Development of complex, sub-vertical layering in the Cortaderas gabbro intrusion, Central Chile

Lithos 340–341 (2019) 124–138

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Development of complex, sub-vertical layering in the Cortaderas gabbro intrusion, Central Chile Michael D. Higgins a,⁎, Diego Morata b a b

Sciences de la Terre, Université du Québec à Chicoutimi, Chicoutimi G7H 2B1, Canada Departamento de Geología, Centro de Excelencia en Geotermia de los Andes (CEGA), Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Santiago, Chile

a r t i c l e

i n f o

Article history: Received 1 February 2019 Accepted 8 May 2019 Available online 13 May 2019 Keywords: Igneous layering Gabbro Non-dynamic layering Chile

a b s t r a c t The Cortaderas intrusion is a small gabbroic unit spatially associated with the Lower Cretaceous felsic Illapel Plutonic Suite of central Chile. It is unmetamorphosed, largely unaltered and not significantly tilted. Most of the intrusion has well-developed, vertical to steep layering dominantly defined by variations in abundance in plagioclase and amphibole. In some places the layers are regular and rhythmic, with a spacing of about 0.5–2 cm, but elsewhere layering is more complex with solitary plagioclase-rich bands ~1 cm wide. In some areas layers are convoluted, sometimes with channel-like structures, apparent cross-cutting relationships, and balloon-like structures. Pegmatoids occur as patches and layers, in some places with miarolitic cavities. Plagioclase and amphibole dominate the mineralogy, with minor clinopyroxene as relicts in amphibole, and minor olivine. Chemical and isotopic data from different samples cannot be linked by any single, simple process and several different magma sources are necessary. We propose that the intrusion is a fossil conduit originally linking magma sources possibly associated with the base of the Illapel Plutonic Suite to basaltic-andesite to andesitic volcanism. The first magmas flowed along a major fault, heated the host rocks and solidified as unlayered gabbros. Later magmas flowed up the same structure but cooled slowly in the preheated host rocks, producing the layered parts of the intrusion. Layering was produced by non-dynamic processes, perhaps similar to that which produces Liesegang banding. Plagioclase and pyroxene crystallised first to produce a loose framework close to the walls of the conduit. Latent heat from growth of amphibole may have facilitated equilibration by development of layering and coarsening. The channel-like and cross-cutting structures may have resulted from interactions of chemical diffusion waves with discontinuities in permeability of the crystal framework. During the development of the crystal framework, the interstitial silicate liquid became enriched in water, partly dissolving earlier structures and crystallising as pegmatoids, some with miarolitic cavities. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Layered plutonic rocks have always been of interest for petrologists, as shown by three volumes on the subject (Cawthorn, 1996; Charlier et al., 2015; Wager and Brown, 1968). Layering is commonly defined as the repetitive occurrence of layers or layer sequences with similar phase proportions, textures, crystal sizes or combinations of these features on a scale of millimetres to metres. Most early explanations of shallow-dipping igneous layering invoked dynamic processes that were similar to those that produce layering in sedimentary rocks (e.g. Wager and Deer, 1962). While this may be correct in some situations, it is always important to remember that many magmas solidify in a regime with a low Reynold's number, the ratio between the inertial and viscous forces (Glazner, 2014). ⁎ Corresponding author. E-mail address: [email protected] (M.D. Higgins).

https://doi.org/10.1016/j.lithos.2019.05.008 0024-4937/© 2019 Elsevier B.V. All rights reserved.

At values less than about 2000 sedimentary-type processes cannot be important and flow is lamellar. In some layered rocks, the presence of features such as strongly rhythmic layering and steeply-dipping layering, have been interpreted as evidence of the action of non-dynamic processes. Explanations for such non-dynamic layering generally involve processes that dissipate energy (Boudreau, 2011; L'Heureux, 2013; Namur et al., 2015), which are sometimes grouped under the term ‘geochemical self-organisation’ (Ortoleva, 1994). Most examples are from larger intrusions, and Boudreau (2011) comments that ‘well-developed modal layering is relatively rare in thin or small intrusions’. Hence, the interest in the Cortaderas intrusion, the subject of this paper, as it is a small intrusion in the Coastal Cordillera of Central Chile, and its layering is sub-vertical and spatially complex. The Cortaderas intrusion has been briefly described in Namur et al. (2015) under the name ‘La Cordadera Intrusion’ (a misnomer), but here we give more details and explore more completely the origin of the intrusion. An important goal of this work was to publish an

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example of complex layering in a gabbroic intrusion. Our earlier work revealed that at least one other small intrusion has similar structures and we hope that other workers will be able to identify other examples. The Cortaderas intrusion is considered to be part of the Lower Cretaceous Illapel Plutonic suite (IPS; Fig. 1), a N-S elongated body that intrudes Upper Jurassic to Lower Cretaceous volcanic and volcaniclastic rocks in the Coastal Cordillera of central Chile (Ferrando et al., 2014; Morata et al., 2010; Parada et al., 1999; Rivano et al., 1985). This plutonic suite ranges from mediumgrained gabbro to trondhjemites, with hornblende and biotite ± clinopyroxene bearing tonalites and granodiorites as the most abundant lithologies (Fig. 1). There are also small mafic intrusions, such as the Cortaderas intrusion, as well as sparse centimetre to decimetre-sized mafic enclaves and small mafic dykes. There are four magmatic pulses from 117 Ma to 90 Ma (Ferrando et al., 2014). The Cortaderas intrusion is thought to be contemporaneous

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with the adjacent Mafic Unit of the IPS, which has an age of 115 Ma (Morata et al., 2010). Modelling suggests that the felsic rocks of the IPS could be generated by progressive, low-pressure fractional crystallisation of mafic melts similar in composition to the mafic components of the IPS without or with only small contributions from the continental crust (Parada et al., 1999). The IPS is thought to have been emplaced at a shallow depth equal to pressures of 1.6–1.7 kbar (Varas et al., 2012). An important tectonic feature of this region is the presence of major north-south normal and reverse faults. These can extend for over 100 km and define some of the margins of the IPS (Fig. 1). These faults may predate the emplacement of the IPS and have guided the magmas towards the surface. They may have been reactivated later, cutting the IPS and other units. The Cortaderas intrusion lies close to the extension of one of these major lineaments, the Silla de Gobernador reverse fault (Arancibia, 2004), which may have guided its emplacement.

Fig. 1. Simplified geological map of the Lower Cretaceous Illapel plutonic suite, with the four main units indicated in tones of mauve and other plutonic rocks in pink (Ferrando et al., 2014). The Cortaderas intrusion lies close to the northern margin of the plutonic suite.

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ILL09149

V V

V V

Ro ad

Cortaderas Intrusion 60

V

ILL09150

V

ILL09146

ILL09147

V VV V V

70

V V VV V V V V

ILL09148 70

ILL08063 V V V VV VV V

31.44º S

Mina Proyect Cortaderas

Layered gabbro

V

50 m

V V

71.205º W Unlayered gabbro outcrop/inferred

North

V V V V V

Lower Cretaceous andesites

Fig. 2. Geological sketch map of the Cortaderas intrusion and its host andesites. Exposure is almost perfect along the road, but sparse elsewhere. Areas of dispersed outcrops are shown schematically. Where possible, the nature of the float was used to refine the shape of the intrusion. Field measurements indicate that layering is mostly vertical to sub-vertical.

