spinels is clear from Fig. 2C,where two distinct Cr–Al trends are displayed (note that the trends in Fig. 2C areequivalent to ‘‘Cr–Al’’ and ‘‘Fe–Ti’’ trends defined by Barnes and Roeder (2001)on the Cr–Al–Fe3þ and Fe3þ# vs. Fe2þ# diagrams).

The Cr–Al negative correlation characteristic of Type I and Type II spinelsoccurring inside olivine phenocryst cores (Fig. 2C) suggests that Cr$Al sub-stitution is the dominant mechanism accounting for their chemical variability.Accordingly, they display a trend parallel to the octahedral CrþAl saturationline (i.e. CrþAl¼ 16); however, CrþAl ranges from 10.722 up to 13.369 (Fig.2C) suggesting the presence of Ti (0.202– 0.892 p.f.u.) and Fe3þ (2.106–4.120 p.f.u.) in Type II spinels. Their high Ti contents are identical to thosereported for oceanic island basalt Cr-spinels (e.g. Arai, 1992; Kamenetsky et al.,2001).

In contrast to the former spinels, those of Type III display a Cr–Al positivecorrelation in Fig. 2C and have lower CrþAl contents (0.195–8.562 p.f.u.), indi-cating that substitution mechanisms other than Crþ Fe$AlþMg were also sig-nificant. The trend towards lower Cr and Al values (and increasing Ti, Fe3þ andFe2þ occupancy) in Type III spinels (Figs. 2C, D) suggests progressive Cr and Aldepletion in Madeira host magmas, possibly resulting from contemporaneous frac-tionation of clinopyroxene which may include significant amounts of Cr2O3 andAl2O3 under the low aSiO2

and high fO2 conditions characteristic of Madeira mag-mas (see below and also Mata and Munha, 1998).

Mg vs. Fe2þ relations (Fig. 3A) reinforce the difference between Type II (and I)and Type III spinels, revealing two distinct groups, although for spinels with Fe2þ

between 5 and 6 p.f.u. the difference is not so evident. Whereas those of Type II


(and I) display almost constant MgþFe2þ sums (�8) characteristic ofnormal spinels, Type III MgþFe2þ values are higher and more variable; higherMgþ Fe2þ values reflect the predominance of magnetite – ulvospinel components,which confer an inverse structure to Type III spinels. Different trends are alsodepicted for spinel inclusions within the same olivine phenocryst in Fig. 3B,

Plate I. A, B Typical texturalrelationships of spinel inclu-sions in olivine phenocrystals;core inclusions are of Cr–Al,Type II, spinels whereas chro-mian–titanomagnetite, TypeIII spinels, occur in the olivinerim and groundmass (see B).Scale bars¼ 1 mm. C Type Ispinels inclusions in olivinexenocrystal displaying defor-mation bands (crossed nicols).Scale bar¼ 1.3 mm


clearly indicating that the compositional contrasts between Type II and Type IIIspinels should reflect the evolving nature of the host magma.

Some of the Type III spinels have intermediate compositions (Al¼ 2.149 to4.890 p.f.u.; Cr¼ 1.682 to 4.064 p.f.u.; see Table 3b) between chromian spinels andthose plotting in the field of magnetite-ulvospinel solid solution series. Given that(compositionally) intermediate spinels display the same trends as shown by mem-bers of the magnetite-ulvospinel solid solution series, we classify them as titano-magnetites with the prefix chromian owing its high Cr2O3 contents (chromiantitanomagnetites).

Type IV spinels

Textural relationships characterising Type IV spinels as inclusions in green diop-side xenocrysts show they are independent from the other groups.

