Author's personal copy
Earth and Planetary Science Letters 286 (2009) 208–218
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Earth and Planetary Science Letters
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Zircon and titanite recording 1.5 million years of magma accretion, crystallization and
initial cooling in a composite pluton (southern Adamello batholith, northern Italy)
Urs Schaltegger a,⁎, Peter Brack b, Maria Ovtcharova a, Irena Peytcheva a,b,c, Blair Schoene a, Andreas Stracke b,
Marta Marocchi d, Giuseppe M. Bargossi d
a
Section des Sciences de la Terre et de l'environnement, Université de Genève, 1205 Genève, Switzerland
Departement Erdwissenschaften, ETH Zürich, 8092 Zürich, Switzerland
Geological Institute, Bulgarian Academy of Science, Sofia, Bulgaria
d
Dipartimento di Scienze della Terra e Geologico-Ambientali, Università di Bologna, P.zza P.ta S. Donato,1-40126, Bologna, Italy
b
c
a r t i c l e
i n f o
Article history:
Received 4 February 2009
Received in revised form 15 June 2009
Accepted 20 June 2009
Available online 17 July 2009
Editor: R.W. Carlson
Keywords:
Adamello batholith
U–Pb dating
zircon
titanite
hafnium isotopes
zircon residence time
a b s t r a c t
The southern part of the Adamello batholith (the so-called “Re di Castello unit”) is an example of a composite
pluton, ranging from gabbro to granodiorite in composition. U–Pb dating of single-zircon crystals from four
tonalitic to granodioritic lithologies reveals that zircon crystallization is protracted in all studied lithologies,
showing apparent durations of growth between 90 and 700 ka. The youngest zircons crystallized near the
solidus and yield identical or slightly older ages than titanite. The formation of these autocrystic zircons is
considered to approximate the age of emplacement of the melt and its final crystallization, in contrast to
antecrystic zircons present in the same sample, which had formed earlier in the magmatic column or were
derived from re-mobilized earlier magma. The autocryst-derived “emplacement” ages range from 42.43 ±
0.09 Ma to 40.90 ± 0.05 Ma, recording 1.5 Ma of intrusion and crystallization history. We anticipate that
extended periods of zircon crystallization may be common in silicic rocks, whereas the zircons from residual
melts from initially undersaturated mafic liquids should yield far more precise emplacement ages within our
present analytical uncertainties of 0.1–0.2% in 206Pb/238U age. Decreasing Th/U ratios of dated zircons within
one melt batch document the depletion of the residual melt portion in Th due to the contemporaneous
crystallization of titanite. Preliminary Hf isotopic compositions of the dated zircon grains suggest that the
early stage melts of the southern Re di Castello unit represent hybrid melts with an important crustal
component (εHf between − 2.8 and +3.0). Subsequently emplaced melts are more juvenile at εHf values at
+6.4 to +8.9 and may thus reflect the addition of large volumes of mafic melt to the magmatic system.
© 2009 Elsevier B.V. All rights reserved.
1. Introduction
Middle-to-upper crustal plutons provide a complex integrated
picture of tens of thousands to million year-long evolution of melt
accretion, melt depletion and crystallization. This often complex history
can be reconstructed from analyzing magmatic mineral textures and
syn-magmatic deformation structures at micro-to-macroscale. A growing body of age determinations provide evidence that the emplacement
of even apparently homogeneous plutons occurs by the sequential
injection of multiple magma pulses, as well as lateral or vertical
accretion of magma batches, both in the subvolcanic environment
(Bacon et al., 2007, Charlier et al., 2008) and in the middle-to-upper
crustal level (Coleman et al., 2004; Matzel et al., 2006; Michel et al.,
2008). Zircon is a suitable mineral to reconstruct these processes, since it
records crystallization at different stages of magmatic evolution: in the
source, during the ascent and at the final level of emplacement, starting
⁎ Corresponding author. Tel.: +41 22 379 66 38; fax: +41 22 379 32 10.
E-mail address: [email protected] (U. Schaltegger).
0012-821X/$ – see front matter © 2009 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2009.06.028
at the moment the melt has reached zircon saturation (e.g., Belousova
et al., 2006). The emplacement of magmatic liquids in the intermediate
and upper crust occurs over variable timescales: batholiths can grow
over hundred thousands to millions of years (e.g., Matzel et al., 2006)
and be composed of individual magmatic pulses that intruded over
much shorter time spans, i.e. in the range of several tens to hundred
thousand years (e.g., Michel et al., 2008).
Can we quantify the timescale during which an individual pluton is
assembled? A challenge of modern high-precision U–Pb geochronology
is to resolve the emplacement ages of individual magma pulses within a
larger magmatic context by analysis of single-zircon crystals. We assume
that zircon may crystallize over a considerable time in the same batch of
magma, even prior to the final emplacement. This may be startling for
people who used to think of the emplacement “age” of a pluton. It has
been made possible by recent improvement in analytical precision and
accuracy in U–Pb zircon dating: development of annealing–leaching
(chemical abrasion) techniques (Mattinson, 2005) to mitigate the
problem of lead loss, improved precision by using a new 205Pb–233U–
235
U tracer solution (Condon and Members of the Earthtime Working
Author's personal copy
U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
Group, 2005) and reducing tracer calibration uncertainties (Schoene
et al., 2006; Schmitz and Schoene, 2007), and thorough monitoring of
analytical conditions during mass spectrometric analysis (described in
Sláma et al., 2008). Part of this effort is a community-wide intercalibra-
209
tion effort for removing inter-laboratory biases (Condon and Members
of the Earthtime Working Group, 2005; Sláma et al., 2008; see most
recent status on www.earth-time.org). Furthermore, we are able to
quantify the initial subsolidus cooling of a crystallized magma body
Fig. 1. Geological map of the Adamello batholith in N Italy. Inset: Map of Re di Castello unit, southern Adamello, with (1) Western Adamello tonalite, (2) Mte. Re di Castello tonalites,
(3) Lago d'Arno/Lago Boazza leucotonalites, (4) Badile granodiorite, (5) Bruffione granodiorite, (6) Listino tonalite, (6a) Listino poprphyry ring structure, (7) Galliner tonalite,
(8) Passo del Termine/Val Paghera leucotonalites, (9) Lage della Vacca tonalites, (10) Blumone gabbros and quartz-diorites, (11) Alta Guardia tonalite, (12) Val Fredda tonalite.
Sample localities are indicated by asterisks.
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210
U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
Table 1
U, Pb isotopic results and ages for zircons from Southern Adamello.
Samplea
Monte
1
2
3
4
5
6
7
8
9
10
11
12
13
Weight
Concentrations
(mg)
U
Th/Ub
Pb
Pb
Rad.
(ppm)
Nonrad.