2. Cortaderas intrusion 2.1. Field geology The Cortaderas intrusion is an ovoid body measuring ~400 by ~400 m outcropping near the northern edge of the IPS. It is named for

a small copper mine (‘Mina Proyect Cortaderas’; Cortadera = Pampasgrass) near the southern limit of the intrusion. A section through the western part of the intrusion is very well exposed along a N-S road cut, but elsewhere outcrops are very limited, and all are shown in Fig. 2. Between the outcrops, the identity of the bedrock can commonly be determined from the abundant float, which has generally not been moved far from its original site. Outcrops are generally quite fresh, but many have a veneer of dark, weathered material a few mm thick. In some places, the veneer has flaked off, leaving patches of fresher, paler material underneath. Because of such variation in surface quality, the trajectory of linear structures has been traced on the photos of natural outcrops. The Cortaderas intrusion is hosted by Lower Cretaceous andesites that lie about 1 km north of the margin of the IPS. The intrusion is undeformed and does not appear to be cut by post-intrusion major faults. Only one felsic dyke, metric in size and of granodiorite composition, was observed (Fig. 3a). Magnetic data show that the IPS and hence the Cortaderas intrusion have not been tilted significantly (Ferrando et al., 2014). As will be shown later, the configuration of the layering in parts of the intrusion means that it is not possible to tilt the intrusion and make all the layers horizontal. Most of the intrusion has well-developed layering, except an areas close to the northern margin which is dominated by massive, dark, fine to medium-grained gabbro. The layering is commonly defined by variations in the proportions and sizes of plagioclase and hornblende. Although lamination or lineation of mineral phases was not observed in the field, weak lamination parallel to phase layering is visible in some thin sections. Field measurements of layering show it to be generally vertical to subvertical with a mean strike to the northeast (Fig. 2). In some places, layers are almost perfectly planar, with regular repeated cycles of plagioclase-rich rock alternating with amphibole-rich rock (Fig. 3). The best examples of this facies are in the road cut at the discovery outcrop (Fig. 2) and on the top of a small hill above the road, allowing limited three-dimensional observations which permit a clearer understanding of the different field relationships. Over a three-

A) Regular, parallel layers

Felsic dyke

50 cm

C) Pegmatoid cutting layering

B) Parallel layers with terminations

1 cm

2 cm

Fig. 3. Outcrop-scale structures in the Cortaderas gabbro intrusion seen in vertical sections. The mineralogy is dominated by plagioclase (pale phase) and amphibole (dark phase). a) Regular, vertical layering is well displayed in the discovery outcrop along the road cut. Layers are parallel and repeat 5 to 20 mm intervals. b) Some layers are not laterally continuous, but branch and terminate. c) Some layers are disrupted by late pegmatoid patches.

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metre long block, the layers are parallel and rhythmic, with a vertical E-W orientation (Fig. 3a). In the southern part of the block the layer repeats are about 5 mm, increasing steadily to the north to about 20 mm. Layers are continuous over tens of centimetres, but have some branches and terminations (Fig. 3b). Another well-layered outcrop occurs the other side of the road about 8 m away, where the layering in this outcrop is also vertical, but N-S. There is no evidence for a fault under the road between the two outcrops (Fig. 2). This is significant because the ensemble of outcrops cannot be tilted in any way to make the layering horizontal. Therefore, even though we have no evidence for tilting, if there has been minor tilting then some components of the layering are still steep. Layering is more irregular elsewhere in the intrusion. In some areas plagioclase-rich layers are 10 to 20 cm apart, but still have the same thickness of about 5-10 mm. Elsewhere, there are solitary, thin (5–10 mm) layers of almost pure plagioclase set in poorly layered or unlayered amphibole/plagioclase rock (Fig. 4b). Ten-cm scale ‘wavy’ layering is also present and is truncated by irregular layers in some places (Fig. 4b). Some layers terminate abruptly against other planar layers without evidence of mechanical displacement along such contacts. The geometry of these surfaces does not indicate a unique ‘polarity’, as would be expected for dynamically produced cross-cutting relationships, such as erosion by vertical flow. In some areas there are strongly curved layers resembling channels which appear to cut into older layered rocks. However, it should be remembered that the whole ensemble is sub-vertical and hence these cannot be channels developed by dynamic processes (Fig. 4a). One such pseudo-channel (not shown) has an angle of over 150° between the two extremities. One outcrop has unusual balloon-like structures: Narrow (10– 20 cm) zones of gabbro cut layering and project into massive gabbro (Fig. 5). Layering occurs in some of the balloons parallel to their edges. a) Channel-like structures

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Balloon-like structures

20 cm

Fig. 5. Intersecting ‘balloon-like’ structures in a horizontal outcrop of layered gabbro.

Several generations of such structures seem to cut each other. Earlier structures do not seem to be displaced. Very rarely layers are offset by synmagmatic faults, however, there are some cross-cutting bands rich in amphibole which may be former fractures healed at magmatic or high sub-solidus temperatures. There no relicts of interstitial liquids. At some locations normal layering defined by phase abundances can be accompanied by abrupt changes in grain-size parallel to layering. There is no evidence of movement along any of these structures in the form of C\\S fabrics. Gabbroic pegmatoids are common in parts of the intrusion. They have similar overall mineralogy as the layered rocks: dominantly b) Complex structures Wavy layers

Solitary layers Channel-like structures

20 cm

10 cm

c) Pegmatoid layers

10 cm

Regular layers

d) Miarolitic cavity

Epidote

5 mm

Fig. 4. a) Cross-cutting relationships with a channel-like structure, containing coarser-grained rocks, seen in a horizontal surface. Surfaces have been emphasised by red dashed lines. b) The range of petrographic structures within a single horizontal outcrop can be very large. Here wavy layering (left) abuts channel-like cross-cutting structures. Further to the right solitary layers transition into regular layering with a repeat of 30–40 mm. c) Pegmatoid layers alternate with fine-grained layers in this outcrop. Some of the pegmatoid layers have epidote, sometimes in miarolitic cavities. d) A well-developed miarolitic cavity in a pegmatoid layer with facetted crystals of plagioclase (white), hornblende (black) and epidote (green).

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plagioclase and amphibole, but without pyroxene. Amphibole crystals form idiomorphic prisms up to 80 mm long; some are hollow, suggesting rapid growth. In some places they occur as patches cutting layering (Fig. 3c). They can also occur as larger (metre-scale) irregular pods and segregations, particularly in the road cut to the west. Layered pegmatoids are quite common: pegmatoid layers are typically 5–10 cm wide, with intervening layers of gabbro with normal grain-size (Fig. 4c). Some layered and massive pegmatoids have miarolitic cavities lined with amphibole, plagioclase and epidote crystals up to several millimetres long (Fig. 4d). 2.2. Thin section petrography The regular layering is not as well-defined when viewed in thin section (Fig. 6a, b) as it is in outcrop (Fig. 3). This is partly because some amphibole crystals are so large (up to 5 mm) that occupy the entire width of a ‘dark’ layer and are discontinuous along the layers. Plagioclase is the most abundant phase in most samples. The degree of sericite alteration is variable, but low in most examined sections and all of the chemically analysed samples. The texture of plagioclase is very variable between and within samples and will be described first from the layered samples (Fig. 6c-f). Plagioclase crystals enclosed within the amphibole macrocrysts are significantly smaller (~0.5 mm) than those in the pale layers and typically make up 25% of the crystals (Figs. 6e-7f). There is no preferential alignment of crystals. Some crystals are rounded, but many are euhedral, with aspect ratios around 1:5:5. Larger crystals tend to be more equant. Most crystals have little zoning with typical compositions of about An56 from measurements of extinction angles. Plagioclase crystals have albite, Carlsbad and pericline twinning, but the latter does not appear to be related to deformation. Plagioclase from the plagioclase-rich layers has a much wider range of sizes, 0.1–3 mm, than that enclosed in the amphibole crystals. In some sections plagioclase crystals have a weak alignment which is parallel to phase layering. There are no bent or broken crystals. Plagioclase is broadly similar in the unlayered rocks to that in the layered rocks, except that the crystals tend to be larger and more uniform in size in the unlayered rocks. Mafic minerals are amphibole, pyroxene and olivine, and are similar in both layered and unlayered gabbros. Amphibole is generally the second most abundant phase, with modes of 20 to 50% (Fig. 6). It is pleochroic from beige to mid-brown to mid-green-brown. Most amphibole occurs as large (b5 mm) poikilitic crystals enclosing plagioclase, pyroxene and oxides. Clinopyroxene occurs almost exclusively within amphibole. Ragged crystals of clinopyroxene are commonly in optical continuity with the surrounding amphibole, suggesting that some of the amphibole has grown at the expense of clinopyroxene (Fig. 6g, h). Rare rounded olivine crystals (b0.5 mm) are present along contacts with plagioclase and commonly show fine-grained orthopyroxenemagnetite symplectite rims. The rock is completed by Ti-magnetite, together with minor apatite and zircon.