Type IV spinel green diopside hosts have chemical characteristics clearlydistinct from those typical of Al–Ti diopside phenocrysts in Madeira lavas

Fig. 3. Mg vs Fe2þ (p.f.u.) (A) and Mgþ Fe2þ (p.f.u.) vs Mg# (B) relationships forMadeira spinels. Circumferences in diagram (B) represent spinels occurring inside thesame olivine grain


(see Table 2). Madeira green diopsides are similar to low-Al clinopyroxenesdescribed by Duda and Schmincke (1985) and Dobosi (1989) as mantle xeno-crysts. They are identical to metasomatised clinopyroxenes in residual harzbur-gite xenoliths from Madeira Island (Munha et al., 1990). Given thesesimilarities, green diopside xenocrysts in Madeira lavas were interpreted byMata and Munha (1998) as dismembered Upper Mantle xenoliths, and theirparticular composition was attributed to mantle metasomatic processes (see alsoPilet et al., 2002).

Chemical data on Type IV spinels (Table 3b) supports the xenolithic characterattributed to their green diopsides hosts. Accordingly, Type IV spinels are differentfrom those occurring in Madeira lavas clinopyroxene phenocrysts, as they havehigher Fe3þ=Fe2þ and MnO contents.

Chromian spinel – titanomagnetite solid solution in Madeira lavas

A prominent solvus inside the spinel compositional space, particularly for inter-mediate compositions between chromian aluminous compositions and members ofmagnetite-ulvospinel solid solution, has been suggested (e.g. Sack and Ghiorso,1991a).

However, in Madeira lavas, we detected chromian titanomagnetites (see Fig. 1Band Table 3b) included in olivine phenocrysts suggesting their crystallisation frommoderately evolved melts. Different types of spinels are characterized by distinctmodes of occurrence; type II Cr–Al spinels are included in the core of olivinephenocrysts, members of magnetite-ulvospinel solid solution series in the ground-mass, and chromian-titanomagnetites included in the rims of olivine phenocrysts.These observations indicate that the entire compositional range shown by spinelscrystallising from Madeira magmas reflects continuously changing crystallisationparameters (see also Wilkinson and Hensel, 1988). The inference is supported byexperimental data (Hill and Roeder, 1974), implying solid solution between chro-mite and titanomagnetite in the studied lavas.

Indeed, Hill and Roeder (1974) demonstrated the occurrence of extensive chro-mite-titanomagnetite solid solution for fO2 conditions above that defined by theNNO buffer whereas, at lower oxygen fugacities, chromite crystallisation should beinterrupted by the precipitation of clinopyroxene. Under these conditions, spinelprecipitation is only re-initiated when melt Fe2O3 and TiO2 concentrations becomesufficiently high resulting in a marked compositional discontinuity of spinels.Sack (1982) demonstrated a strong temperature dependence of the activity coeffi-cients for (Fe,Mg)2TiO4 and (Fe,Mg)Al2O4 components in chromian spinel duringtransition to titanomagnetite, providing a plausible explanation for the scarcity ofrepresentative analyses that are intermediate between chromian spinel and titano-magnetite in Fig. 1B. Taking this into account, the data presented for Madeiranspinels strongly suggests significant solid solution from chromian spinel to titano-magnetite. This observation, coupled with Hill and Roeder (1974) experimentaldata, implies that Madeira magmas should have evolved under relatively oxidisingconditions. Otherwise, more reduced conditions would drive the spinel composi-tional trends through spaces, inside the spinel prism, that should be characterisedby pronounced miscibility gaps (see Sack and Ghiorso, 1991a, b).