(pg)
Frerone aplite granite dykes (GMF)
z
0.0070
954
6.75
0.42
z
0.0020
785
7.78
0.47
z
0.0016
861
6.12
0.39
z
0.0049
1094
7.43
0.53
z
0.0049
1061
8.33
0.48
z
0.0028
726
5.23
0.86
z
0.0014
983
6.94
1.43
z
0.0023
1451
10.39
0.20
z
0.0022
1207
8.17
0.54
z
0.0023
941
8.78
0.56
z
0.0009
878
6.20
0.53
z
0.0013
1188
8.42
2.37
t
0.0125
85
0.83
13.50
Atomic ratios
Apparent ages
206/
204c
207/
235d,e
Error
0.62
0.48
0.64
0.43
0.46
0.54
0.60
0.43
0.44
0.43
0.61
0.62
2.08
6926
2114
1564
4444
5425
1065
418
59,148
2073
2204
642
289
52
0.04275
0.06544
0.04270
0.04337
0.05051
0.04468
0.04296
0.05235
0.04294
0.06249
0.04255
0.04240
0.04232
206/
238d
Error
0.23
0.33
0.44
0.25
0.22
0.31
1.06
0.10
0.26
0.31
0.90
1.64
5.21
0.00661
0.00951
0.00660
0.00667
0.00762
0.00685
0.00663
0.00686
0.00663
0.00909
0.00662
0.00663
0.00662
2σ
(%)
Error
206/238
2σ
(%)
Corrected for disequilibrium
0.20
0.20
0.20
0.20
0.19
0.20
0.10
0.06
0.06
0.09
0.10
0.13
0.37
0.04688
0.04988
0.04689
0.04719
0.04804
0.04727
0.04696
0.05539
0.04701
0.04989
0.04659
0.04635
0.04639
0.11
0.25
0.38
0.15
0.11
0.23
0.99
0.08
0.24
0.29
0.85
1.54
4.89
42.50
61.05
42.43
42.82
48.97
44.04
42.63
44.04
42.57
58.30
42.56
42.63
42.51
42.50
64.36
42.45
43.11
50.03
44.38
42.71
51.81
42.69
61.55
42.31
42.16
42.09
43.06
189.43
43.64
58.86
101.20
62.86
47.18
427.98
49.70
189.89
28.30
15.83
18.16
0.87
0.65
0.51
0.80
0.87
0.67
0.73
0.62
0.44
0.36
0.55
0.79
0.88
2σ
(%)
207/235
207/206
Error
corr.
207/
206d,e
Leucotonalite of Alta Guardia (TAG)
14*
z
0.0031
1163
7.78
15*
z
0.0069
354
3.29
16*
z
0.0075
418
3.13
17*
z
0.0064
438
3.07
18*
z
0.0046
634
4.24
19
z
0.0034
694
4.84
20
z
0.0074
356
2.35
21
z
0.0090
308
2.09
22
z
0.0097
511
3.39
23
z
0.0130
292
1.98
24
z
0.0102
301
2.02
25
z
0.0092
513
3.49
26
t
0.0226
329
1.99
27
t
0.0204
453
3.50
28
t
0.0275
709
4.40
29
t
0.0186
897
3.60
30
t
0.0377
616
3.74
2.56
2.13
1.10
2.50
2.60
0.71
0.72
0.63
0.94
0.89
1.50
1.13
53.20
45.44
56.85
43.31
43.49
0.47
0.10
0.87
0.48
0.47
0.64
0.46
0.55
0.46
0.36
0.44
0.55
0.14
1.07
0.22
0.14
0.13
596
697
1238
492
478
1414
1570
1901
2148
1837
859
1708
76
103
160
124
241
0.04186
0.06436
0.04202
0.04342
0.04174
0.04169
0.04176
0.04191
0.04187
0.04447
0.04263
0.04200
0.04237
0.04218
0.04202
0.04192
0.04203
0.83
0.74
0.70
1.11
1.11
0.38
0.40
0.37
0.33
0.28
0.56
0.32
0.89
0.76
1.87
1.05
0.38
0.00649
0.00935
0.00650
0.00668
0.00647
0.00645
0.00644
0.00646
0.00646
0.00680
0.00657
0.00646
0.00643
0.00643
0.00644
0.00642
0.00649
0.35
0.37
0.38
0.35
0.35
0.27
0.21
0.20
0.06
0.07
0.08
0.07
0.07
0.08
0.12
0.09
0.07
0.04681
0.04992
0.04689
0.04715
0.04679
0.04684
0.04699
0.04702
0.04697
0.04746
0.04710
0.04715
0.04781
0.04755
0.04733
0.04732
0.04694
0.71
0.62
0.57
1.01
1.01
0.26
0.33
0.30
0.30
0.26
0.52
0.30
0.84
0.72
1.76
0.99
0.35
41.67
59.99
41.76
42.91
41.58
41.48
41.41
41.54
41.54
43.66
42.18
41.52
41.31
41.34
41.38
41.28
41.73
41.64
63.33
41.79
43.16
41.52
41.47
41.54
41.69
41.65
44.18
42.39
41.78
42.14
41.95
41.80
41.70
41.80
39.56
191.36
43.79
56.71
38.33
41.20
48.67
50.27
47.70
72.47
54.28
56.65
89.80
77.01
65.88
65.60
45.98
0.53
0.55
0.58
0.43
0.43
0.73
0.55
0.59
0.43
0.40
0.55
0.41
0.73
0.54
0.89
0.71
0.50
Tonalite of Malga Listino (TML)
31*
z
0.0015
847
32*
z
0.0040
500
33
z
0.0026
192
34
z
0.0010
325
35
z
0.0012
563
36
z
0.0014
607
37
z
0.0024
420
38
z
0.0040
312
39
z
0.0036
213
40
z
0.0025
232
41
z
0.0028
600
42
z
0.0015
816
43
t
0.0253
332
44
t
0.0184
353
45
t
0.0169
444
46
t
0.0273
396
47
t
0.0298
393
5.68
4.83
1.33
2.21
3.69
4.20
2.67
2.18
1.49
1.70
3.99
5.74
3.51
3.60
4.00
3.90
3.92
2.09
7.79
0.65
0.59
0.55
0.88
0.34
0.79
0.79
0.92
0.87
0.86
28.19
20.73
30.31
29.21
38.58
0.53
0.52
0.63
0.56
0.45
0.66
0.36
0.69
0.68
0.83
0.51
0.72
2.74
2.52
1.83
2.28
2.38
265
123
249
339
528
420
1265
656
406
274
793
589
141
146
119
170
142
0.04126
0.04166
0.04178
0.04170
0.04133
0.04121
0.04152
0.04141
0.04175
0.04242
0.04127
0.04119
0.04118
0.04129
0.04180
0.04156
0.04150
1.59
2.42
1.79
1.66
0.92
0.61
0.53
0.76
1.15
1.91
0.64
0.81
0.84
0.68
0.89
0.76
1.16
0.00639
0.00644
0.00638
0.00644
0.00637
0.00638
0.00637
0.00642
0.00643
0.00647
0.00640
0.00641
0.00637
0.00637
0.00636
0.00638
0.00638
0.35
0.36
0.22
0.87
0.45
0.24
0.21
0.08
0.10
0.16
0.08
0.08
0.09
0.08
0.08
0.07
0.18
0.04680
0.04694
0.05
0.04692
0.04704
0.04683
0.04731
0.04675
0.04706
0.04753
0.04673
0.04658
0.04691
0.04702
0.04770
0.04722
0.04721
1.49
2.27
1.70
1.29
0.75
0.52
0.48
0.71
1.08
1.79
0.60
0.76
0.80
0.64
0.85
0.72
1.09
41.08
41.36
41.02
41.48
40.95
41.01
40.90
41.28
41.34
41.60
41.16
41.21
40.91
40.93
40.85
41.03
40.97
41.05
41.44
41.56
41.48
41.12
41.01
41.30
41.20
41.53
42.18
41.06
40.98
40.98
41.08
41.58
41.35
41.29
39.26
46.00
72.59
45.30
51.30
40.60
64.92
36.43
52.22
75.91
35.40
27.73
44.69
50.07
84.20
60.13
59.90
0.38
0.48
0.62
0.46
0.59
0.54
0.43
0.66
0.72
0.77
0.55
0.68
0.54
0.54
0.53
0.58
0.46
Leucotonalite of Cima di Vallone
48
z
0.0038
653
49
z
0.0011
1143
50
z
0.0013
674
51
z
0.0046
271
52
z
0.0044
267
53
z
0.0042
239
54
z
0.0049
306
55
z
0.0042