Bas et al. (1986) all compositions fall into the basalt field except for one which has a sufficiently low SiO2 content that it falls in the picrobasalt field. Layered samples seem to be systematically poorer in silica than massive samples for equivalent alkali content. MgO is very variable, but there is no consistent variation between layered and unlayered samples (Fig. 7b). H2O contents (as estimated from LOI) vary from 1.0 to 3.5% (Fig. 7c), and again there is no correlation with SiO2 content or presence of fabric. In contrast, Zr is well correlated with SiO2 (Fig. 7d). Trace element abundances are quite variable (Fig. 7e, f). All three massive samples have generally similar straight REE spectra, except for their La and Ce contents, without significant Eu anomalies. The layered samples have a wider range of compositions. Sample ILL0149 resembles the unlayered samples. ILL0863 has only about half the REE of the other samples but is otherwise similar. It also has the lowest SiO2 content and the least radiogenic Sr isotopic composition (see below). Sample ILL09150 has a more pronounced slope and is the only sample with a significant positive Eu anomaly. It is from a plagioclase-rich layer. Trace element abundances normalised to primitive mantle {(Sun and McDonough, 1989) have significant deficiencies in Nb and Ta, and positive anomalies in K, Pb and Sr.

2.4. Isotope geochemistry Sr and Nd isotopic compositions were measured on a Finnigan MAT 262 thermal ionization mass spectrometer (TIMS) with variable multicollector detector and RPQ at the Centro de Instrumentación Científica, University of Granada (Spain), following the procedures described in Montero and Bea (1998). Normalization value for 87Sr/86Sr was 88Sr/86Sr = 8.375209 and the reproducibility under successive determinations of the NBS-987 dissolved standard was better than 0.0007% (2σ). For the Nd determinations, the normalization value for 143Nd/144Nd was 146Nd/144Nd = 0.7219, with a precision better than 0.0016% (2σ) calculated under successive measures of the WSE power standard. The reproducibility of successive measurement of the La Jolla dissolution standard was better than 0.0014% (2σ). Initial isotopic ratios were calculated for an age of 115 Ma (Morata et al., 2010). Sr and Nd initial isotopic ratio are very variable (Fig. 8a). Samples ILL09149 and ILLo9150 were taken from the same outcrop 2 m apart and their dispersion illustrates the geological and analytical variability of the data. The three samples from the eastern part of the intrusion are very similar to each other, but very different from those in the western part. Layered samples seem to have lower Sr initial ratios than the unlayered samples, but the same contrast is not seen in the εNd values. There is a correlation between the Sr isotopic composition and SiO2 content (Fig. 8b), but not between with Nd isotopic composition and SiO2 (not shown). There is no correlation between Sr initial isotope ratios and Eu anomalies (Fig. 8c).

2.3. Whole-rock geochemistry 2.5. Mineral chemistry The whole-rock major composition of six samples were measured by inductively coupled plasma atomic emission spectrometry (ICPAES, Perkin-Elmer P-430) at the Departamento de Geología, Universidad de Chile, Chile (major elements) using USGS BCR-1 (basalts) and G-2 (granite) as standards. Trace element determinations were done by ICP-MS using Rh as an internal standard. Samples were dissolved in HNO3 + HF at a pressure of 180p.s.i. in a microwave digestor, dried and redissolved in 4% HNO3. Precision, estimated from the analyses of 10 replicates of one sample, was better than ±2% and ± 5% for analyte concentrations of 50 and 5 ppm respectively. There is a wide range in whole-rock chemical compositions (Table 1). Silica varies from 41 to 50%; total alkalis (K2O + Na2O) from 1.5 to 5% (Fig. 7a). Using the classification of Le

Mineral compositions were analysed using a scanning electron microscope – a Zeiss Sigma 300 VP at IOS Geoservices, Canada. Analytical data are in electronic appendix 1. Cores of plagioclase crystals were very calcic with compositions from An86 to An91, whereas the rims varied from An44 to An63 (Fig. 9). Most pyroxene crystals that were analysed were enclosed by amphibole and had a narrow range in composition with a mean composition of Wo44En44Fs12 in the field of augite and a range in Mg number from 0.77 to 0.82. Rare pyroxene primocrysts had similar compositions. Amphiboles were classified using the scheme of Hawthorne et al. (2012) and the program of Locock (2014): All analyses fall in the field of magnesio-hornblende.

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a) Layers - unpolarised light

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b) Layers - cross-polarised illumination

Dark layer D 5 mm c) Pale layer- unpolarised light

d) Pale layer - cross-polarised illumination

1 mm e) Dark layer - unpolarised light

f) Dark layer- cross-polarised illumination

1 mm g) Relict pyroxene

1 mm

h)

Plag

Cpx Amph

Fig. 6. a) Rhythmically layered gabbro from the outcrop in Fig. 3A in unpolarised light and b) linearly cross-polarised illumination. Layering is most clearly defined by the amphibole oikocrysts and the centre of the layers are shown by red dashed lines. c), d) Magnification of part of a pale layer. e), f) Magnification of part of a dark layer, from the bottom-left of panel A. g) Layered gabbro with plagioclase (plag), poikilitic amphibole (amph) and relic clinopyroxene (cpx) in plane polarised light and h) in linearly cross-polarised illumination.

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6

12

a) An56

5

b)

10

ILL09149

ILL09150 8 ILL0863

ILL09148

MgO

Na2O+K2O

ILL09146

Layered samples

4

ILL09147 Massive samples

3

ILL09149

2

ILL09147

6

4 ILL09146

ILL0863 1

ILL09148

2 Picrobasalt

ILL09150 An56

Basalt

0

0 40

42

44

46

48

50

SiO2

54

52

40

42

44

46

60

4

c)

48

50

52

54

SiO2 d)

ILL09150

ILL09147

50 ILL09149

ILL09146 40

Zr (ppm)

H 2O

3

ILL09149 ILL09148 2

ILL09148 ILL09150 30

ILL0863

ILL09147 20

ILL0863 1

ILL09146 10

0 40

42

44

SiO2 46

48

50

100

1000

e) 1

42

44

SiO2

46

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f)

ILL0863 ILL09150

Rock/Primitive Mantle

ILL09147

Rock/ Chondrites 10

0 40

52

100

10

ILL0863

1

1

Nd

Ce La

Pr

Sm Pm

Gd Eu

Tb

Dy

Yb

Er Ho

Tm

Lu

Cs

Rb Th Nb K Ce Pr P Zr Eu Dy Yb U Ta La Pb Sr Nd Sm Ba Ti Y Lu

Fig. 7. Whole-rock compositions of layered gabbros (open symbols), unlayered gabbros (filled symbols). a) Total alkalis versus SiO2. The classification of Le Bas et al. (1986) is shown by dashed lines and stoichiometric plagioclase of composition An56 is indicated by an asterisk. b) MgO versus SiO2. c) H2O versus SiO2. d) Zr versus SiO2. e) Chondrite normalised REE content {(Sun and McDonough, 1989). f). ‘Spidergram’ diagram. Trace element abundances normalised using primitive mantle abundances (Sun and McDonough, 1989).

3. Discussion 3.1. Development of the Cortaderas magmas It is not the purpose of this paper to discuss the origin of mafic magmas in the Andes as this has been treated in detail in many other papers (e.g. Hickey et al., 1986; Parada et al., 1999). We are concerned here

more with the local context: What can outcrop structures, mineral chemistry, whole-rock major element, trace element and isotopic compositional variations tell about how the intrusion formed and how it relates to the rest of the Illapel Plutonic suite? The compositions of the plagioclase cores are much more calcic than most gabbros reflecting the low abundance of sodium in these rocks. In contrast pyroxene and amphibole compositions resemble that of other

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Table 1 Cortaderas analyses. UTM coordinates are in Zone 19J Sample

ILL08063

ILL09146

ILL09147

ILL09148

ILL09149

ILL09150

Description UTM Easting UTM Northing SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total Ba Rb Th Nb Sr Zr Y Cr V Ni Co Sc Pb Cu Zn Li Cs Be Ga Mo Sn Tl Ta Hf U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu* (87Sr/86Sr)o eSr (143Nd/144Nd)o eNd Age (Ma)