fO2 and P-T-aSiO2conditions during crystallisation of Types I

and II spinels: the initial stages of Madeira magma evolution

Relatively high oxygen fugacities during crystallisation of Types I and II spinelscan be assessed by using the olivine-orthopyroxene-spinel oxygen barometer(O’Neil and Wall, 1987; Matiolli and Wood, 1988; Ballhaus et al., 1990, 1991;Wood, 1990; Nell and Wood, 1991). Reversed re-equilibration experiments of Wood(1990) and Ballhaus et al. (1991) demonstrate that Nell-Wood (Wood, 1990; Nelland Wood, 1991) and Ballhaus et al. (1991) versions of the oxygen barometer yieldconsistent results (Wood, 1991), providing better overall agreement between cal-culated and known fO2 compared with the other available models. Accordingly,results reported in Table 3a were calculated on the basis of Ballhaus et al. (1991)calibration of the oxygen barometer; they also incorporate the necessary correc-tions due to orthopyroxene undersaturation (see Ballhaus et al., 1991) in Madeiralavas. Application of the Nell-Wood (Wood, 1990; Nell and Wood, 1991) model toMadeira spinels yields essentially the same results as reported here and the con-clusions are not affected by the choice of models. In order to facilitate comparisonbetween data we have compared them all to the FMQ equilibrium (O’Neil, 1987a;see Table 3a and Fig. 4), because relative oxygen fugacities [�log( fO2)FMQ¼log( fO2)sample� log( fO2)FMQ] are largely insensitive to errors in calculated

Fig. 4. Oxidation states of Madeira spinels plotted against Cr=(CrþAl) in spinel. Filledand empty triangles represent Type I and DWWC high-pressure spinels, respectively;circles – Type II liquidus spinels; filled squares – spinels in harzburgite=lherzolitelithospheric xenoliths (data from Munha et al., 1990) exhibiting the typical ‘‘oxidationtrend’’ of Ballhaus (1993). Calculations at 1200 �C, 1.3 GPa for Type I=DWWC spinels,and liquidus T (Helz and Thornber, 1987), 0.1 GPa, corrected for the absence of orthopyr-oxene in Madeira lavas (see text), for Type II spinels. Error bars illustrate typical deviationsfrom calculated fO2 due to thermal resetting


temperatures (�100 �C shift relative fO2 by less than �0.2 log units) and have onlya small dependence on pressure (�0.3 log units=�1 GPa).

When calculating relative oxygen fugacities from Type I spinelsþ olivinexenocrysts we have assumed that they do indeed represent dismembered DWWCxenoliths; thus, aSiO2

-P-T were preset assuming olivine-orthopyroxene saturation(O’Neil and Wall, 1987; Ballhaus et al., 1991) at 1.3 GPa, 1200 �C, which are thetypical conditions inferred by Munha et al. (1990) to have prevailed during ‘‘pri-mary’’ crystallisation of DWWC cumulates.

Type II spinel crystallisation conditions were estimated by assuming that spi-nels (included in the core zones of olivine phenocrysts) probably crystallised atpressures-temperatures not significantly different from those characterising theirhost phenocrysts at saturation. Liquidus temperatures were calculated according toHelz and Thornber (1987) geothermometer, by using whole rock compositions(Table 1) that were previously readjusted (by iteration with fO2 estimates) to properFe2O3=FeO values (Klinic et al., 1983; see also Ottonello et al., 2001) and cor-rected for potential olivine accumulation (e.g. Carmichael, 1991). Estimated liqui-dus temperatures display a limited range, from 1161 �C to 1221 �C (Table 3a). Incontrast, temperatures derived from Fe–Mg exchange between Type II spinels andcoexisting olivine phenocrysts (Ballhaus et al., 1991) show a much wider range(721 �C–1205 �C; Table 3a) and indicate that most spinels are reset to lower thanliquidus temperatures; errors in calculated fO2 because of Fe,Mg readjustmentduring thermal resetting are, however, likely to have been small (less than �0.2 logunits relative to FMQ according to Ballhaus, 1993). Geobarometric estimates wereobtained by comparing DWWC and phenocryst olivine CaO contents (see Munha

et al., 1990 and Table 2;D

ol-phenoc=cpx

Ca

Dol-DWWC=cpx

Ca

>1:5), implying �P< �� Tð�KÞ

562 lnð1:5Þ

��

�1.1 GPa at T¼ 1160–1220 �C (Kohler and Brey, 1990), which is supported by theaverage P¼ 0.14� 0.04 GPa of Nimis (1999) single-clinopyroxene geobarometerresults for Madeira clinopyroxene phenocrysts. Accordingly, we suggest that crys-tallisation of Type II spinels occurred within shallow magma chamber(s), probablyat pressures that did not significantly exceed 0.1 GPa. Finally, amelt