321
56
z
0.0040
178
57
z
0.0080
92
(VAL)
5.04
7.59
4.52
1.86
1.83
1.61
2.09
2.33
1.17
0.62
0.88
0.80
0.62
2.36
0.76
0.79
0.64
0.68
1.07
1.39
0.50
0.52
0.55
0.66
0.64
0.56
0.63
0.71
0.48
0.57
1323
645
585
228
636
526
956
832
284
229
0.04916
0.04141
0.04144
0.04123
0.04138
0.04132
0.04116
0.04309
0.04137
0.04177
0.34
0.75
0.99
2.34
0.80
0.90
0.46
0.55
1.52
2.08
0.00739
0.00637
0.00637
0.00637
0.00636
0.00638
0.00636
0.00659
0.00637
0.00638
0.06
0.09
0.12
0.17
0.09
0.09
0.07
0.07
0.13
0.16
0.04822
0.04711
0.04715
0.04691
0.04716
0.04694
0.04691
0.04740
0.04707
0.04745
0.32
0.71
0.94
2.23
0.75
0.84
0.43
0.52
1.43
1.94
47.49
40.96
40.96
40.96
40.90
43.14
40.90
42.37
40.96
41.03
48.73
41.20
41.23
40.96
40.90
41.03
40.90
42.37
41.16
41.56
110.07
54.85
56.80
44.67
57.37
46.15
44.67
69.38
52.79
71.90
0.41
0.55
0.51
0.65
0.62
0.68
0.51
0.49
0.71
0.88
a
b
c
d
e
z = zircon, t = titanite; all zircons annealed-leached, all single grains, titanites are multigrain fractions; * = measured with
Calculated on the basis of radiogenic Pb208/Pb206 ratios, assuming concordancy.
Corrected for fractionation and spike.
Corrected for fractionation, spike, blank and common lead (Stacey and Kramers, 1975).
Corrected for initial Th Disequilibrium, using an estimated Th/U ratio of 4 for the melt.
205
Pb–235U tracer.
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U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
211
down to 600 °C (Cherniak, 1993) by dating titanite with the same
technique.
The results presented here are from a study of rocks from the Re di
Castello (RdC) pluton in the southern Adamello batholith (northern
Italy), and provide a first accurate quantification of the duration of
zircon growth in individual magma pulses of a composite magmatic
unit. The RdC pluton is a complex assembly of up to several km3-sized
bodies of gabbros, diorites, tonalites, granodiorites and minor granites
(Brack, 1985; Ulmer et al., 1985), showing spectacular field evidence
for the sequential emplacement of distinct magmatic units, mingling
and mixing of magma at variable volume fractions of crystals
(crystallinity) as well as syn-magmatic ductile-to-brittle deformation
features (Brack, 1985; Blundy and Sparks, 1992; John and Blundy,
1993). The present data demonstrate that state-of-the-art singlezircon preparation techniques and analytical precision are capable of
resolving the residence time and/or the period of zircon growth in a
given population. Titanite U–Pb ages allow assessing whether the
youngest zircon dates indeed represent the youngest zircon growth in
the cooling magma or if data are biased by lead loss. We also combine
the hafnium isotope composition of zircon with their crystallization
age in order to characterize the source of melts as a function of time.
sediments up to several kilometers wide and exhibiting features of
emplacement-related deformation (Brack, 1981). Limbs of the same
sediments also separate different magmatic bodies. The main
magmatic lithologies are fine-grained tonalites to granodiorites that
are in some cases homogenous and abut sharply against one another.
Elsewhere, however, contacts between individual magmatic pulses
can only be discerned through subtle differences in texture and
mineralogy. Gabbros and diorites occur as coherent and often sheetlike intrusive bodies especially along the margins of the RdC, but are
also present as small enclaves in the acidic rocks (Blundy and Sparks,
1992). Structures of syn-intrusive (forceful) deformation are observed
in several magmatic bodies in the southern RdC (Brack, 1985; John and
Blundy, 1993). The mutual relationships in outcrop suggest that
immediately after their emplacement the different intrusions at least
locally coexisted as largely crystallized and cooling bodies, with
residual interstitial liquid, deforming in a plastic way during the
emplacement of subsequent melt batches (e.g., John et al., 1997).
Radiometric ages available to date (Del Moro et al., 1985; Villa, 1985;
Hansmann, 1986, Hansmann and Oberli, 1991) fully overlap and
indicate a narrow time span of less than 4 Ma for the emplacement
and cooling of the southern RdC intrusive rocks.
2. The Adamello batholith
3. U–Pb and Hf isotope results
The Adamello batholith is exposed over an area of ~670 km2 with
up to 3 km of vertical relief in northern Italy (Fig. 1; see Callegari and
Brack, 2002 for further references). It is the largest of the Tertiary
Periadriatic intrusions of the Alps and the Adamello magmas were
emplaced during the Middle/Late Eocene–Early Oligocene into the
South Alpine Variscan basement and its non-metamorphic PermoTriassic cover rocks. The outer borders of the igneous complex are
largely primary intrusive contacts. Along its north-eastern corner the
batholith is tectonically bound, i.e. by the late- to post-magmatic
Tonale Line and by the Miocene Giudicarie Line, and its interior parts
are crossed by only a few late to post-magmatic fault zones
(Pennacchioni et al., 2006).