Layered gabbro 290,467 6,519,430 41.3 1.14 17 8.05 6.88 0.18 8.03 13.3 1.27 0.25 0.01 1.06 98.47 72.7 4 0.5 0.7 454.9 25 7.8 99.2 671.1 32.1 55.9 58.7 1.9 569.3 68.5 3.5 0.2 0.2 18.4 3.3 0.6 0 0.1 0.5 0.1 2.81 6.85 1.01 5.13 1.49 0.54 1.57 0.25 1.45 0.29 0.75 0.11 0.67 0.1 0.35 0.70335 −15.8 0.51271 4.2 115.7

Massive gabbro fine grained 290,516 6,519,634 50.2 0.8 20.3 3.39 5.72 0.17 4.05 10 2.92 0.92 0.16 1.18 99.81 225.6 27.8 1.5 1.3 438.9 48 15.4 17 275.7 10.6 27.4 27.8 1.7 51.2 52.9 7.2 2.5 0.4 18.7 0.2 0.6 0.1 0.1 0.8 0.3 7.19 16.82 2.46 10.85 2.75 0.98 2.79 0.45 2.66 0.59 1.61 0.24 1.51 0.23 0.35 0.70361 −8.3 0.51276 5.3 115.7

Massive gabbro, medium grain 290,483 6,519,571 49.6 0.6 17.2 1.94 6.2 0.2 7.61 11.66 2.32 0.97 0.19 1.42 99.91 217.2 35.8 1.5 1.1 390.6 48 15 174.3 258.9 51.4 34.6 38.2 1.3 57.8 51.2 7.1 2 0.6 15.4 0.3 0.5 0.1 0.1 0.8 0.3 13.78 27.49 2.96 13.7 3.52 1.08 3.49 0.51 2.9 0.59 1.7 0.24 1.42 0.21 0.31 0.70367 −5.6 0.51271 4.4 115.7

Layered gabbro 290,612 6,519,537 44.5 0.68 25.6 4.15 3.52 0.08 2.74 12.9 2.76 0.58 0.12 1.94 99.57 177.4 17.4 0.6 1.8 672.7 33 17 11.5 235.5 8.4 17.2 21.2 1.7 231.8 29.5 4.4 0.3 0.6 20.1 0.3 0.7 0 0.1 0.9 0.2 6.75 16.78 2.68 12.46 3.21 1.03 3.42 0.54 3.19 0.64 1.74 0.26 1.57 0.23 0.31 0.70343 −13.4 0.51279 5.9 115.7

Massive gabbro 290,776 6,519,683 46.4 0.98 12.1 5.62 7.76 0.24 10.11 12.01 1.54 0.8 0.13 2.19 99.88 146.2 22.8 1 0.9 231 44 18.4 19.1 472.1 20.3 48.2 80.1 1.7 77.2 76.6 8.6 0.3 0.3 14.1 0.4 1.3 0 0.1 1.2 0.4 5.61 13.47 2.09 10.41 3.06 0.94 3.36 0.55 3.37 0.73 1.86 0.28 1.76 0.25 0.29 0.70349 −7.7 0.5128 6.1 115.7

Layered gabbro 290,776 6,519,683 46.8 0.47 25.6 1.95 3.6 0.07 2.23 10.67 3.45 1.47 0.11 3.47 99.89 396.9 50.8 1 0.9 723.6 36 8.1 11.1 169.8 7.9 13.3 11.5 1.8 67.5 22.3 15.3 2.5 0.4 18.7 0.2 0.3 0.1 0.1 0.6 0.4 5.44 11.32 1.58 7.08 1.64 0.7 1.57 0.25 1.48 0.31 0.8 0.13 0.81 0.11 0.43 0.70344 −10.2 0.51278 5.7 115.7

gabbros. These magmas may have become depleted in sodium by exchange with felsic magmas: Sodium would have been partitioned into alkali feldspars in the felsic magmas and removed from the mafic magmas. Other evidence of mixing will be discussed below. We will now consider the composition of the Cortaderas samples to determine if their variability was produced within the volume currently occupied by the intrusion from an originally homogeneous magma, or if multiple external sources were required. Ideally, the action of single, well-understood petrological processes — such as fractional crystallisation — should produce geochemical variations that can be easily interpreted to reveal the underlying parameters of the process. The

weak correlation between total alkalis and silica (Fig. 7a) could perhaps be partly explained by an accumulation of plagioclase with a composition close to An56, which is the typical crystal core composition found in these gabbros. However, this simple explanation is not supported by the REE data (Fig. 7e): only one sample has the significant positive Eu anomaly expected with such a process, and it lies off the correlation line. Dispersion of points on plots of MgO versus silica (Fig. 7b), and Zr versus silica (Fig. 7d) are also not consistent with a model of only plagioclase accumulation. Variations in Sr and Nd initial isotopic compositions show that there are many different magmatic sources for this small intrusion (Fig. 8a),

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0.7037

a) ILL09147 0.7036

c) ILL09147

ILL09146

ILL09146

ILL09146

Massive samples ILL09149

0.7035

0.7034

Layered samples

ILL0863

0.7033 3.0

3.5

4.0

4.5

5.0

ILL09149

ILL09149

ILL09150

8

( 7Sr/86Sr)i

ILL09147

b)

ILL09148

ILL09150

ILL09150

ILL09148

ILL09148 ILL0863

5.5

40

6.0

ILL0863

42

44

46

48

50

0.8

SiO2

Nd

1.0

1.2

1.4

Eu/Eu*

Fig. 8. Initial Sr and Nd isotope geochemistry of gabbros from the Cortaderas Intrusion for an age of 115 Ma. Filled symbols = massive samples; Open symbols = layered samples. a) Sr initial ratio versus epsilon Nd. b) Sr initial ratio versus SiO2.c) Sr initial ratio versus Eu/Eu*.

but can these different magma sources be related by a simple model? The correlation between Sr isotopic compositions and SiO2 (Fig. 8b) could be explained by the mixing of two components with approximately equal amounts of Sr, such as a gabbroic magma and a mass of plagioclase crystals. However, in such a case we would expect to see a correlation between Eu anomalies (from the added plagioclase) and Sr isotopic compositions, which we do not (Fig. 8c). If a third component is postulated then it would have to have had relatively little effect on the Sr-SiO2 correlation. A more likely story is that the Cortaderas intrusion was produced by multiple, independent injections of magma, each with isotopically different sources.

Plagioclase compositions (An) 0.40

0.50

0.60

0.70

0.80

0.90

1.00

Pyroxene compositions Hd

Di

3.2. Origins of layering in igneous rocks

Fs

En

Ca-amphibole 1.0

A* sum: A(Li + Na + K + 2Ca + 2Pb)

Edenite

Pargasite

Sadanagaite

Magnesio-hornblende

Tschermakite

0.5

Tremolite 0.0 0.0

The isotopic compositions of the Cortaderas samples are similar to those of the intermediate and acidic parts of the Illapel Plutonic suite (Parada et al., 1999) and do not resemble pristine mantle values. This suggests that the magmas that fed the Cortaderas intrusion were variably contaminated by crustal material, such as that which was melted to produce the Illapel Plutonic suite. The hydrous nature of the source magmas (Fig. 7c) is confirmed by the abundance of amphibole in both layered and unlayered rocks. The presence of late pegmatoids indicates that the water content of the magma increased during solidification (Fig. 3c). The presence of epidote and miarolitic cavities (Fig. 4d) shows that final solidification must have occurred at relatively low temperatures and pressures where water vapour could exist as a separate phase. The presence of many isotopically unrelated magmas in the Cortaderas intrusion accords with a model of the intrusion as an open magma conduit. The similarity of isotopic compositions of the Cortaderas magmas to felsic magmas of the Illapel Plutonic suite suggests that there may have been a link: Perhaps mantle-derived mafic magmas mixed with felsic magmas produced by crustal fusion.

0.5 1.0 1.5 C* sum: C(Al+Fe 3++Mn3++Cr+V+Sc+2Ti+2Zr) -WO-CLi

Fig. 9. Plagioclase, pyroxene and amphibole compositions.