SiO2(and the

derived correction on calculated fO2) were estimated from Madeira magma com-positions at P-T conditions specified above. Thus, whole rock compositions(SiO2¼ 42.76–46.21 wt%; Table 1) together with Ghiorso’s silicate melt solutionmodel (Ghiorso et al., 1983; updated by Ghiorso and Sack, 1995), fix amelt

SiO2for

Madeira lavas at 0.33� 0.03; this is about 0.2 log units bellow that defined byolivine-orthopyroxene equilibrium (T¼ 1160–1220 �C, P¼ 0.1 GPa; O’Neil andWall, 1987) and the corresponding correction (see Ballhaus et al., 1991) has beenapplied throughout calculations on estimating fO2 from Type II spinels.

Even at such low aSiO2, all Type II spinels indicate fO2 conditions significantly

more oxidised than FMQ (�log(fO2)FMQ¼ 1.6� 0.5), identical to high-pressureType I spinels in cognate DWWC cumulate xenoliths (Munha et al., 1990) andolivine xenocrysts (�log(fO2)FMQ¼ 1.5� 0.6; see Table 3a and Fig. 4). Type IIspinel �log(fO2)FMQ values increase with decreasing MgO-in host whole rock(Helz and Thornber, 1987) liquidus temperatures (R2¼ 0.72; see Table 3a), sug-gesting that early low-pressure fractionation of Madeira magmas proceeded under


Fig

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increasing oxidising conditions. As olivine phenocrysts, with little or no ferric iron,largely predominate over Type II spinels in the early crystallisation of Madeiramagmas, it is feasible that residual liquids would have become increasinglyenriched in Fe3þ=Fe2þ, and in consequence enhanced their relative oxygen fuga-city (Christie et al., 1986; Snyder et al., 1993). However, high �log( fO2)FMQ valuesexhibited by cognate DWWC high-pressure cumulates and the occurrence of oxi-dised harzburgite=lherzolite lithospheric xenoliths (see Fig. 4), indicate that therelatively high redox state of Madeira magmas should have originally beenacquired at depth. Moreover, the incompatible element ratio (Nb=Th, Ce=La,La=Sm) variations shown in Fig. 5 cannot be fully explained by olivineþspinel� orthopyroxene fractionation processes. Thus, it is likely that some of thecovariance between element ratios and �log( fO2)FMQ values exhibited by Madeiramagmas (Fig. 5) was inherited from their respective source(s); such geochemicalfeatures probably reflect the development of mantle heterogeneities (e.g. Matiolliet al., 1989; Amundsen and Neuman, 1992), as well as Dsolid=liquid variations duringpartial melting (e.g. Canil, 1999, 2002), induced by oxidising metasomatism oper-ating at the source of Madeira magmas.

Despite subtle variations, the overall oxygen barometry data firmly support con-clusions reached previously, that oxidation states of Madeira magmas were high.

Thus, experimental work (Hill and Roeder, 1974), geothermobarometry (Helzand Thornber, 1987; Ballhaus et al., 1991) and spinel mineral chemistry (thisstudy) are mutually congruent, indicating that the early stages of crystallisationin Madeira magmas proceeded under relatively high oxidising conditions. Calcu-lated oxygen fugacities clearly exceed those defined by the NNO buffer(0.2 � �log( fO2)NNO � 1.8; see Table 3a), even exceeding values reported byBallhaus (1993) for ocean island basalts elsewhere.

fO2: influence on the crystallisation sequencein magnetite – ulvospinel solid solution