The Adamello Batholith is divided into four plutons (Re di Castello,
Adamello, Avio, and Presanella plutons; Callegari, 1985; Callegari and
Brack, 2002) that encompass around 10 Ma of magmatic activity (43
to 33 Ma). The individual plutons are composite bodies and were
emplaced sequentially from the oldest units in the south (~43 Ma) to
the youngest units in the north (Del Moro et al., 1985). Geochemical
(Dupuy et al., 1982; Macera et al., 1985) and isotopic (Cortecci et al.,
1979; Del Moro et al., 1985) results demonstrate a roughly northward
increase of the initial 87Sr/86Sr (Sri) and 18O/16O (δ18O) ratios, along
with the concentrations of incompatible elements such as U, Cs and K,
with maximal values in the Avio pluton. This points to increasing
crustal contamination during the interval of magmatism. Mafic rocks
of the southern Re di Castello unit (cumulate olivine–pyroxene
wherlites to hornblendites to hornblende gabbros and diorites;
Ulmer et al., 1985; Blundy and Sparks, 1992) have Sri and δ18O close
to mantle values (Sri: 0.7036–0.7038, δ18O: +5.9–6.0), overlapping
with those of the associated tonalitic–granodioritic rocks, whereas Sri
ratios in the north range up to 0.711 (Cortecci et al., 1979; Del Moro
et al., 1985). Altogether, the basic and felsic rocks define a typical calcalkaline fractionation trend (e.g., Macera et al., 1985; Ulmer et al.,
1985). On the basis of these results Taylor (1980), Macera et al. (1985),
Bigazzi et al. (1986) and Thompson et al. (2002) proposed and refined
an assimilation fractional crystallization (AFC) model for the derivation and evolution of the Adamello, with the digestion of progressively
larger amounts of lower to middle crustal material into fractionating
mantle-derived magmas.
The southernmost unit of the Adamello batholith, i.e. the Re di
Castello (RdC) pluton, is a complex assembly of gabbros, diorites,
tonalites, granodiorites and minor granites, (Fig. 1, lower panel). These
magmatic units are bordered by rims of contact-metamorphic Triassic
Four lithologies have been sampled for this study, which cover the
entire age range of the southern RdC unit and most of which exhibit
unequivocal relative age relationships with the neighbouring rocks and
each other (Fig. 1; descriptions and coordinates in Table S1). Zircon
separation was carried out at the University of Bologna following the
procedure of Marocchi et al. (2008). Analytical techniques of the U–Pb
age determinations of zircon follow those described in Schaltegger et al.
(2008), those of titanite are outlined in Chiaradia et al. (2009). Part of
the analyses were carried out using the 205Pb–233U–235U tracer solution
(Condon and Members of the Earthtime Working Group, 2005), which
has been internationally intercalibrated and proven to yield 206Pb/238U
inter-laboratory reproducibility to better than 0.1% (see Sláma et al.,
2008), the other part with a 205Pb–235U spike from ETH Zürich (analyses
marked in Table 1 by asterisks). All analyses were carried out on the
Triton thermal ionization mass spectrometer at University of Geneva,
using a discrete-dynode MasCom® secondary electron multiplier in ion
counting mode for the analysis of Pb. Uncertainty ellipses of individual
analyses in Fig. 2 are at 2σ level and do not include the uncertainty of
tracer calibration or non-blank common Pb composition. For the
calculation of average 206Pb/238U titanite ages, an additional estimated
uncertainty for the initial Pb composition has been added to analytical
uncertainty of the 206Pb/238U ratio (2% for 206Pb/204Pb, 1% for 207Pb/
204
Pb, 2% for 208Pb/204Pb).
The Hf fraction was isolated using Eichrom™ Ln-spec resin, and
measured in static mode on a NuPlasma™ multi-collector ICP-MS using an
Aridus nebulizer for sample introduction. 176Lu/177Hf ratios of analyzed
zircons were not determined but 176Hf/177Hf ratios were age corrected for a
typical value of 176Lu/177Hf in zircon of 0.0005 (compare, e.g., to values in
Kemp et al., 2008; or in Miskovic and Schaltegger, 2009). The 176Lu decay
constant of Scherer et al. (2001) was used for calculation; the correction
stayed within limits of analytical precision of the measured 176Hf/177Hf
ratios in all cases. The Hf isotopic ratios were corrected for mass
fractionation using a 179Hf/177Hf value of 0.7325 and normalized to
176
Hf/177Hf =0.282160 of the JMC-475 standard (Blichert-Toft et al. 1997).
The JMC-475 standard was measured every fourth position in similar
measurement conditions as the zircon samples and yielded values
between 176Hf/177Hf =0.282130 to 0.282190 according to the daily instrument tuning. 176Lu and 176Yb present in analysis never amounted to more
than 0.0003% or 0.01%, respectively, of the intensity of the 176Hf beam and
were corrected for using isotopic ratios for 176Lu/175Lu=0.02656 and
176
Yb/172Yb=0.586155. Errors of the measured 176Hf/177Hf ratios are
either given as external 2 σ reproducibility of standard measurements (i.e.
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U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
±0.5 ε units) or individual 2 σ errors, whichever is larger. εHf values and
TDM model ages were calculated with (176Hf/177Hf)CHUR(0) =0.282785
(Bouvier et al., 2008) and use present-day depleted mantle values of 176Hf/
177
Hf =0.28325 (Nowell et al., 1998), 176Lu/177Hf =0.0385 (Griffin et al.,
2000), and a crustal 176Lu/177Hf =0.015 (Griffin et al., 2002).
Prior to the selection of zircon crystals for analysis, a morphological
classification into acicular, long prismatic, prismatic, short-prismatic and
equant morphologies was applied. In addition, the selected zircons were
shortly described in terms of inclusions, milky-translucent vs. transparent
crystal domains, presence of cracks, and/or even characterized by simple
drawings. Despite this effort, no clear dependence between external
morphology and apparent 206Pb/238U age can be recognized, except for
indications of inherited material in the core (circular round inclusions,
central milky domain, or radial cracks). For better characterization of the
dated material, a micro-photographic documentation will have to be
established for such detailed studies in future.
Granitic aplite dykes of the Monte Frerone granite (GMF) are found
as a complex network of apophyses emerging from a several hundred
meter long body crossing the pre-Adamello folds of the carbonates
and hornfelses of the inner contact aureole in the southwestern flank
of Monte Frerone. The central body of GMF is cut by an outlier of the
Valfredda leucotonalite (Brack, 1984) and peripheral aplitic dykes are
locally affected by deformation induced by a subsequent intrusion.
The GMF is hitherto the oldest magmatic lithology that can be clearly
assigned to the southern RdC pluton and accordingly it is the oldest
known magmatic product of the entire Adamello batholith.