2.0

Layering in igneous rocks may be produced by dynamic or nondynamic processes (Cawthorn, 1996; Charlier et al., 2015; Wager and Brown, 1968). The most commonly invoked dynamic processes involve accumulation of crystals from episodic currents in a magma chamber. Such currents can arise from within the chamber in response to thermal or crystallisation effects (e.g. Couch et al., 2001), or can be imposed from outside by recharge events. Crystal accumulations can form by sinking or floatation, depending on density differences. As such, they broadly resemble sedimentary processes with clearly notable differences due to viscosity and density contrasts. It should be possible to recognise these processes by variations in mineral proportions (if more than one mineral is present) and crystal size distributions: denser and larger crystals should be more abundant close to the accumulation surface (Higgins, 2002). Recharge events may produce erosional surfaces, compositional variations and/or ‘chilled zones’ with an abrupt change in mean crystal size. These criteria are not always easy to apply in practice as they may have been obscured by subsequent textural equilibration (Higgins, 2002). This type of dynamic layering can only be observed where crystals accumulate on shallowly inclined walls or the floor, or the roof for floatation cumulates. It also needs certain fluid properties that allow differential movement of crystals and magma. Dynamic magmatic processes are controlled by inertial and viscous forces and the Reynold's number (Re) is the ratio of these parameters. Glazner (2014) has pointed out that processes such as erosion, which

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depend on inertia, cannot occur in a low Reynold's number environment, so it is instructive to calculate this parameter for the Cortaderas magma. The Reynold's number (Re) is defined as Re = (ρlLU)/η, where, ρl = liquid density; L = length scale; U = velocity; and η = dynamic viscosity. The liquid density is taken to be 2700 km/m3. The length scale is that of the crystals which must be dislodged from the framework by inertial forces: a value of 5 mm is chosen. The dynamic viscosity is taken to be 50 Pa.s at 1000C (Giordano et al., 2008). For a velocity of 1 m/s Re = 0.27 and the magma flow will be laminar. Laminar flow will start to transition to turbulent flow at about Re = 2000. This would necessitate a magma velocity of about 7 km/s, which is unreasonable. Hence, flow is laminar at the crystal scale and erosion of existing crystal frameworks by turbulent forces cannot occur. The steep orientations of layering in the Cortaderas intrusion and the physical properties of the magma behoove us to consider non-dynamic explanations for layering. Layering in rocks can also be produced by non-dynamic processes generally termed geochemical self-organisation (Dewers and Ortoleva, 1990; L'Heureux, 2013; Ortoleva, 1994): One of the best known, and simplest, examples is the rhythmic ‘Inch-scale’ layering of the Stillwater intrusion (e.g. Boudreau, 1995). Non-dynamic layers can originate early in solidification by kinetic oscillations in crystal nucleation and growth (kinetic mechanism), or later in partly-crystallised rocks by textural equilibration (equilibrium mechanism) (Boudreau, 1995; Boudreau, 2011; Stern, 1954). A fundamental difference between these two mechanisms is the temperature: Kinetic processes occur at significant undercooling, whereas equilibration happens close to the mineral saturation temperature. It is possible that both explanations are correct in some natural situations: for example, equilibration may later enhance kinetically produced layers. Kinetic mechanism layering may develop in response to ‘double diffusive’ gradients of temperature and composition in a magma (McBirney and Noyes, 1979). Although such layering can develop in a stagnant, low-crystallinity magma, this situation is unlikely to be encountered in natural situations and an initial, homogeneous crystal framework commonly usually proposed (e.g. Ballenegger, 2016). Layering develops by crystallisation of a second mineral phase within this framework. Initially, the system undercools until this phase nucleates. It then grows fast, depleting the adjacent zone of the magma in this component, inhibiting nucleation in that region. Nucleation will then step out to an undepleted region away from the first layer. Repetition of this process will produce a layering of the second phase within a homogeneous framework. Equilibration mechanism layering develops in igneous rocks at low degrees of undercooling, that is just below the saturation temperature of a phase. Some fluid component, such as a silicate liquid, is necessary to facilitate movement of components between crystals. In this situation energy is dissipated by minimisation of the surface energy, rather than overall crystallisation of phases. Some surface energy can be lost by the simultaneous solution of small crystals and the growth of larger crystals, a process known as coarsening (see review by Higgins, 2011). Additional energy can be lost by the minimisation of the area of heterogeneous crystal contacts by the development of mineralogically more homogeneous layers. The balance of layer development versus coarsening is controlled by the relative surface energies of like (e.g. plagioclaseplagioclase) and unlike (e.g. plagioclase-amphibole) mineral interfaces. The surface energy of phase contacts is related to the mismatch between the crystal lattices. If the energy of the unlike-mineral interface is much greater than that of the like-mineral interface, then more energy will be dissipated by layer formation as compared to simple coarsening. Layering can also be promoted by a template – that is an existing planar discontinuity, such as the magma-mush interface or a mineral layer. Hence, equilibration can reinforce existing layers produced by other processes. There is also a kinetic component to the development of layering by equilibration. Thin layers can develop faster as they involve the least

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amount of transport, but they also yield the least reduction in total surface energy. Thicker layers take more time to develop, all other factors being equal, but are more effective in energy reduction. Of course, transport speed is strongly controlled by the porosity and permeability of the framework. Boudreau (2011) has pointed out that equilibration layering is quite rare in small intrusions, as the system generally cools too rapidly for equilibration to occur. Hence, minerals do not have enough time at temperatures close to their liquidus for equilibration to occur. There is, however, another possibility for equilibration that is well illustrated in metamorphic rocks. There, some minerals form by reactions from other minerals. Coarsening and layer formation are quite common under these circumstances (e.g. Bowes and Park, 1966; Thompson, 1975). A similar process may also occur in igneous rocks during conversion of one phase to a more thermodynamically more stable phase: an example that may be pertinent here is the conversion of pyroxene to amphibole. Sedimentary rocks can develop planar structures during lithification and similar processes may occur in igneous rocks. Modern use of the term Liesegang banding refers to rhythmic variations in the amount of cementing phases in some sedimentary rocks (Fu et al., 1994; Stern, 1954). Liesegang banding may cross-cut original sedimentary features with simple, parallel structures or complex banding (Fig. 12d). Discontinuities in the banding may represent the interaction of diffusion fronts on either side of a discontinuity in network permeability or the effects of multiple diffusion fronts that rework existing layering. Models involving kinetic processes (supersaturation/nucleation cycles) or equilibration have been proposed. In either case variations in network permeability will affect the development of Liesegang banding via variations in diffusive transport. Glazner (2014) has suggested that similar processes might have produced complex layering in granitoid intrusions: we suggest that such an explanation may also extend to some mafic intrusions. Another model has been developed for the generation of layering during crystallisation of phases from aqueous fluids in an overpressured (stressed) environment (Kelka et al., 2017). Such layering can be modelled by cnoidal waves, in which the layer spacing is controlled by intrinsic parameters: overpressure and permeability. Such a model can produce regular, closely-spaced waves as well as isolated, solitary waves. It is possible that a similar mechanism operates in magmatic systems, as overpressure and permeability can vary in the conduit of an open system. 3.3. Layered rocks in the Punta Falcone and Stillwater intrusions We have shown that the Cortaderas Intrusion has many special petrographic features, but these are not all unique: one other intrusion has so many similarities that it deserves special mention at this point. The Punta Falcone intrusion is a small (200 by 600 m), predominately mafic intrusion that lies within the Cretaceous Sardinia-Corsica Batholith (Ballenegger, 2016; Hauser and Bussy, 2014; Tommasini and Poli, 1992). The Punta Falcone intrusion (PFI) is well exposed along a 150-m coastal section, but outcrops inland are sparse and small (Ballenegger, 2016). Twelve mafic units have been distinguished on the basis of their field characteristics. Contacts between the different units are commonly sub-vertical, as are contacts with the host granites, and the intrusion has not been tilted significantly since emplacement. Contacts between units may be sharp or gradational and hybridisation has occurred between the mafic and felsic rocks. The short time span between the emplacements of mafic and felsic rocks is confirmed by radiometric dating. Ballenegger (2016) interprets the intrusion as having formed from 6 different magmatic pulses, each of which was emplaced within an earlier magmatic pulse. The mafic rocks are hydrous gabbros, dominated by plagioclase and amphibole. Some of the mafic units have regular, fine-scale layering defined by alterations of plagioclase and amphibole/clinopyroxene rich