In the preceding discussion it has been claimed that early crystallisation of Cr–Al(type II) spinelsþ olivine in Madeira magmas has proceeded under increasingoxidation conditions. However, at more advanced stages of (olivineþclinopyroxeneþ spinel� plagioclase) crystallisation, Cr–Al spinels gave way totitanomagnetites (type III spinels) that display significant increase in ulvospinelcomponent with diminishing Mg# (Fig. 6); moreover, there is a positive correlationbetween (maximum) ulvospinel (35–77 mol%) and fayalite (15–49 mol%) contentsin coexisting titanomagnetite and olivine (see Table 3b), suggesting that the rela-tive oxidation in Madeira magmas has decreased during the late stages of crystal-lisation (Frost and Lindsley, 1991). To access this apparent inflection on the fO2

crystallisation trend we have assumed that (for each sample) the ulvospinel-richestFe–Ti oxides are in equilibrium with the most Fe rich silicates that are the (seeFrost and Lindsley, 1992) and estimated �log(fO2)FMQ (from coexisting olivine-titanomagnetite mineral analyses) by using the equation:

� log ð fO2ÞFMQ ¼ 2 log aðFe3O4Þspinel � 3 log aðFe2SiO4Þolivine

þ 3 log aðSiO2Þmelt ð1Þ


where the standard state thermodynamic properties and solid solution activityterms were computed from the equations embodied in Andersen et al. (1993)QUILF software. Although, P, T and amelt

SiO2variations are (a priori) unknowns in

Eq. (1), the constancy of ameltSiO2

in Madeira magmas (at 1000 �C�T�1200 �C, ameltSiO2

remains at 0.3–0.4 for Madeira basalt to mugearite compositions; (Ghiorso andSack, 1995) and the very small dependence of �log( fO2)FMQ on pressure (less than�0.01 log units=�1 GPa) and temperature (�0.1 log units=�100 �C) allow calcu-lations to be performed at fixed P¼ 0.01 GPa, T¼ 1100 �C and amelt

SiO2¼ 0.35.

The results in Fig. 7 (see also Table 3b) are consistent with previous type II spineloxygen barometry determinations, supported by titanomagnetite-ilmenite fO2

(�1.6<�log( fO2)FMQ<1.2; unpublished) data from Madeira essexitic rocks thatcorroborate the estimated trend of reducing conditions during late-stage differen-tiation. Titanomagnetites in Madeiran lavas indicate oxygen fugacities that vary

Fig. 6. A Ti (p.f.u.) vs Mg# and B Fe3þ=Fe2þ vs Mg#¼Mg=(Mgþ Fe2þ) relationships forType III spinels


over three orders of magnitude (Fig. 7); �log(fO2)FMQ values decrease withincreasing olivine fayalite contents (an evolution index), from �log(fO2)FMQ¼ 2.22.2 at X(Fe2SiO4)olivine¼ 0.15 to �log(fO2)FMQ¼�1.1 at X(Fe2SiO4)olivine¼ 0.49.The situation depicted in Fig. 7 is reminiscent of that found in many tholeiiticmagma bodies, where there is a trend of rising relative oxygen fugacity duringearly silicate fractionation followed by falling relative oxygen fugacity after Fe–Tioxide saturation (e.g., Byers et al., 1984; Christie et al., 1986; Juster et al., 1989;Snyder et al., 1993). Data in Fig. 7 suggests that titanomagnetite saturation inMadeira lavas occurred at (FeO=MgO)melt � 1.0–1.1 wt% (calculated fromX(Fe2SiO4)olivine � 0.15–0.16 and Kd (Fe=Mg)ol=liq¼ 0.30; Roeder and Esmile,1970), which is much lower than in more reduced, iron enriched, tholeiiticmagmas.