Apart from an Early Paleozoic inheritance (analyses 2, 5, 6 and 10
yielding an upper intercept age at 469±45 Ma), analytically concordant
zircon data of GMF yield a scatter of 206Pb/238U dates from 42.43 to
42.63 Ma, i.e. record an age dispersion of 200 ka (Fig. 2a). From 6
clustering analyses a mean 206Pb/238U age of 42.57±0.05 Ma may be
calculated; the MSWD value of 5, however, indicates that the points are
not equivalent. Therefore, the youngest age of 42.43±0.09 Ma is taken as
a first approximation for the age of autocrystic zircon growth and of
emplacement, whereas the other analyses date prolonged crystallization
of autocrystic zircon (crystallized in the same magma at an earlier stage),
or of antecrystic zircon (earlier crystallization in the same magmatic
system), see discussion below. There is no relation between external
morphology and 206Pb/238U age. The sample yielded very few titanite
crystals, one of which could be dated at a 206Pb/238U age of 42.51±
0.16 Ma; it is thus coeval with the youngest zircon within analytical
uncertainties.
The tonalite of Alta Guardia (TAG; nr. 11 in Fig. 1) is a peripheral
intrusion, lithologically similar to but possibly distinct from the Val
Fredda leucotonalite. The latter unit very likely predates the formation
of the Listino ring (see Fig. 1) and its interior.
Apart from Paleozoic inherited components yielding an upper
intercept age of 437 ± 37 Ma (analyses 15, 17, 24), eight TAG zircons
crystallized over 350 ka (between 41.41 and 41.76 Ma; Fig. 2b). A
mean 206Pb/238U date of 41.53 ± 0.04 Ma with an MSWD of 3.1 may be
derived from these analyses; we assume, however, that the youngest
date of 41.41 ± 0.09 Ma (analysis 19) reflects autocrystic zircon
growth and approximates the age of emplacement. Four out of five
titanite analyses yield an average 206Pb/238U date of 41.31 ± 0.03 Ma
(MSWD = 1.03; with enhanced Pb common uncertainty), which is
insignificantly younger than the age for the youngest zircon. A fifth
titanite analysis plots at an older age of 41.73 ± 0.03 Ma. There is no
relation between external morphology and 206Pb/238U zircon age.
The tonalite of Malga Listino (TML; nr. 6 in Fig. 1) forms the core of
the enigmatic Listino ring structure, a zone of highly deformed
tonalite with abundant mafic enclaves and sediment inclusions,
crosscut by syn-magmatic dykes (Brack, 1984, 1985). The TML core,
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however, is composed of an undeformed tonalite which shows some
variation in grain-size (fine to medium grained) and containing
comparably few mafic enclaves but no sedimentary xenoliths.
The results of 14 zircon crystals scatter between 40.90 and
41.60 Ma, indicating apparent prolonged zircon growth over 700 ka
(Fig. 2c). Analyses 29 and 31 yielded 206Pb/238U dates of 40.90 and
40.95 ± 0.09 Ma and are thus considered to represent autocrystic
zircon growth and approximate the age of emplacement of this
tonalite. The lack of enclaves and xenoliths is also reflected by the lack
of inheritance in the analyzed zircon grains. Five titanite analyses
scatter between 41.03 and 40.85 Ma, averaging at 40.94 ± 0.09 Ma
(MSWD = 4.3; with 0.1% enhanced Pb common uncertainty). Excluding the oldest or youngest of the analyses does not yield a significantly
different result (40.97 ± 0.10 and 40.90 ± 0.07 Ma, respectively). The
titanite dates thus agree with the date of the youngest zircons. There is
a tendency that the youngest zircon 206Pb/238U ages are found in
longprismatic to acicular zircon grains (analyses 31, 33, 35–37).
The leucotonalite of Cima Vallone (VAL; nr. 3 in Fig. 1) crosscuts the
Listino ring structure and its tonalitic core (TML) and likely belongs to
the youngest plutonic rocks in the southern part of the RdC pluton.
Two zircons from VAL incorporated a Paleozoic inherited component (analyses 53 and 55; upper intercept age of 405 ± 54 Ma),
whereas 7 concordant grains cluster around a mean 206Pb/238U date of
40.93 ± 0.04 Ma (Fig. 2d). The elevated MSWD of 4.4, however,
indicates that the ages are not equivalent and in turn record zircon
growth over a timespan of 130 ka (41.03–40.90 Ma). The youngest
zircons at 40.90 ± 0.03 Ma (analyses 41 and 43) record growth of
autocrystic zircon and thus approximate the age of emplacement.
There is no relation between external morphology and 206Pb/238U age.
Initial Hf isotopic compositions of magmatic zircons from TML and
VAL range between εHf values of +6.4 and +8.9, characterizing a
rather juvenile source for these two intrusions. The zircons of GMF
and TAG, however, have initial Hf isotopic composition equal to εHf
values between − 2.8 and +3.0, indicating a more hybrid character of
those melts (Table 2). These values provide evidence for the existence
of different magma sources within the same magmatic system. The Hf
Table 2
Hf isotopic compositions of dated zircons.
176Hf/
177Hf
±2 σ
176Hf/177Hf
eps Hf
(T)
(T)
±2 σ
T2 (DM)
(Ga)
GMF Granite Monte Frerone
1
0.282825
0.000004
3
0.282782
0.000021
4
0.282829
0.000008
0.282824
0.282782
0.282829
2.8
1.3
3.0
0.5
0.5
0.5
0.87
0.95
0.86
TAG Leucotonalite Alta Guardia
14
0.282832
0.000008
15
0.282722
0.000005
16
0.282744
0.000006
23
0.282803
0.000004
24
0.282667
0.000004
25
0.282808
0.000003
0.282832
0.282722
0.282744
0.282803
0.282667
0.282808
3.0
− 0.8
− 0.1
2.0
− 2.8
2.2
0.5
0.5
0.5
0.5
0.5
0.5
0.85
1.08
1.03
0.91
1.19
0.90
TML Tonalite Malga Listino
31
0.282993
0.000013
32
0.282980
0.000013
38
0.282926
0.000006
41
0.282948
0.000023
0.282993
0.282980
0.282926
0.282948
8.7
8.3
6.4
7.1
0.5
0.5
0.5
0.5
0.52
0.54
0.66
0.61
VAL Leucotonalite Cima di Vallone
48
0.282999
0.000006
49
0.282993
0.000006
50
0.282970
0.000012
0.282999
0.282993
0.282970
8.9
8.7
7.9
0.5
0.5
0.5
0.50
0.52
0.56
Fig. 2. Concordia diagrams and ranked 206Pb/238U age plots containing the results of zircon U–Pb dating of four samples from the southern Re di Castello unit, Southern Adamello.
(a) GMF, granite of Monte Frerone, (b) TAG, tonalite of Alta Guardia, (c) TML, tonalite of Malga Listino, (d) VAL, leucotonalite of Cima di Vallone. Gray ellipses and black bars denote
titanite analyses, open ellipses and bars zircon analyses.
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Fig. 3. Epsilon Hf versus 206Pb/238U age diagram for selected zircons.
isotopic values could point to an evolution towards more juvenile
compositions with time (Fig. 3).