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bands (25–40 mm layer spacing; Fig. 10a). The ‘strength’ of the layer definition is determined by the degree of mineralogical contrast within the bands. Layering may be truncated by layered or unlayered rocks (Fig. 10b) but channel structures similar to those seen in the Cortaderas intrusion (Fig. 4a) have not been observed. In some areas regular finescale layering is cut by irregularly spaced amphibole-rich bands (Fig. 10c). Elsewhere there are solitary plagioclase-rich bands up to 10 mm thick (Fig. 10d). In some places, these are parallel to fine-scale layering which may be well or poorly defined. The grain-size may vary across these bands, with the development of pegmatoid patches. The mineralogy of the regularly layered rocks is dominated by plagioclase, clinopyroxene and amphibole (Ballenegger, 2016). Plagioclase has cores with An80–90 and rims with An40–60, pyroxene has a Mg number of 0.70–0.78 and amphibole is largely Mg-hornblende. All these compositions are similar to those of the Cortaderas intrusion, especially the anomalously calcic plagioclase. In the mafic layers, plagioclase crystals 0.1–1 mm long are enclosed by poikilitic clinopyroxene and/or amphibole crystals up to 10 mm long. In some areas, clinopyroxene has narrow amphibole rims where in contact with other phases, but elsewhere amphibole replaced clinopyroxene completely. Plagioclase crystals in the pale layers are somewhat larger than in the dark layers, typically 1–2 mm long. Quartz also occurs as a sparse, late phase. There does not appear to be a significant fabric in these rocks. Overall, petrology of these layers resembles those seen in the Cortaderas intrusion. The source magmas of the PFI appear to have many isotopic similarities with those of the Cortaderas intrusion. The isotopic composition of different units in the PFI tells a broadly similar story to that of Cortaderas, of multiple magma sources (Ballenegger, 2016). Although both Nd and Sr isotopic compositions are much more radiogenic, they both have a significant range in compositions for the different units of the intrusion. And again, both the mafic rocks of PFI and their host granites a) Quasi-regular planar layering

20 cm

c) Two intersecting layering orientations

have broadly similar isotopic compositions. The presence of epidote and miarolitic cavities in the PFI again suggest that final solidification of a hydrous magma occurred high in the crust. Mineral and isotopic compositions suggest that original mafic magma interacted strongly with the felsic magmas of the batholith. The inch-scale layering in the Stillwater intrusion (Boudreau, 1987; Hess and Smith, 1960) superficially resembles regular layering in the Cortaderas and Punta Falcone intrusions. In this unit, regularly spaced layers rich in pyroxene and plagioclase alternate with spacing of about 24 mm. Viewed from a distance it is clear that the pyroxene-rich layers are doublets, but this is not evident in thin section. The layers are currently vertical and there is no evidence that this is not close to their original orientation. However, the grain-size of the minerals in this unit (up to 10 mm) is larger than that observed in the Cortaderas and Punta Falcone layered rocks. In addition, the mafic phase is pyroxene, rather than amphibole (Boudreau, 1995). Similar rhythmically layered rocks occur in many other mafic intrusions, but the original orientation is shallow or unclear.

3.4. Development of layering in the Cortaderas intrusion The limited amount of exposure means that it is not possible to construct a definitive, overall petrologic history of the intrusion. Instead, we will use field relationships and chemical analyses to show how parts of the intrusion may have formed and extend these ideas to an overall model. The unlayered and layered parts of the intrusion have similar chemical, mineral and isotopic compositions. As the unlayered rocks are a minor component, we will concentrate on the development of the layered rocks. The vertical to steep orientation of most of the layering observed in the field (Fig. 2) indicates that settling or floatation of crystals due to density differences cannot have created the layering. b) Truncated layering

10 cm

d) Thin plagioclase-rich and pegmatoid layers

Plagioclase

Pegmatoid

10 cm

10 cm

Fig. 10. Layered structures in the Punta Falcone intrusion, Sardinia. The mineralogy is dominated by plagioclase (pale phase) and amphibole (dark phase). a) Repetitive, parallel layers are defined by variations in plagioclase content. Some layers divide and combine along strike. b) Regular layering can be truncated by other layered rocks. c) In some parts of the outcrop there are late amphibole-rich layers that cut earlier regular layering. d) Narrow plagioclase-rich undulate across some parts of the outcrop. The grain-size may vary markedly across the plagioclase-rich bands and be pegmatitic in places.

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Plagioclase in A

Kinetic growth B

Pyroxene in C

25% crystallised

3.4.1. Regular layering The 3-m section shown in Fig. 3a shows some of the clearest evidence for layering by geochemical self-organisation. The layers are almost perfectly rhythmic and parallel. The spacing of the layers varies regularly from one side of the outcrop to the other. A simple physicochemical model best explains all these features. The process of solidification of the magma can be interpreted from the final product, which is the rock, supplemented by intermediate textures preserved within amphibole macrocrysts. We propose that the process started with the emplacement of a batch of magma into cooler country rocks, which then stagnated and started to crystallise plagioclase, followed by minor clinopyroxene, olivine and oxides. The system continued to cool until amphibole started to crystallise. Amphibole crystals grew rapidly in layers, displacing the silicate liquid and preserving the texture of the early phases as chadocrysts. Solidification terminated with the crystallisation of residual liquid, largely as rims on plagioclase and amphibole. Chadacrysts in amphibole can tell us much about the state of the magma just before the formation of the mafic-rich layers. Plagioclase crystals have a complete range in size from (sparse) microlites to 1 mm long. Crystals are not generally euhedral and some occur in clusters. The lack of foliation of plagioclase chadacrysts indicates that the magma had a low content of this phase when emplaced. However, the irregular shapes of the plagioclase crystals do suggest that there was some disruption before the texture was sealed by growth of amphibole. The plagioclase chadacrysts are evenly distributed throughout the macrocryst and have an overall abundance of about 25%. The presence of ragged pyroxene crystals in the centre of many amphibole macrocrysts shows that amphibole grew partly by conversion of earlier-formed pyroxene. However, the size of the amphibole crystals suggests that most of the amphibole crystallised directly from the silicate liquid. The large size of the amphibole crystals, and the paucity of small crystals suggest that that this phase has been coarsened (Higgins, 2011). This process is only significant in igneous systems if the magma remains at the crystallisation temperature for a protracted period. Aspects of the final stages of solidification can be determined from the textures of the felsic layers. These layers are very rich in plagioclase crystals, which are generally larger than those trapped in the amphibole macrocrysts (Fig. 6). Plagioclase must have continued to grow after crystallisation of amphibole; hence these layers must be the last part of the rock to solidify. Very small plagioclase crystals are present showing that this phase is not coarsened here. Some parts of the plagioclase-rich layers have signification foliation defined by a preferred orientation of the shape of the plagioclase crystals. Magmatic foliation can be produced by compaction or shear (Higgins, 1991). Compaction alone is not efficient at producing a significant fabric, unless combined with pressure-solution (Meurer and Boudreau, 1998). The sub-vertical orientation of the layers precludes compaction driven by gravity, but a type of filter-pressing produced by changes in overall stress are possible. Pressure-solution has not been observed but cannot be ruled out. Shear of crystals mushes is a much more efficient at producing significant foliations. The crystal mush is likely to be weak during solidification, especially parallel to the amphibole-rich layers which were cemented earlier. Hence, stress with a component parallel to the layers could produce shear in the plagioclase-rich layers. It may well be that both mechanisms are important here. A cooling and solidification model for the regularly layered rocks can now be proposed (Fig. 11). Crystallisation of plagioclase starts at point A, either in the conduit or in the current position. Rapid cooling increases the undercooling, leading to a kinetically controlled texture (Higgins, 2011). At point B pyroxene saturates and cooling continues