It is widely accepted that oxygen fugacity exerts an important control on Fe–Tioxide stability (e.g. Hill and Roeder, 1974; Thy and Lofgren, 1994). Moreover,Snyder et al. (1993) and Toplis and Carrol (1995) experimentally demonstratedthe strong fO2 dependence on the order of appearance and relative abundance ofFe–Ti oxide minerals. These studies demonstrated that above the QFM buffer,crystallisation of (magnetite-ulvospinel)ss should precede the appearance of (ilme-nite-hematite)ss; this sequence being inverted at more reduced conditions.

Clearly, under constant ameltTiO2

, rising fO2 values will favour the early crystal-lisation of magnetite rich titanomagnetites; then, crystallisation of such Fe3þ richspinels will rapidly decrease Fe3þ=Fe2þ ratios and, consequently, the relative oxy-

Fig. 7. Oxidation states for Type III (empty squares) spinels plotted against fayalitemolar fractions in coexisting olivines. Calculations assume 1100 �C, 0.01 GPa andaSiO2

(melt)¼ 0.35 (see text). Oxidation states for Type II spinels are also plotted forcomparison. Thick vertical arrow illustrates the trend of decreasing oxidation estimatedfrom titanomagnetite–ilmenite fO2 (unpublished) data on Madeira granular (essexite)rocks. Error bars illustrate typical deviations on calculated �log(fO2)FMQ due to tem-perature and aSiO2

variations


gen fugacity in residual melts. In tholeiitic magmas titanomagnetite saturation willdrive the residual melt to the stability field of ilmenite (e.g. Snyder et al., 1993), butunder the low amelt

SiO2prevailing in Madeira (and other alkaline) magmas ilmenite

crystallisation is suppressed (e.g. Carmichael et al., 1974); thus, TiO2 will beexclusively partitioned between Ti-bearing silicates (e.g., Ti-augite) and spinel,resulting in a net increase in the rate of ulvospinel enrichment during subsequentcrystallisation of titanomagnetite. The overall situation may be represented byreactions (2) and (3)

2 FeTiO3ilmenite

þFe2SiO4olivine

¼ 2Fe2TiO4spinel

þSiO2melt

ðsee Lindsley and Frost; 1992Þ ð2Þ

CaAl2Si2O8plagioclase

þ 2Fe3O4spinel

þ 2TiO2melt

¼ CaTiAl2O6pyroxene

þ 2Fe2SiO4olivine

þ Fe2TiO4spinel

þO2 ð3Þ

in which low silica and oxygen activities will drive reactions to the right; withelimination of ilmenite, a solid assemblage of plagioclase, Ti-augite, fayalitic oli-vine and ulvospinel enriched titanomagnetite will result.

Thus, high oxygen fugacities estimated for the early stages of crystallisation inMadeira magmas (above NNO; see Fig. 4) are responsible for the titanomagnetitecrystallisation sequence, which is characterized by significant ulvospinel enrich-ment. It should be noticed that for similar oxygen fugacity, an increase inFe3þ=Fe2þ will be favoured by high degree of melt depolymerization (Mysen,1990), i.e. high ratios of non-bridging oxygens vs. bridging oxygens per tetrahed-rally coordinate cations; therefore, the SiO2 undersaturated character of Madeiramagmas, (favouring melt depolymerization), cannot be ruled out as co-responsiblefor the magnetite rich spinel precipitation.

In conclusion, we suggest that magnetite-ulvospinel trends described forMadeira spinels may be explained by a two step process. Initially high fO2 leadsto preferential precipitation of magnetite; then, the decrease in oxygen fugacitywill favour the incorporation of ulvospinel component into crystallising titanomag-netite. Under more reduced conditions, and or for more silica rich magmas, initialenrichment in Fe2TiO4 component is favoured. This leads to titanomagnetite com-positional variations marked by an evolutionary trend from ulvospinel rich tomagnetite rich.