4. Discussion of the isotopic data and their implications
4.1. Age range of magmatism
The new age data cover the majority of the lifetime of the multicomponent southern RdC pluton, which was assembled during some
1.5 Ma. The leucocratic dykes of GMF were emplaced at 42.43 ±
0.09 Ma and likely predate the Lago della Vacca intrusive body (Fig. 1;
Brack, 1984; John and Blundy, 1993). Our emplacement age of 41.41 ±
0.09 Ma for the peripheral TAG intrusion is younger than the U–Pb
thorite age of 42.9 ± 0.2 Ma (Hansmann, 1986) for the Vacca tonalite.
The intrusions of Vacca and Galliner tonalites, and the Blumone
gabbros and diorites are clearly cut by the Listino ring structure and its
core, the TML tonalite at 40.90 ± 0.09 Ma. The VAL leucotonalite is
coeval with the latter at 40.90 ± 0.03 Ma and crosscuts deformed
tonalites of the Listino ring; TML and VAL ages are thus in line with
field relationships.
4.2. How to treat precise U–Pb age data?
The data reveal that age dispersions of up to several 100 ka exist
among a zircon population from the same rock, which may indicate
prolonged growth of autocrystic zircon in the same magma batch, or
incorporation of older, antecrystic zircon from different magma
batches of the same magma system, and may additionally be biased
by incorporation of small inherited cores in the magmatic zircon. Such
a result may be achieved by higher precision through the new and
improved analytical techniques described above, and foster some
fundamental questions of how to interpret high-precision zircon U–Pb
ages from magmatic rocks. Zircon forms in complex magmatic
systems at different levels of the crust. Small melt portions in the
source area within the lower crust may already saturate and crystallize
zircon due to feldspar crystallization lowering the K and Na activity in
the melt (Watson and Harrison, 1983), whereas other melt batches do
so only during the ascent through lower and middle crust, during
further crystallization and cooling. Such zircon has been termed
“antecrystic” (see Miller et al., 2007 and further references therein); it
gets entrained and incorporated into the magma emplaced at a
shallow crustal level and will record the integrated history of melt
extraction, ascent, possibly intermediate crystallization at N50%
crystallinity and subsequent re-mobilization of such “proto-plutons”
before final emplacement. Since magma is emplaced with less than
50% crystals (convection threshold; Marsh, 1981; Lejeune and Richet,
1995), there will be further zircon formation in the stagnant
interstitial melt, leading to mantling of preexisting grains as well as
possible nucleation of new, smaller grains (Charlier and Zellmer,
2000). The last autocrystic zircons crystallizing at the final level of
emplacement would date or at least approximate the emplacement, or
possibly even be slightly younger. The age dispersion that we will
determine by using a representative suite of zircons therefore depends
on what moment the system reaches saturation for zircon. Since we
have analyzed entire grains, integrating growth over a certain period
of time, we will only get a maximum age for the youngest autocrystic
zircon. Furthermore, we have to be aware that we may still record
subordinate effects of lead loss, despite our efforts of chemical
abrasion pre-treatment. The U–Pb results of magmatic titanite from
TAG and TML lithologies (Fig. 2b, c) demonstrate that titanite ages
may serve as a test for this hypothesis: titanite has been shown to have
a closure temperature of ca. 650 ± 50 °C for grains with an effective
diffusion radius between 100 and 1000 µm and 2–10 °C/Ma cooling
rate (Cherniak, 1993); magmatic titanite therefore records its crystallization close to the solidus or initial cooling below the solidus. In both
samples TAG and TML the titanites are within analytical uncertainty of
the youngest zircon (Fig. 2b, c), arguing that our annealing–leaching
pre-treatment is indeed effective at minimizing the post-crystallization lead loss. One older titanite analysis of sample TAG (Fig. 2b)
asks, however, for caution and indicates that titanite may grow earlier
and be possibly mantled by refractory phases in order to keep its
isotopic system closed over 105 a.
Our results of zircon U–Pb dating in fact roughly represent absolute
durations of zircon growth. Since we are dealing with single grain
analyses recording an integrated signal across all growth zones, we
may consider it as a minimum duration of growth. On the other hand,
since we cannot rule out that an antecrystic or even xenocrystic
component is affecting the apparent age range, we may also consider
obtaining periods of time that are longer than the durations of zircon
crystallization. In the case of samples GMF, TAG and VAL (Fig. 2a, b, d,
respectively) inherited Pb is evidenced by significantly higher 206Pb/
238
U ages (analyses 2, 5, 6, 8, 10, 15, 17, 24, 48, 53 and 55). Sample TML,
however, does not show any sign of a significantly older inherited
component that would contribute to the apparent c. 700 ka-long
zircon growth period. Generally speaking, most intermediate and acid
rocks will reach zircon saturation early in magma evolution, zircon
will record a prolonged period of growth (possibly biased by ante- and
xenocrystic components) and the emplacement age may be best
approximated by the youngest dates only. With a significant number
of analyses we may use the value of mean square of weighted deviates
(MSWD) of the mean 206Pb/238U age to assess whether the observed
scatter is beyond purely analytical variation. It is self-understanding
that ID-TIMS techniques are too difficult and time-consuming to ever
achieve a number of analyses from one sample that is representative
for the entire population. If a dataset of seven concordant analyses –
e.g., in case of the sample VAL (Fig. 2d) – already shows nonequivalency of points and a MSWD of 4.4, we can be sure that there is
geological complexity beyond analytical scatter. Such information for
rocks of similar age cannot be obtained from low-precision U–Pb
analyses such as by LA-ICP-MS or SIMS, because their large
uncertainty on individual data points would suggest equivalency of
points (with MSWD's around unity) and hence yield a potentially
inaccurate weighted mean age for the emplacement, even outside of
ICP-MS or SIMS analytical uncertainty.
4.3. Implications for zircon crystallization and survival of
xenocrystic zircon
Through crystallization, initially zircon-undersaturated mafic magma
batches become volatile-saturated and eventually rise to intermediate
crustal levels due to the lowered viscosity and higher buoyancy. They may
only saturate zircon due to cooling and oversaturation of the interstitial
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U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
fluid-saturated residual melt. The first precipitating zircons date in such a
case the moment of zircon saturation. Only these kinds of samples are
anticipated to yield perfect clusters of zircon 206Pb/238U ages within
present analytical errors and confirm that analytical reproducibility is
indeed within the claimed analytical uncertainty (see examples in
Schoene et al., 2006).