to point C, the amphibole peritectic. The rock is now about 20% crystallised. Here some pyroxene is converted to amphibole and more amphibole crystallises directly from the silicate liquid. The temperature is buffered from C to D by the conversion of pyroxene to amphibole and the release of latent heat by amphibole crystallisation. Amphibole equilibrates both by coarsening and by development of amphibole-rich layers. At point D most amphibole has crystallised and cooling recommences, increasing the size of plagioclase crystals until point E when the rock is completely solid. Foliation in the plagioclase-rich layers may have been produced by shear deformation of the magma between points D and E. The model proposed above requires a period of slow cooling, which is not generally encountered in a small intrusion. One possibility is that Cortaderas is the remains of a magmatic conduit, which was open to the surface where it fed lava flows. Repeated input of low-crystallinity magma may then have mimicked the thermal effects of a large intrusion. The significant variation in layer spacing over a distance of a few metres suggests that there was a strong gradient in thermal and chemical fields and hence that the contact with the feeder was very close. There is also another possibility which is compatible with this model: conversion of pyroxene to amphibole may have also helped the equilibration process. The abundance of amphibole in the layered rocks suggests that amphibolitisation of pyroxene may have played a role in the development of the layering. Mineralogical reorganisation of the rock expressed by the conversion of pyroxene to amphibole during solidification, a peritectic reaction called hydration crystallisation by Beard et al. (2004), could enable the system to approach textural equilibrium more readily that simple cotectic (saturation) crystallisation of a phase. An important aspect of the proposed model is that layering starts to develop at relatively low degrees of crystallinity, as shown by the abundance of plagioclase preserved within the amphibole oikocrysts. The amount of a phase necessary to make a rigid framework is also dependant on the shape of the crystals and their distribution (Jerram et al., 2003; Philpotts et al., 1999). The amount indicated here, 25%, is close to the lower limit. However, even if the network was not complete it may have given a certain rigidity to the magma that was sufficient to enable layer formation. Ballenegger (2016) has proposed that regular layering in the PFI was produced by kinetic processes based on a double diffusive model (McBirney and Noyes, 1979). She rejects equilibration-based models for layer development on the basis of dihedral angles: she considers that that these have impingement textures and hence show that equilibration has not occurred. However, the petrographic context of the

Schematic temperature

We must therefore consider models in which the geology and textures of these rocks result from non-dynamic processes.

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D

Coarsening and layer development

Amphibole peritectic Kinetic growth E

Schematic time Fig. 11. Proposed model for the development of regular layering with schematic time and temperature scales, following Higgins (2017). The position of the saturation and peritectic temperatures are also schematic. Reference points A to E are discussed in the text.

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observations must be considered. In the model proposed here for Cortaderas initial plagioclase growth is at high undercooling and this kinetically-controlled texture is preserved by the rapid growth of amphibole macrocrysts by coarsening. This situation could also be applied to the PFI. This does not mean that Ballenegger's kinetic model for layering is not correct, just that the dihedral angle between a plagioclase chadocryst and its host amphibole is not a good criterion to reject equilibration-based processes for layer development. It is quite possible that layering is produced by kinetic processes early in the solidification of both Cortaderas and PFI and this is amplified by later equilibration processes. The former process would mostly affect plagioclase crystals and may not be preserved in the final rock, whereas equilibration controls amphibole textures which could be well preserved. 3.4.2. Solitary layers One of the more enigmatic features in the Cortaderas intrusion is the isolated plagioclase-rich bands termed solitary layers. They can be parallel to regular layering or components of irregular layers. Structures petrographically similar to solitary layers have been observed in the PFI (Ballenegger, 2016). These features appear to have formed at a high temperature and are unrelated to currently visible fractures in the crystal framework or rock. However, it is not impossible that some bands developed from earlier discontinuities such as fractures or shear-zones, and this will be discussed below. If the solitary layers are just an extreme form of geochemical selforganisation then it may be instructive to think of these structures as chemical solitary waves or solitons. Solitons are single waves or wave packets that maintain their shape during propagation due to a cancellation of nonlinear and dispersive effects. The original idea was developed from observations of water waves (Drazin and Johnson, 1989), but has been extended to many other environments, including partially molten rocks (Scott and Stevenson, 1984). The first applications were for the transport of magma through the lithosphere, but it may also be applicable at a small scale. It is not easy to model these processes as we have so little information on the basic parameters during solidification of the rock. Solitary layers may also be a form of Liesegang banding (see below) or possibly a response to overpressure in a permeable system (Kelka et al., 2017). In the latter model cnoidal pressure waves could lead to local solution of amphibole resulting in narrow plagioclase-rich layers like those observed at Cortaderas and PFI. Ballenegger (2016) considered that solitary layers formed by percolation of evolved liquids through partially solidified rock and are perhaps related to hybridisation of mafic and felsic rocks. This origin may not be completely incompatible with the solitary wave theory proposed above. However, Ballenegger's modelling of layer development also produced irregularly spaced solitary layers, so their origin may not be very different from that of the regular layers. 3.4.3. Irregular layering and cross-cutting structures Although the regularly layered rock can be modelled by geochemical self-organisation, the simple action of this process in a homogeneous medium cannot produce the abrupt truncations of layers that are so commonly observed in the Cortaderas intrusion and the PFI. We will consider two models for the origin of these features: erosion by new influxes of magma and reactive flow in a crystal framework. The simplest explanation for these discordant surfaces is erosion of existing layered rocks by turbulent influx of a new magmatic pulse into the feeder dyke. The presence of such truncations confirms that the development of layering must have been very early and that the rock was not completely solidified at that time. Discordant surfaces are common in the Cortaderas intrusion: in some cases, they resemble sedimentary crossbedding structures (Fig. 4), elsewhere they resemble channel structures. However, the channels are vertical and hence cannot have developed by the same processes that produce channels in sedimentary rocks. The sides of some channels are at a large angle to

the centre, covering almost 150° in some cases. Estimates of magma properties and flow rates discussed earlier give a Reynolds number that is much too low for turbulent flow. Hence, this proposition must be rejected. Reactive flow can produce complex structures in sedimentary, metamorphic and igneous systems, some of which can be referred to as Liesegang banding (L'Heureux, 2013). An example of such complexity is shown in Fig. 12d for haematite-rich bands in sandstone. The models presented in Fig. 12a-c show the kind of structures that may be produced by the interaction between reactive flow and variations in the permeability of the crystal framework. Low permeability is taken to inhibit the production of layering as movement of species in the magmatic fluid (silicate melt) between the crystal framework is retarded. The discussion will be in terms of propagation of layering and coarsening waves. These models do not consider the possibility of multiple events, which may rework existing layers and lead to more complex structures. In Fig. 12a, part of a propagating wave of regular layering encounters a region of lower permeability. The wave is retarded in that area and lateral propagation of the unretarded wave produces a discontinuity that resembles the cross-cutting structure seen in Fig. 3c. In Fig. 12b, a propagating wave of regular layering encounters a discontinuous surface with lower permeability. The wave that propagates through the gap spreads out to make a structure with a geometry similar to that of a channel, except that the sides can have a much wider angle. Such a structure resembles those seen in Fig. 3d. Fig. 12c illustrates a more extreme form of the structure shown in Fig. 12b, but the gap is smaller and the zone of low permeability is much deeper. Here, the waves that propagate through the gap spread out to make a balloon-like structure similar to that seen in Fig. 5. 3.4.4. Fine-grained layers Fine-grained layers are a widespread but uncommon feature in the Cortaderas intrusion. Such surfaces in other intrusions are commonly interpreted as internal chill zones, where a new, macrocryst-free magma is cooled rapidly against an existing rigid crystal framework. However, if this were so, then we would expect to see fine-grained microdykes extending from the surface into the solidified material. These have not been observed, hence there may be another explanation: these surfaces may represent high-temperature mylonites formed late when the rock had little or no silicate liquid present. The emplacement of the Cortaderas intrusion along a major active fault means that the solidified part of the intrusion must have been repeatedly disrupted, and perhaps these structures are the vestiges of these shear-zones, now partly coarsened. These surfaces could also act as templates for the development of non-dynamic layering, in which case their fine-grained character may have been completely erased. 3.4.5. Development of pegmatoids and miarolitic cavities The presence of extensive pegmatoid layers and patches suggest that in some circumstances the volatile content of the magma in the conduit could increase significantly, reducing magma viscosity and hence increasing crystal growth rate. The presence of miarolitic cavities in some pegmatoid patches confirms the high volatile content of these regions (Fig. 6b) and suggests that the intrusion solidified at sufficiently low pressures for a volatile phase to exsolve. The presence of epidote shows that the process continued to low temperatures, when the surrounding rock was largely solidified. The pegmatoids have several different forms, suggesting different petrologic paths. The layered pegmatoids (Fig. 6a) may have formed by the same physicochemical process that created the regularly layered gabbro and possibly the solitary layers. However, in this case the initial volatile content of the magma may have been higher than elsewhere. During textural equilibration, crystallisation of layers increased the volatile content of the intervening zones so that pegmatoids could form and, in some cases, also a separate vapour phase that filled the miarolitic cavities. The patch pegmatoids (Fig. 3c) cut the layering and must have