Pressure effects on chromian-spinel compositions

Melt composition, oxygen fugacity, pressure and temperature are thought toinfluence the chromian-spinel compositions (e.g. Roeder and Reynolds, 1991;Poustovetov and Roeder, 2000; Kamenetsky et al., 2001). Given the considerableoverlap in Mg# and Feþ3=Feþ2 values for Type I and Type II spinels (Figs. 2A, B),neither the fractionation extent nor the oxidation state of the liquids in which theyprecipitated provide adequate variable parameters to explain the observed differ-ences in their Cr and Al contents; for the same Mg# values Type II spinels havehigher Cr# (higher Cr and lower Al contents) than Type I spinels (Fig. 2A).


Temperature related compositional differences are also likely to have been small,since the initial crystallisation temperatures of DWWC xenoliths (�1200 �C;Munha et al., 1990), which include Type I spinels, were similar to those calculatedfor Type II spinels (Table 3a).

However, Type I spinels must have crystallised at higher pressures (1.2–1.5 GPa; Munha et al., 1990) than Type II spinels (probably�0.1 GPa); thus, itis relevant to discuss the influence of pressure on the compositional differencesobserved.

There is some experimental evidence for Al increase and Cr decrease in spinelswith pressure (Green et al., 1971; Jaques and Green, 1979; Murck and Campbell,1986; Thy, 1991); however, the issue is still a matter of debate. Under sub-solidusconditions (xenoliths) spinel Al content variations are interpreted as the result ofequilibria between Al-bearing pyroxenes, olivine and Mg-Al spinels,

CaMgSi2O6 þ MgAl2O4 ¼ Mg2SiO4 þ CaAlAlSiO6 ð4Þ

whereas in lavas spinel Al contents are related to changes in the solubility of Cr-rich minerals (Allan et al., 1988),

ðMgAl2O4Þsp þ ðFeCr2O4Þliq ¼ ðFeCr2O4Þsp þ ðMgAl2O4Þliq ð5Þ

In contrast to previous experimental indications (op cit) �V values for reaction (4)(Robie et al., 1978; Holland and Powell, 1998) seem to be very small (see Fig. 8);moreover, Roeder and Reynolds (1991) demonstrated that pressure variations from0.01 to 1.0 GPa have only marginal effects on Cr=(CrþAl) contents of spinel

Fig. 8. P-T-aMgAl2O4relations for Al–clinopyroxeneþ olivineþ spinel equilibria (calcu-

lated by using Holland and Powell (1998) thermodynamic data base and PM-18 DWWCxenolith mineral chemistry reported by Munha et al., 1990). Note that decreasing tempera-tures could lead to increase in Al contents of spinel in equilibrium with Al–pyroxene;however, unless cooling rates are very slow, the reaction progress will typically remain farfrom equilibrium due to the sluggishness of Al diffusion in clinopyroxene


crystallising from basaltic melts. Thus, compositional effects on Al–Cr contentsare likely have been negligible for the pressure range encompassed by theMadeiran spinels.

Nevertheless, pressure may exert a crucial influence on spinel compositionthrough modification of the crystallisation sequence (and consequent changes ofmelt Al–Cr contents; see also Barnes and Roeder, 2001). Indeed, Green andRingwood (1967) demonstrated that important changes in solid-liquid phase rela-tions should take place between 0.9 GPa and 1.35 GPa, with the olivine field con-tracting while the pyroxene fields undergo rapid expansion (see also Cox et al.,1979); these same authors also remarked that, at 1.2–1.5 GPa, orthopyroxene couldbe the primary liquidus phase in alkali basalt magmas. Accordingly, the observedcrystallisation sequence in DWWC xenoliths (Munha et al., 1990) is (Type I)spinelsþ orthopyroxene, followed by olivine; in contrast, lower pressure crystal-lisation of Type II spinels occurred simultaneously with olivine, being followed byclinopyroxene. Given solid-liquid partition coefficients, Dol

Cr DpxCr (e.g. Green,

1994), it is to be expected that, for the same extent of fractionation (as measuredby #Mg values), high-pressure (>0.9 GPa) crystallising magmas should have higherAl=Cr ratios when compared to their low-pressure equivalents. For a similar tem-perature and oxygen fugacity range of conditions, these differences in melt Al=Crvalues will correlate with coexisting spinel compositions (Maurel and Maurel,1982a, b; Roeder and Reynolds, 1991); thus, explaining the observed compositionalcontrasts between high-pressure, low-Cr #, Type I spinels and low-pressure, high-Cr #, Type II spinels in Madeiran magmas.