Such a case has not been found among the four studied Adamello
rocks, which all show age variation largely in excess of analytical
scatter (130 to 700 ka) with an analytical uncertainty of ±30–90 ka in
the 206Pb/238U age of an individual analysis. The studied tonalites and
granodiorites have saturated zircon already at an early stage of their
geochemical evolution, because Zr acts as an incompatible element
during crystallization of pyroxene, amphibole and plagioclase. Maximum bulk rock Zr concentrations are reached at around 64% SiO2
(Ulmer, 1986). We can envisage that some melt portions already
started to crystallize zircon at lower crustal levels, despite the
continuous influx of mafic magma batches. In these rocks we
anticipate finding an extended history of zircon crystallization over
200–300 ka, as has been shown for various volcanic cases, e.g., by
Brown and Fletcher (1999), Charlier and Zellmer (2000), Bacon and
Lowenstern (2005), Bachmann et al. (2007), Simon et al. (2008), and
many others. For longer durations of autocrystic zircon growth (as
displayed by sample TML, if we assume that the age range is not due to
ante- and xenocrystic material; Fig. 2c), it would only be possible if the
necessary heat is sustained by repeated influx of hot mafic magma
into the source region, keeping temperatures above the solidus and
adding more and more juvenile material (Annen et al. 2006). We
would find even more likely that the antecrystic zircons may be
derived from already partly solidified magma at different crustal levels
of the magmatic system (“proto-plutons”) that became (possibly
repeatedly) remelted and with zircon crystals that were entrained
into the leaving melt fraction. Such a hypothesis may be supported by
field observations (antecrystic amphibole and plagioclase crystals
entrained in later liquids), but is, however, not supported by our U–Pb
and chemical data.
Furthermore, we would assume that residual liquid (residual from
basalt fractionation) would very likely mix with partial crustal melts
and carry their zircon xenocrysts in suspension; in samples GMF, TAG
and VAL xenocrystic zircons were indeed detected. In sample TML,
however, they were not. The nature of the digested crustal component
in the RdC pluton did not change during the 1.5 Ma lifetime of the
studied magmatic system, as shown by the similar age for the
inherited component in GMF, TAG and VAL (upper intercept ages of
469 ± 45, 437 ± 37 and 405 ± 54 Ma, respectively). This age range is
typical for ubiquitous Ordovician metamorphism and magmatism in
the Southalpine basement (Boriani et al. 1995; Zurbriggen et al. 1997)
and also coincides with an Rb–Sr isochron age of 460 Ma for the Edolo
schists, used to calculate AFC models of Adamello intrusive rocks
(Bigazzi et al. 1986). Our Hf isotope analyses, however, distinguish
between a more hybrid crustal component for GMF and TAG (with
intermediate, partly negative εHf), and a rather juvenile source for
TML and VAL (with εHf around +8).
215
age between GMF and TML (between 42.4 and 40.9 Ma; Fig. 1b): They
were repeatedly injected by gabbroic to dioritic dykes and sills,
namely the Cadino and Mattoni gabbros intruding into the Val Fredda
tonalites (Blundy and Sparks, 1992) or the stocks and dykes of the
Blumone gabbros (Ulmer et al. 1985). As a result of this massive and
geochemically more juvenile mafic input, the following batches of
evolved magma (samples TML and VAL) have significantly more
juvenile Hf isotopic compositions (εHf values up to +8.9, see Fig. 3).
This is, however, not true for sample TAG, an intrusion more
peripheral to the magmatic center of the pluton, which is displaying
less radiogenic Hf isotopic values and therefore obviously remained
unaffected by the juvenile input. According to the “hot-zone” model of
Annen et al. (2006, 2008) we may argue that the onset of magma
production in the RdC unit was dominated by partial crustal melts
(sample GMF) which may be regarded as low-degree partial melts in
the lower crust produced by hot and dry mantle melts; this peak in
crustal melting would have been followed by pulses of wetter mantle
melts and resulting hybrid liquids became more and more juvenile.
This source evolution is only valid for the southern Re di Castello unit,
in that the more northern units of the Adamello batholith show
increasing crustal influence (or of more radiogenic/older crust) as
traced by Sr and O isotopes with ongoing magmatism between 38 and
33 Ma (Cortecci et al., 1979; Del Moro et al., 1985).
4.5. Zircon as a tracer for magma fractionation processes
We have shown that zircon crystals form during a certain period of
time during magma assembly, ascent and final emplacement at
intermediate to shallow crustal level. By using the variable initial Hf
isotope composition of individual zircons we may be able to trace the
magma mingling of hybrid melt batches and/or exchange of
antecrystic zircon between melt batches in certain cases. In complement to Hf isotopes that serve as a source indicator, we may
additionally use the model Th/U ratio (Th concentration calculated
on the basis of 208Pb rad. assuming concordancy between 206Pb/238U
and 208Pb/232Th) of single-zircon crystals to trace magma fractionation. Fig. 4 shows Th/U ratios of zircon and titanite of all 4 samples
dated in this study and leads to the following conclusions: GMF and
VAL show restricted variations of the Th/U from 0.65 to 0.45 that are
independent of measured age, probably indicating short-lived crystallization without significant fractionation due to precipitation of
concurrent U and Th-bearing phases, respectively. TAG and TML
show much larger scatter (between 0.85 and 0.35), which in the case
of sample TML is age dependent, pointing to decreasing Th/U during
4.4. Heat and magma mingling
Field relationships between some of the magma pulses suggest
that the intrusion of each respective pulse at the present-day
observable structural level occurred while the previous batches had
already cooled and mostly solidified. This interpretation is in line with
our U–Pb titanite ages indicating rapid cooling down to 600 °C within
the analytical uncertainty of the U–Pb age of the youngest autocrystic
zircon (c. ±0.09 Ma; Fig. 2b, c).
We therefore consider the intrusions of GMF, TAG and TML + VAL,
respectively, as three separate magma batches, which did not mingle
with the previous magmas. Mingling is recognized in lithologies of the
Blumone and the Val Fredda Complexes, which are intermediate in
Fig. 4. Th/U versus 206Pb/238U age diagram for dated magmatic zircons and titanites; the
latter are labelled “ti”.
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zircon crystallization. Uranium is known to be four times more
compatible in the zircon lattice than is Th (Mahood and Hildreth
1983). A reduction of the Th/U ratio in the crystallizing zircon and in
the magma therefore asks for contemporaneous precipitation of a Thbearing phase. The TML titanites incorporated Th/U at ratios up to 2.7
(Fig. 4 and Table 1), suggesting that concurrent titanite crystallization
may have led to melt depletion in Th. We therefore may interpret the
largest spread and the lowest values for Th/U in zircon in sample TML
by the highest Th concentrations (790–910 ppm) and highest Th/U
ratios of co-precipitating titanite in this lithology. The situation is
somewhat reversed in the case of sample TAG: zircon Th/U are
elevated at 0.0.36–0.87, whereas titanite has, with one exception, very
low Th/U of 0.22–0.13, concurrent with a trend towards higher Th/U in
co-precipitating zircon.
5. Description of CL textures of dated samples
Cathodoluminescence imaging (or backscattered electron imaging) is often used to visualize in an empiric way the result of growth
processes in zircons (e.g., Corfu et al. 2003). Do the cathodoluminescence images of zircons analyzed in this study really reflect the
observed differences in the crystallization/ fractionation paths? With
our U–Pb zircon dataset from the southern Re di Castello unit we try to
reconstruct the history of zircon growth and pluton assembly in great
detail and may test the validity of the empirically interpreted CL
information (Fig. 5).