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a) Crossbedding-like structures

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b) Channel-like structures

Arrows indicate the directions of progressive development of waves Low-permeability surface

c) Balloon structures

Low-permeability surface

d) Liesegang layering in sandstone

10 mm

Low-permeability zone

Fig. 12. a), b), c) Schematic model for the origin of macroscopic structures by processes similar to that which produces Liesegang banding. The process occurs within a crystal framework that pervades the whole region shown here. It is illustrated schematically by the position of coarsening waves and their interaction with zones of lower permeability. The arrows do not imply movement of magma, but only of wave development. a) A low-permeability surface retards part of the coarsening wave. The other part continues and encroaches along the rear of the discontinuity to create a structure that partly resembles crossbedding. b) Coarsening waves propagating through a wide gap in a low-permeability surface produce a structure that resembles a channel. There are no physical limits on the angle subtended by the sides of the ‘channel’. c) Propagation of coarsening waves through a narrow gap in a low-permeability zone produces a balloon-like structure. d) Liesegang banding in sandstone from Kanab, Utah, USA. Bands are produced by differing amounts of iron oxides in the cement.

developed later. It appears that the hydrous silicate liquid digested existing layers of plagioclase and amphibole at the same time as the crystallisation of the larger amphibole and plagioclase crystals of the pegmatoid took place. 3.5. Emplacement model We can now integrate these ideas into a simplified model for the magma formation, emplacement and development of the main features of the Cortaderas intrusion. Mafic magma originated in the upper mantle or lower crust. It rose until it encountered the base of the IPS batholith, where it stalled and interacted with the felsic magmas that would later form the plutons of the batholith. The mafic magma lost sodium and may have gained water. There may also have been isotopic exchange of Sr and Nd. The magma then continued to rise in an existing fault or discontinuity and partly solidified to form a relatively narrow gabbroic dyke. The orientation of the Cortaderas intrusion and the layering suggest that this earlier structure was oriented approximately NE-SW (Fig. 2), which is parallel to the nearby 100 km long Silla de Gobernador fault. The presence of miarolitic cavities in parts of the intrusion suggests that the rocks observed here crystallised close to the surface, perhaps at the 5 km depth suggested for the IPS (Varas et al., 2012). The dyke may have opened towards the surface and fed volcanic activity. Many different batches of gabbroic magma rose in this structure and partly solidified on the walls. The first batch was emplaced into cool rocks and solidified relatively rapidly to produce the massive gabbro in the north and west parts of the intrusion. It initially crystallised plagioclase, with minor pyroxene and olivine. Further cooling promoted crystallisation of amphibole, partly fed by solution of pyroxene.

Further batches were emplaced into warmed host rocks and hence cooled more slowly. It initially crystallised a framework of plagioclase, pyroxene and amphibole. The slow cooling enabled crystallisation under conditions close to equilibrium; hence the texture could coarsen and develop layering as a way of dissipating excess surface energy associated with smaller crystals. The complex arrangement of layers observed in parts of the Cortaderas intrusion may have been formed in response to variations in the permeability of the framework, in a similar way to the formation of Liesegang banding. Layering in small mafic intrusions is relatively rare, possibly because there is insufficient time for it to develop before solidification (Boudreau, 2011). The model proposed here circumvents this problem for this intrusion in two ways. 1) The intrusion is a fossil conduit, kept warm by many magma pulses. Such an emplacement model been proposed for other, much larger plutons (e.g. Glazner et al., 2004). 2) Equilibration layering can be promoted by peritectic crystallisation of amphibole. The layering process described shares aspects of modal igneous layering developed in magma chambers (Boudreau, 2011) and metamorphic differentiation, which occurs during sub-solidus changes in the mineralogy of a rock (Bowes and Park, 1966). 4. Conclusions The Cortaderas intrusion is the solidified remains of a conduit that conveyed hydrous mafic magma along the margin of a major felsic plutonic suite towards the surface. The isotopic, chemical and volatile composition of the magma changed many times, reflecting different sources. The petrological complexity seen in the field reflects these compositional changes, modulated by variations in flow and cooling

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rates of the magmas. The conduit seems to have followed an active fault whose synmagmatic movements may have further modified to the intrusion. The Punta Falcone intrusion in Sardinia has many similarities with Cortaderas, both in its structure, composition and tectonic setting (Ballenegger, 2016; Hauser and Bussy, 2014; Tommasini and Poli, 1992). This suggests that layering in small mafic intrusions may be common within a specific context: high-level conduits of hydrous mafic magmas associated with large granitic batholiths. Two examples are known, but there may be others that have been observed, but not studied, as they do not seem to contribute much to the overall history of the batholiths in which they seem to occur. The complexity of the rhythmic structures and the lack of a simple explanation may also have made publication difficult. Layering in these small intrusions can be simple or extremely complex. However, in both cases the dominant layering appears to form by non-dynamic processes during crystallisation of amphibole. This may build on earlier kinetic layering, which is now concealed. Layer development starts at about 20% crystallinity, which is lower than that usually considered necessary to form a rigid framework. The complex patterning in some parts of the Cortaderas intrusion has topological similarities with Liesegang banding and a similar origin is proposed. However, there remains a considerable challenge to model such processes and produce patterns similar to those seen in the field. There are many similarities between the vertical ‘trough bands’ seen here and sub-horizontal structures seen in other intrusions, such as the Skaergaard intrusion, which may necessitate the re-examination of the latter as the result of non-dynamical processes. Acknowledgements This research has been funded by the Chilean National Science Foundation (FONDECYT) Project 1080468 to DM and a ‘Discovery’ grant from NSERC (Canada) to MDH. MDH thanks the Université du Québec à Chicoutimi and the Departamento de Geología, Universidad de Chile for granting and hosting his sabbatical research leaves. Anne-Cecile Ballenegger (formerly Hauser) is thanked for sharing her knowledge of the Punta Falcone intrusion. This is a contribution to the CONICYT/ FONDAP project #15090013. Appendix A. Mineral compositions Supplementary data (mineral compositions) to this article can be found online at https://doi.org/10.1016/j.lithos.2019.05.008. References Arancibia, G., 2004. Mid-cretaceous crustal shortening: evidence from a regional-scale ductile shear zone in the Coastal Range of Central Chile (32° S). J. S. Am. Earth Sci. 17, 209–226. Ballenegger, A.-C., 2016. Fine-Scale Rhythmic Magmatic Layering by a Double Diffusion process: The Example of the Mid-crustal Punta Falcone Mafic Intrusion (Sardinia, Italy). PhD Thesis. Department of Geology. University of Lausanne, Lausanne, Switzerland. Le Bas, M.J., Maitre, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification of volcanic rocks based on the total Alkali-Silica Diagram. J. Petrol. 27, 745–750. Beard, J.S., Ragland, P.C., Rushmer, T., 2004. Hydration crystallization reactions between anhydrous minerals and hydrous melt to yield amphibole and biotite in igneous rocks: description and implications. J. Geol. 112, 617–621. Boudreau, A.E., 1987. Pattern forming during crystallisation and the formation of finescale layering. In: Parsons, I. (Ed.), Origins of Igneous Layering. D Reidel, Dordrecht, pp. 453–471. Boudreau, A.E., 1995. Crystal aging and the formation of fine-scale igneous layering. Mineral. Petrol. 54, 55–69. Boudreau, A., 2011. The evolution of texture and layering in layered intrusions. Int. Geol. Rev. 53, 330–353. Bowes, D.R., Park, R.G., 1966. Metamorphic segregation banding in the Loch Kerry Basite Sheet from the Lewisian of Gairloch, Ross-shire, Scotland. J. Petrol. 7, 306–334. Cawthorn, R.G., 1996. Layered Intrusions. Elsevier, Amsterdam; New York. Charlier, B., Namur, O., Latypov, R., Tegner, C., 2015. Layered Intrusions. Springer. Couch, S., Sparks, R.S.J., Carroll, M.R., 2001. Mineral disequilibrium in lavas explained by convective self-mixing in open magma chambers. Nature 411, 1037–1039.

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