These features explain the (‘‘old’’) general observation that high-pressure spi-nels tend to be Al-rich and, if properly used, may provide a diagnostic tool forpolybaric crystallisation in alkali basaltic magmas elsewhere.

Conclusions and implications

1) Four discrete groups of spinels were observed: I) Al-rich (Cr#<0.15) chro-mian-spinel; II) Al-poor (Cr#>0.31) chromian-spinel, magnesiochromite andchromite; III) chromian-titanomagnetite (8<Cr2O3 wt%<19) and titanomag-netite; and IV) rare, TiO2-poor (<9 wt%), MnO-rich (<1 wt%), titanomagnetiteinclusions in green-core clinopyroxene xenocrysts.

2) Spinel groups I to III largely reflect the physical conditions and extent offractionation of the crystallising Madeiran magmas: Al-rich chromian-spinels(51<Mg#<62) are found as inclusions in olivine xenocrysts from high-pres-sure, cognate, ultramafic cumulates, Al-poor chromian-spinel=magnesio-chromite=chromite (30<Mg#<61) occur in olivine phenocryst cores,chromian-titanomagnetites (11<Mg#<34) in olivine phenocryst rims, andchromian-titanomagnetites to titanomagnetites occur in clinopyroxene pheno-crysts (6<Mg#<25) and in the groundmass (3<Mg#<36).

3) Extensive substitution of FeþTi for CrþAl as chromian-spinel – ulvospinel –magnetite solid solution is widespread in Madeiran spinels. Experimentalevidence suggests that the chemical continuum of spinels requires relativelyoxidising conditions during crystallisation; spinel Fe2O3=FeO ratios and wholerock chemistry indicate that group II (liquidus) spinels crystallised from


Madeiran magmas under fO2 values above those of the NNO buffer (0.2��log( fO2)NNO�1.8).

4) Relatively high fO2, coupled with low aSiO2of magmas, also explain the

observed ulvospinel enrichment evolutionary trend of Madeiran titanomagne-tites; contrary to the trend usually considered (e.g. Carmichael and Ghiorso,1990) more common.

5) Al contents of liquidus spinels are largely dependent on the Al content of themelt; high-Al spinels deviating from this relationship are interpreted to havecrystallised at high-pressures. It is shown here that differences in Al=Cr ratioAl=Cr between liquidus, chromian-spinels occurring in high-pressure olivinecognate xenocrysts and those included in the core zones of olivine phenocrystsis not a direct consequence of the influence of pressure on the stabilisationof high-Al spinels. Such differences probably reflect compositional effects ofdistinct crystallisation sequences during polibaric fractionation of Madeiramagmas.

Acknowledgements

This work was supported by FCT (Portugal) through Research Project (POCA – Centro deGeologia da Universidade de Lisboa; POCTI=FEDER). We are indebted to T. Palacios forkeeping the electron microprobe in running conditions. Thoughtful reviews by Professors C.Ballhaus, E. Stumpfl and an anonymous referee have led to a substantial clarification of theideas expressed here. D. Costa, M. L. Duarte, I. M. Almeida, L. T. Martins, A. Pinto and S.Martins are acknowledged for their support in various phases of this work.

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Authors’ address: J. Mata (e-mail: [email protected]) and J. Munha (e-mail: [email protected]),Centro de Geologia, Departamento de Geologia da Universidade de Lisboa, Faculdade deCieencias, Bloco C2, 5� Piso, 1749-016 Lisboa, Portugal.


madeira island alkaline lava spinels: petrogenetic implications

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