Sample GMF shows well-defined and undisturbed oscillatory zoning
over most of the growth zones (Fig. 5a, b), possibly reflecting the observed
(minimum) 200 ka of uninterrupted crystallization. Some grains show
clear inherited cores (C in Fig. 5a), in accordance to the isotopic results
(analyses 2, 5, 10 with inheritance of old lead; see Table 1), as well as nonplanar, chaotic textures at the interface between core and rim (Fig. 5a). The
latter may be related to post-crystallization processes mobilizing radiogenic Pb; significant Pb loss is, however, excluded by the titanite analysis.
Sample TAG shows the same uninterrupted oscillatory textures
integrating 350 ka of growth (Fig. 5c,d), ending with a high–U zone in
case of Fig. 5c, and very likely featuring small inherited cores in both
Fig. 5c, d, which is in line with the detected inheritance (analyses 15,
17, and 23).
A clearly different picture is presented by zircons of sample TML.
Many zircon grains of this sample contain a sector-zoned central part,
which was formed at high temperatures in deep crustal levels. Sector
zoning is typical for high-T melts, such as in granulites (e.g., Vavra et al.
1999, Schaltegger et al. 1999), gabbros (Peressini et al. 2007), or, more
generally speaking, in high-temperature mantle-derived residual liquids
(examples of zircon in oceanic arc melts in Heuberger et al. 2007). Sector
zoning is produced by an equilibrium process partitioning heavy ions
between prism and pyramid faces (see summary of models in Corfu et al.
Fig. 5. Cathodoluminescence images of representative zircons from sample GMF (a, b), TAG (c, d), TML (e, f), and VAL (g, h). C = core; F = post-crystallization surface-bound
replacement front; P = abrupt change of growth rate for a U-rich (low-CL) prism face; R = resorption interface. Scale bars = 50 µm.
Author's personal copy
U. Schaltegger et al. / Earth and Planetary Science Letters 286 (2009) 208–218
2003). The prism faces in Fig. 5e show enrichment in U and acceleration
of growth (P in Fig. 5e) until the moment when sector zoning gets
replaced by oscillatory (rapid) growth, and grain morphology changing
from a more equant shape with {211} pyramids to a prismatic shape with
{101} pyramids. These observations agree with our observation that
zircon grew over prolonged periods of time; crystallization started a
deeper level, the magma of TML thus saturated deeper in the crust than
the other samples. However, we do not have a direct age control on such
sector-zoned crystals. The zircon imaged in Fig. 5f shows oscillatory
zoning from center to rim, but disturbed by phases of resorption (R) and
surface-bound replacement fronts (F). Both crystals end their growth in
high-U rims.
The crystals of VAL (Fig. 5g, h) show very fine-banded oscillatory
growth, indistinguishable from TAG. The zircon in Fig. 5 h has a small
inherited core, whereas the zircon on Fig. 5g is representative for
abundant secondary replacement structures (F) in zircons of this
sample. The observed age dispersion may therefore be explained by
either prolonged growth of autocrystic zircon or (and?) disturbance of
the lattice during or shortly after crystallization and may be at the
origin of age variation at the 10 ka level.
The conclusion from the CL imaging is that we may find useful
information for the crystallization sequence of the dated zircon
population and these are in line with the results of the age determinations. A direct application will, however, only be possible once we have
direct CL information on the very dated zircon grains.
6. Conclusions
Mid-to-upper crustal intrusions are assembled incrementally by
individual magma pulses rather than being single batches of magma
intruding at one time (see e.g., Michel et al 2008). Our data
demonstrate that the composite southern RdC pluton of the Adamello
batholith was emplaced over 1.5 Ma, with each magma batch of 10− 1
to 100 km3 volume recording single-crystal zircon ages spanning from
130 to 700 ka. This age span may be interpreted in terms of i)
prolonged growth of autocrystic zircon, ii) incorporation of antecrystic
zircon from the same magmatic system, iii) incorporation of minute
amounts of inheritance, or – most likely – a combination of all.
It is very likely that each magma pulse (yielding a coherent
lithological unit) consists of a series of even smaller increments only
visible through careful observation of magmatic fabrics. The availability
of precise U–Pb dates from magmatic zircon with permil analytical
uncertainties in 206Pb/238U age forces us to better understand the
systematics of the zircon crystals that are present in a magmatic rock, in
order to define the date of the youngest autocryst as a best
approximation for the emplacement age of the respective magma batch.
Each of the dated intrusions from the southernmost part of the
Adamello Batholith (from oldest to youngest: Granite of Monte Frerone –
GMF; Tonalite of Alta Guardia – TAG; Tonalite of Malga Listino – TML;
Granodiorite of Cima Vallone – VAL) was emplaced into already fully
crystallized rocks, as indicated by the titanite ages of TAG and TML, which
are coeval or only slightly younger than the youngest zircon age and date
cooling to 650±50 °C, i.e. below the solidus. Substantial mingling
phenomena are known from the southernmost RdC pluton, from
intrusions intermediate in age between GMF and TML (Valfredda and
Vacca to Blumone suites; Blundy and Sparks 1992; John and Blundy 1993).
Hf isotopes of individual, dated zircons act as a source indicator and,
between the intrusion of GMF and TML, document a drastic change in
source and/or in the ratio of partial crustal versus residual mantle melt in
the Re di Castello unit. The geochemical evolution of the melts can be
described in terms of the hot-zone model of Annen et al. (2006, 2008).
The first basaltic magmas intruding into the lower crust lead to a pulse of
crustal partial melts that mix with mafic residual melts to produce the
strongly hybrid GMF, and later TAG magmas. As indicated by juvenile Hf
isotopic values, the mantle component increases during the subsequent
evolution of the system, possibly due to the influx of large volumes of
217
mafic melts into the magmatic system, which triggers the intrusion of
the Blumone gabbro intermediate in age between GMF and TML.
The increased precision in U–Pb ID-TIMS dating allows observation
and documentation of magmatic processes at the 104–105 yr level and
starts to form an observational link between geological processes
acting in the past over 106–108 Ma, and timescales of melt assembly in
large modern volcanic centers.
Acknowledgements
Support of the Geneva isotope laboratory through several projects
of the Swiss National Research Fund is acknowledged. M. Senn and M.
Chiaradia contributed with technical help and advice to the success of
this research. Comments of C. Annen (Geneva) on an early version
were very helpful. The two journal reviews of M. Schmitz and F. Corfu
were extremely thoughtful and are highly appreciated. This study was
initially supported by a APAT (CARG, RL-PAT) for the 079 Bagolino
Sheet funded to G.B and M.M.
Appendix A. Supplementary data
Supplementary data associated with this article can be found, in
the online version, at doi:10.1016/j.epsl.2009.06.028.
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